Palaeogeography, Palaeoclimatology, Palaeoecology 337–338 (2012) 88–98

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Palaeogeography, Palaeoclimatology, Palaeoecology

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High-resolution isotopic record of C4 photosynthesis in a Miocene grassland

Jennifer M. Cotton a,⁎, Nathan D. Sheldon a, Caroline A.E. Strömberg b

a Department of Earth and Environmental Sciences, University of Michigan, Ann Arbor, MI 48109, USAb Department of Biology & Burke Museum of Natural History and Culture, University of Washington, Seattle, WA 98195, USA

⁎ Corresponding author.E-mail address: jcot@umich.edu (J.M. Cotton).

0031-0182/$ – see front matter © 2012 Elsevier B.V. Alldoi:10.1016/j.palaeo.2012.03.035

a b s t r a c t

a r t i c l e i n f o

Article history:Received 17 November 2011Received in revised form 23 March 2012Accepted 27 March 2012Available online 4 April 2012

Keywords:PaleosolsC4 photosynthesisPaleoclimateStable isotope geochemistry

The origin and spread of grasslands is one of the key evolutionary events in the Cenozoic, which ischaracterized by long-term cooling and drying climatic conditions. One way to examine the relationshipbetween vegetation and climate is to study the isotopic composition of organic carbon preserved in paleosols.Paleosols that preserve organic matter in the form of carbonaceous root traces provide direct evidence of theproportion of C3 to C4 biomass that grew in the soil, in contrast with pedogenic carbonate δ13C values, whichmay also reflect complicating factors including mixing from atmospheric δ13C in low productivityecosystems. A high temporal and spatial resolution reconstruction of past ecosystems was derived fromthirty-five paleosols in a thirty-four meter section of the Sixmile Creek Formation at Timber Hills, Montana(USA) that was deposited during the Miocene (10.2–8.9 Ma ago). Phytoliths were extracted from paleosolsamples to compare vegetation assemblages to inferences based on isotopic compositions, with both proxiesgiving similar results. Isotopic results from organic matter indicate both a small component of C4photosynthetic plants locally prior to their regional expansion to dominance in the late Miocene throughthe early Pleistocene, and large variation in the abundance of C4 plants (0–25%) in this ecosystem bothlaterally and on a 100 Kyr timescale. In contrast, pedogenic carbonate δ13C values from this site indicate ahigh proportion of C4 photosynthesis that is at odds with both phytolith and δ13Corg results, suggesting thatthe carbonate values are biased by diagenesis or diffusion of atmospheric CO2, and that a similar issue mayimpact previous paleovegetation reconstructions based on pedogenic carbonates. Quantitative reconstruc-tions of mean annual temperature and mean annual precipitation indicate little local variability through timeand that the fluctuations in C4 proportion were not climatically driven. Instead, the variable proportion ofC4 photosynthesis is best explained by ecosystem-scale variables such as succession, fluvial avulsion, and fireregime.

© 2012 Elsevier B.V. All rights reserved.

1. Introduction

Grasslands, which currently occupy 25% of Earth's surface (Shantz,1954), serve as an important example of the evolution of vegetationin response to both climate change and to the adaptations ofassociated herbivorous vertebrates (Janis et al., 2002). Based onmolecular clock ages, grasses first appeared as a component ofecosystems between 70 and 55 Ma ago, with the early grasses usingthe C3 photosynthetic pathway and occupying narrow ecologicalniches (Kellogg, 2001). Utilization of the C4 photosynthetic pathwaydeveloped more recently, likely during the Oligocene (Edwards et al.,2010). The C4 photosynthetic pathway employs PEP-carboxylase as aCO2 pump, increasing the efficiency of photosynthesis in times of lowCO2, as well as in arid environments where respiration is costly (Sage,2004). Since their origin, grasses using C4 photosynthesis have spreadto make up the majority of modern grasslands, and now account for

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18% of vegetation on Earth (Melillo et al., 1993; Tipple and Pagani,2007). Many studies record a global expansion of C4 grassesbeginning in the late Miocene and extending to the late Pliocene inmultiple sites around the world from 8 to 3 Ma (Cerling et al., 1997;Latorre et al., 1997; Fox and Koch, 2003; Behrensmeyer et al., 2007),though not all localities show a monotonic increase in abundance(Kingston et al., 1994; Edwards et al., 2010; Uno et al., 2011). Fewstudies have focused on the regional extent of C4 vegetation in NorthAmerica prior to this expansion. The goal of this study is toreconstruct C4 vegetation of an earlier Miocene grassland in Montanain an effort to understand the extent and variability of the vegetationand its dependence on climate on short (100 Kyr) time scales.

It is well known that the C4 photosynthetic pathway existed longbefore the global expansion in the late Miocene through Pliocene.Molecular phylogeny predicts that C4 photosynthesis originated inthe earliest Oligocene (Edwards et al., 2010) and that the pathwaylikely evolved in response to environmental change (Ehleringer et al.,1997; Sage, 2004). Both the oldest C4 phytolith (plant silica bodies;Strömberg, 2005) and oldest isotopic (Tipple and Pagani, 2007)evidence for C4 photosynthesis are dated to the early Miocene.

89J.M. Cotton et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 337–338 (2012) 88–98

Several recent studies corroborate this date of origin, with isotopicand paleobotanical evidence of the existence of C4 plants back as faras the Oligocene (Edwards et al., 2010; Strömberg and McInerney,2011; Urban et al., 2011; Fox et al., 2012). Fox and Koch (2003) andFox et al. (2012) also demonstrate small abundances of C4 vegetationin the Great Plains of the United States throughout the mid to lateMiocene. Thus, it is expected that some amount of C4 vegetationwould be present in the Miocene of Montana, but the extent of thisvegetation during the Miocene is currently unknown.

Due to the differences in isotopic fractionations between theprocesses of C3 and C4 (e.g., most modern grasses) photosynthesis,stable carbon isotopes can be used to estimate with some uncertaintythe amount of C4 biomass present in a particular environment. δ13Corgof modern C3 plant biomass ranges from−22‰ to −34‰ and has anaverage isotopic composition of −27±2.3‰ (V-PDB scale; O'Leary,1988; Diefendorf et al., 2010). Much of the variability in C3 δ13Corgcomposition is a function of moisture availability (Diefendorf et al.,2010; Kohn, 2010). For example, the lightest isotopic values (δ13Corgb

−30‰) are only found in tropical rainforest. Modern C4 biomass ismore enriched in 13C with a range of−9‰ to−16‰with an averageδ13Corg isotopic composition of−13±0.8‰ (Vogel, 1993).Mixed C3/C4ecosystems can be reconstructed using a two end-member mixingmodel (e.g., Fox and Koch, 2003, 2004). As an example, Sheldon (2006)used soil organic carbon δ13Corg preserved in Hawaiian paleosols toreconstruct a pure C3 ecosystem there for the past ~450 Ka, results thatwere supported by palynological data.

Previous studies of older time intervals have reconstructed pastvegetation by relying instead on pedogenic carbonates due to the factthat they are more commonly preserved than organic carbon (Quadeand Cerling, 1995; Latorre et al., 1997; Fox and Koch, 2003, 2004), andbecause the Δ13Corg-carb fractionation is constrained to +14–17‰(typically +15.5‰; Koch, 1998; Cerling, 1984), which allows for atwo-component mixing model based on the pedogenic carbonate endmembers. One complication of using pedogenic carbonates is thatthey can be affected by diffusion of isotopically heavy atmosphericCO2 (e.g., ~−6.0‰ for the Miocene atmosphere; Passey et al., 2009;Tipple et al., 2010) into the soil, which increases in areas of low soilproductivity. Thus, reconstructions using pedogenic carbonates havethe potential to overestimate the abundance of C4 vegetation in agiven environment. Pedogenic carbonate may also be contaminatedby inherited carbonate from the material on which the soil formed.CO2 released from limestone or other sedimentary rock with highcarbonate content during pedogenesis can mix with CO2 from plantroots and influence the carbon isotopic composition of pedogeniccarbonate nodules (Sheldon and Tabor, 2009). Isotopically heavier CO2

derived from limestone incorporated in pedogenic carbonates wouldproduce an overestimation of % C4 plants. Pedogenic carbonates canonly be used to estimate abundance of C4 vegetation with carefulsampling and identification of soil parent material (Cerling and Quade,1993). Fossil tooth enamel from mammals has also been used toreconstruct past vegetation (e.g. Wang et al., 1994; Cerling et al., 1997;Passey et al., 2002) but this may also lead to biased reconstructions dueto the dietary preference of the animal, especially in relatively lowpercentage C4 environments.

For this study, organic material preserved in late Miocenepaleosols was used to produce a high resolution C3/C4 vegetationreconstruction. Preserved organic material represents the isotopiccomposition of bulk soil organic material (SOM) of the original soil.Given that the ultimate source of soil organic carbon in soils isvegetation, organic material preserved in paleosols can be used forC3/C4 vegetation reconstructions (Koch, 1998).

Phytoliths preserved in the same late Miocene paleosols were alsoused for a multiproxy vegetation reconstruction and to validate theisotopic results. Phytoliths are silica bodies produced by many vascularplants, which range in size from 5 to 500 μm and are thought to beproduced for structural support, as a waste product, or to deter

herbivory (Jacobs et al., 1999). Phytoliths are precipitated in intra andintercellular spaces frommonosilicic acid (H4SiO4) that is taken upwithgroundwater (Piperno and Pearsall, 1998). Morphologies of phytolithstend to mimic the shape of their surrounding cell, and are oftentaxonomically distinct. Due to this taxonomic sensitivity, includingdistinct morphotypes created by C3 and C4 grasses, phytoliths can beused for reconstructions of past vegetation (Piperno and Pearsall, 1998;Strömberg, 2002, 2004). Phytoliths have proven to be especially usefulfor grassland vegetation reconstructions because of the lack of macro-scopic grass fossils in the geologic record (Thomasson, 1987; Strömberg,2002). As such, numerous recent studies have focused on the origin andexpansion of the grassland ecosystem using phytolith interpretations(e.g. Strömberg, 2004, 2005; Zucol et al., 2010). Notably, Strömbergand McInerney (2011) use phytoliths to document the shift from C3to C4 grasses in the Great Plains of North America during the Mid-Miocene through Pleistocene, and find an increase in phytolithsproduced by C4 grasses between 8 and 5 Ma.

This study reconstructs the vegetation of an ancient grassland andcorrelated this with paleoclimatic reconstructions in order to under-stand better the climatic conditions associated with the expansion of C4grasses, An outcrop in southwestern Montana dated at 10.2–8.9 Ma(Fritz et al., 2007) was examined in order to study the local distributionof C4 grasses prior to the global expansion that occurred between 8 and3 Ma (Cerling et al., 1997; Fox and Koch, 2003; Behrensmeyer et al.,2007). Studying the dynamics of grassland ecosystems during this timeperiod is essential to understanding the environmental conditionsassociatedwith the global expansion of C4. The isotopic results from the34m thick succession in Montana were also compared to previouslypublished estimates based on pedogenic carbonates from other regionsaround the world (Quade and Cerling, 1995; Latorre et al., 1997; Foxand Koch, 2003).

2. Methods

The sample site is located in southwesternMontana and is part of theSixmile Creek Formation, a unit within the Bozeman Group (Fig. 1;Hanneman andWideman, 1991). The 34 m section (Fig. 2) is composedof calcium carbonate cemented, fluvially derived sediments and 35compound paleosols formed at Timber Hills, Montana (Fig. 1) during theMiocene. In addition to pedogenic features such as horizonation and roottraces, many of the beds are heavily burrowed and contain abundanttrace fossils. The majority of paleosols preserved at this site wereidentified as Inceptisols (Soil Survey Staff, 2010). The section beginswitha tuffaceous bed that has been dated to 10.2±0.4 Ma. This tuffaceousbed sits 112 m below the Timber Hills basalt, which has been dated to5.9±0.2 Ma (Fritz et al., 2007). Approximate ages for each bed (Fig. 2,Table 1) were determined by applying a linear sedimentation rate basedupon the above dates of 26 mMa−1 (Retallack, 2007) to the section. Theresulting age model estimates that this 34-meter section spans 10.2 to8.9±0.4 Ma ago.

After 10 cm of surface material was removed, fresh paleosolsamples were collected both in vertical (to assess changes in theecosystem through time) and lateral sections (to assess spatialvariability). Samples were also collected from depth profiles througheight individual paleosol profiles at 15 cm intervals to determine theδ13Corg relationship with depth in the soil. Lateral transects along fivepaleosols were collected within the top 15 cm of the bed at intervalsof 5 to 10 m. In addition, three modern soils were sampled andanalyzed to confirm the accuracy of the mixing model in predictingpresent-day vegetation cover by comparing δ13Corg with the modernabundance of C4 vegetation.

To prepare for bulk δ13Corg analysis, all samples were firstultrasonicated in methanol to remove any modern organic material.Samples were then treated with a 7% solution of HCl at 40 °C until allcarbonate had been reacted and samples no longer fizzed. Sampleswere then rinsed in deionized water. The dried samples were then

90 J.M. Cotton et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 337–338 (2012) 88–98

homogenized and loaded into tin capsules and analyzed using aCostech elemental analyzer attached to a Finnigan Delta V+ isotoperatio mass spectrometer at the University of Michigan. Each samplewas analyzed in triplicate. Thin sections of pedogenic carbonatenodules were examined for evidence of diagenesis (e.g., sparry ratherthan micritic calcite; Sheldon and Tabor, 2009). The micritic noduleswere then microsampled to avoid spar and analyzed on a Thermo-Finnigan MAT 253 isotope ratio mass spectrometer with KielIVautosampler at the University of Michigan. Carbon and oxygenisotopic ratios are reported in units of per mil (‰) relative to theinternational standard Vienna Pee Dee Belemnite. The internalstandards used to normalize measured values to V-PDB were IAEAsucrose and caffeine for organic carbon analysis and NBS 19 forcarbonate analysis. The analytical uncertainty for the organic carbonmeasurements is b0.10‰ with a mean standard deviation of 0.22‰for triplicate analyses of each sample. The analytical error is 0.03‰ forinorganic carbon and 0.05‰ for oxygen isotopes. Analytical errorswere determined through repeated measurements of referencematerials. Paleosols were identified as containing C4 biomass if theirisotopic composition was greater than −24.7‰, based on the averageisotopic composition of a pure C3 ecosystem in a precipitation regime of1000 mm yr−1 or less, adjusted for the change in δ13C of atmosphericCO2 in the late Miocene (Passey et al., 2009; see discussion below).Estimates for percent C4 vegetation for each paleosol (Table 1) weredetermined using a two-component mixing model with a pure C3 end-member of −24.7‰ and a pure C4 end-member of −11‰.

Phytoliths preserved in paleosols were also used to producequantitative paleovegetation reconstructions to compare with thatproduced from isotopic records. Phytoliths were extracted from the Ahorizons of eight paleosols from the section. Extractions and quantita-tive paleovegetation reconstructions followed the methods outlined byStrömberg et al. (2007). Subsamples of paleosols were crushed andtreated with HCl to remove CaCO3. The 0–250 μm size fraction wasseparated by sieving and clays were removed through centrifuging.The biogenic silica was then isolated by heavy liquid (Zinc Bromide)separation. Once extracted,morphotypeswere counted andmatched tothe corresponding vegetation type and then converted to relativepercentages based on the total number of phytoliths extracted fromeach sample. The abundance of C4 phytoliths (Table 2)was calculated as

Fig. 1. Location of studied trace fossil site (TFFT) in Southwest Montana lo

a range of percentages, because some morphologies are produced onlyby C4 grasses, while othermorphologies are produced by both C3 and C4grasses. Examples of such morphotypes are displayed in Fig. 3. Theminimum abundances represent phytolith morphotypes that are onlyfound in C4 grasses. The maximum abundances include phytolithmorphotypes that are made strictly by C4 grasses and those that aremade by both C4 and C3 grasses and cannot be distinguished, but whichare nonetheless taxonomically indicative of grasses (Table 2). Theabundance of C4 vegetation determined by analysis of organic matterwas compared to that determined by phytolith extraction to validatethe use of isotopes in paleovegetation reconstructions.

3. Results

3.1. δ13C of preserved organic matter

The mean δ13Corg for this section is −23.9±1.0‰ (±1σ) with arange of −26.5‰ to −19.9‰. Fig. 2 presents the mean δ13Corg foreach paleosol based on triplicate analyses. The gaps displayed in thedata (Fig. 2) correspond to four covered beds (3, 6, 27, and 31) and toone paleosol (25) that had too little preserved organic carbon foranalysis. The range in δ13C values within profiles varied from 0.3 to3.5‰ (Fig. 4). Variability of carbon isotopic composition along strikewithin the beds ranged from 1.7 to 4.0‰, with the greatest range inisotopic composition in the longest transect (Fig. 5). Both depthprofiles and lateral transects show substantial isotopic variability.

δ13Corg was also measured in depth profiles of three modernanalog soils from southwestern Montana to compare modern organicmaterial to Miocene preserved organic material. Modern soil organicmatter (mean of three soil profiles) from the same region of Montanaaverages−24.3±0.8‰with a range of 2.6‰ (Supplementary Table 7).Because soil formation is a slow process (hundreds to thousands ofyears), this modern SOM data reflects the pre-industrial Holoceneatmospheric CO2 composition of −6.5‰. The average isotopic compo-sitions of these modern soils are similar to the Miocene paleosols at−23.9‰, but this area is not a pure C3 system, and typically contains10–20% C4 vegetation (Epstein et al., 2002). Thus, based on our isotopicreconstruction it is most likely that the late Miocene contained similaramounts of C4 biomass to the modern ecosystem.

cated at N45° 02′ 26.7″, W112° 15′ 18.3″ with an elevation of 1976 m.

Clay Silt Fine Med. Coarse Conglom.

5

10

15

20

25

30

Meters

Age

(M

a)

δ Corg (‰)

*

°

9.8

9.6

9.4

9.2

9.0

* Profile Sampled

Conglomerate

Sandstone

Siltstone

Mudstone

Burrows

Root Traces

Carbonate

Nodules

Cross Bedding

° Laterally Sampled

Rooted Nodules

Covered Section

10.2-26 -25 -24 -23 -22 -21

+ Phytolith Analysis

+

10.0

+

+

*

*

*

*

*

*

*°

°

°

°

Pure C3Ecosystem

Mixed C3/C4Ecosystem

Sand

10 200% C4

Fig. 2. Left: stratigraphic column for Montana site. The crosses in the stratigraphy represent covered beds. Right: Average δ13Corg for each stratigraphic bed based on at least threereplicate analyses. % Corg for this section ranged from 0.02 to 0.37% with an average of 0.07±0.01%.

91J.M. Cotton et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 337–338 (2012) 88–98

3.2. Phytolith assemblages

Strömberg (2004, 2005) demonstrated that grass phytoliths arepreserved in paleosols and that their different morphologies can beused for paleovegetation reconstructions. Of the eight stratigraphiclevels analyzed for phytolith assemblages, all levels yielded phyto-liths, but only three levels yielded enough phytoliths for quantitativepaleovegetation reconstruction. Fig. 3 shows a comparison betweenmodern grass phytolith morphologies and those observed in thepaleosols from Montana. Fig. 3A and B are crenate grass silica short cell(GSSC), amorphotype produced by pooid grasses, which are all C3 plants.Fig. 3C and D is saddle GSSC, a morphotype produced by chloridoidgrasses, which are all C4 plants. Fig. 3E and F shows cross GSSC, amorphotype produced by panicoid PACMAD grasses, many of which areC4 grasses. For stratigraphic Bed 2 (0.75 m), C4 grassesmadeup 7.0–17.5%of the overall vegetation, and comprised 17.9–44.8% of the grasscommunity. Stratigraphic Bed 19 (10.0 m) contained 0–10.8% C4 grassesthat comprised up to 13.8% of the total grass community. StratigraphicBed 38 (31.75 m) contained 3.4–17.2% C4 grasses in the overall

vegetation and also made up 3.7–18.9% of the grass community.Phytolith assemblages are summarized in Table 2 and SupplementaryTable 7. The low end of the abundance ranges is a conservative estimateof the amount of C4 biomass based on percentages of phytolithsproduced strictly by C4 grasses. The existence of C4 phytoliths in thesepaleosol samples reaffirms the results of the presence of C4 biomass inmany paleosols in this section interpreted from the isotopic data. Theresults of the phytolith reconstructions are compared to the δ13Corg ofthe same stratigraphic level in Fig. 6.

3.3. δ13C and δ18O of pedogenic carbonates

Pedogenic carbonates from three different paleosols were analyzedfor δ13C and δ18O. Three nodules from each paleosol were analyzed toassess isotopic variability within each paleosol. δ18Ocarb ranged from−17.0 to −15.9‰ with an average of −16.5±0.4‰ (V-PDB). δ13C ofthe paleosol carbonates ranged from −4.2 to −2.8‰ with an averageof −3.6±0.5‰. Non-pedogenic matrix carbonate (i.e., cement) wasalso analyzed from three different beds and had an average δ18Ocarb

Table2

Qua

ntitativeph

ytolithreco

nstruc

tion

sof

samples

from

tracefossilsite

inSo

uthw

estM

ontana

follo

wingthemetho

dsou

tlines

byStrömbe

rg(2

004)

.Eight

leve

lswerean

alyz

edforph

ytolithassemblag

es,b

uton

lythreeleve

lsyielde

den

ough

phytoliths

forqu

antitative

vege

tation

reco

nstruc

tion

s.

Bed

numbe

rSa

mple

numbe

rC 4

abun

danc

ein

overallv

egetation

C 4ab

unda

ncewithingrassco

mmun

ity

Minim

umestimate

Max

imum

estimate

Minim

umestimate

Max

imum

estimate

(PAN+

CHLO

R)/(GSS

C−

OTH

G)*

GSS

C/

(FIT

OT+

GSS

C+

AQ)

(%)

95%C.I.

(%)

(PACM

AD

TOT)

/(GSS

C−

OTH

G)*

GSS

C/

(FIT

OT+

GSS

C+

AQ)

(%)

95%C.I.

(%)

(PAN+

CHLO

R)/

(GSS

C−

OTH

G)

(%)

95%C.I.

(%)

(PACM

AD

TOT)

/(G

SSC−

OTH

G)

(%)

95%C.I.

(%)

Bed1

MT0

8-04

17.0

(4.1–9.9)

17.5

(13.1–

22.2)

17.9

(11.2–

25.5)

44.8

(35.8–

54.0)

Bed4

MT0

8-04

2N/A

aN/A

N/A

N/A

N/A

N/A

N/A

N/A

Bed5

MT0

8-04

3N/A

N/A

N/A

N/A

N/A

N/A

N/A

N/A

Bed11

MT0

8-04

4N/A

N/A

N/A

N/A

N/A

N/A

N/A

N/A

Bed12

MT0

8-04

5N/A

N/A

N/A

N/A

N/A

N/A

N/A

N/A

Bed19

MT0

8-04

80.0

0.0

10.8

(6.4–15

.8)

0.0

0.0

13.1

(7.7–18

.5)

Bed38

MT0

8-05

33.4

(0.7–6.4)

17.2

(11.6–

23.4)

3.7

(0.7–6.9)

18.9

(12.8–

25.6)

aN/A

=ph

ytoliths

presen

t,bu

tfewer

than

200diag

nostic

morph

otyp

eswereco

untablein

thesample.

Table 1δ13Corg values for each stratigraphic bed at trace fossil site (TFFT) in southwesternMontana. Bed 1 is a volcanic ash dated at 10.2±0.4 Ma (Fritz et al., 2007). Ages aredetermined by linear interpolation of a regional sedimentation rate from Retallack(2007).

Sample number Bed number Age %Corg δ13Corg St. Dev. n %C4

JTFF-1o 2 10.2 0.37 −26.47 0.82 3 0.0Covered 3JTFF-2o 4 10.1 0.04 −25.21 0.48 3 0.0JTFF-3o 5 10.1 0.11 −26.53 0.19 3 0.0Covered 6 10.1JTFF-4o 7 10.1 0.04 −23.30 0.47 3 10.2JTFF-5o 8 10.1 0.06 −24.87 0.07 3 0.0JTFF-6o 9 10.1 0.03 −23.18 0.07 3 11.1JTFF-7o 10 10.1 0.04 −24.25 0.28 3 3.3JTFF-8o 11 10.0 0.03 −23.57 0.11 3 8.2JTFF-9o 12 10.0 0.04 −23.55 0.29 3 8.4JTFF-10o 13 10.0 0.03 −22.24 0.89 33 18.0JTFF-11o 14 10.0 0.04 −23.20 0.07 3 11.0JTFF-13o 15 10.0 0.04 −23.58 1.30 39 8.2JTFF-14o 16 9.9 0.04 −24.44 0.06 3 1.9JTFF-15o 17 9.9 0.04 −23.36 0.01 3 9.8JTFF-16o 18 9.9 0.06 −24.63 1.07 12 0.5JTFF-17o 19 9.9 0.08 −23.69 0.03 3 7.4JTFF-18o 20 9.8 0.03 −24.51 1.13 15 1.4JTFF-19o 21 9.8 0.07 −23.59 0.57 30 8.1JTFF-21o 22 9.7 0.05 −24.74 0.25 3 0.0JTFF-20o 23 9.6 0.07 −24.63 0.11 3 0.5JTFF-22o 24 9.6 0.05 −23.13 0.13 3 11.5JTFF-23o 25 9.6JTFF-24o 26 9.6 0.10 −24.97 0.09 3 0.0Covered 27 9.6JTFF-25o 28 9.3 0.06 −23.62 0.54 3 7.9JTFF-26o 29 9.3 0.04 −23.95 0.36 18 5.5JTFF-27o 30 9.2 0.03 −23.13 0.04 3 11.5Covered 31JTFF-28o 32 9.2 0.12 −22.64 0.11 3 15.0JTFF-29o 33 9.1 0.11 −23.24 0.40 3 10.7JTFF-30o 34 9.1 0.04 −22.69 0.08 3 14.7JTFF-31o 35 9.1 0.09 −23.23 0.34 12 10.7JTFF-32o 36 9.1 0.04 −23.30 0.96 15 10.2JTFF-33o 37 9.0 0.08 −23.80 1.43 27 6.5JTFF-34o 38 9.0 0.09 −24.02 0.02 3 5.0JTFF-35o 39 9.0 0.18 −25.64 0.17 3 0.0

92 J.M. Cotton et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 337–338 (2012) 88–98

and δ13Ccarb composition of −16.7±0.3‰ and −3.2±0.2‰, respec-tively. The pedogenic carbonate nodules were isotopically indistin-guishable from the surrounding cement.

4. Discussion

4.1. Relationship between C3 photosynthesis and climate

Aridity can positively shift the isotopic composition of C3 plantsdue to water stress (Ehleringer and Monson, 1993). Paleoclimaticreconstructions of mean annual temperature and precipitation forthis site and surrounding areas were used to determine if ariditycould have produced the isotopically heavy carbon measured in thissection, and thus changed the C3 isotopic end member for our % C4reconstructions. The following is a brief overview of methods forpaleoclimatic reconstructions using paleosols (see Sheldon and Tabor(2009) for a detailed review of quantitative reconstructions). Meanannual temperature can be estimated using the Al2O3/SiO2 weath-ering ratio climofunction for Inceptisols described by Sheldon (2006)according to the following equation:

MAT ¼ 49:94C þ 3:99 ð1Þ

C ¼ mAl2O3=mSiO2 ð2Þ

where r2=0.96 and SE=±0.6° (analytical uncertainty and SE totalabout ±2°). C represents the clayeyness of the soil, which is related to

Fig. 3. Comparison of modern grass phytoliths to those extracted from beds at Montana site. A: Genus Nasella C3 pooid grass produces crenate grass silica short cell (GSSC)phytoliths. B: Crenate GSSC extracted from Bed 2 from the stratigraphic section. C: Genus Eragrostis C4 chloridoid grass produces saddle GSSC phytoliths. D: Saddle GSSC extractedfrom sample level 2. E: Genus Arundo C4 panicoid PACMAD grass produces cross GSSC phytoliths. F: Cross GSSC extracted from sample level 2 from the site.

93J.M. Cotton et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 337–338 (2012) 88–98

climate through weathering. Using Eqs. (1) and (2), mean annualtemperature for this section was determined to be 9.2 °C±0.9 °C (±1σ) based on whole rock geochemical analyses of 13 of the paleosolsspanning the entire section (see Supplementary Data). Precipitation canbe estimated using two methods based on paleosols. The depth to Bkhorizon increases with increasing MAP (Retallack, 2005), and is relatedby the following equation:

P ¼ 137:24þ 6:45D−0:013D2 ð3Þ

where r 2=0.52 and SE=±147 mm yr−1 and D represents the depthto Bk horizon. The chemical index of alterationwithout potassium (CIA-K) is an indication of degree of weathering (Maynard, 1992), and is

Fig. 4. Depth profiles of δ13Corg for seven individual paleosol profiles. Maximumvariability in a single profile is 3.51‰.

calculated from the molar ratios of Al2O3 CaO and Na2O determined byXRF whole rock analysis.

CIA� K ¼ mAl2O3

mAl2O3 þmCaOþmNa2Oð4Þ

The CIA-K is related to precipitation by the following equation(Sheldon et al., 2002):

P ¼ 221e0:0197 CIA�Kð Þ ð5Þ

where r2=0.72 and SE=±182 mm yr−1.Retallack (2007) reported paleoprecipitation estimates based on

11 depths to Bk horizons another nearby Montana site contemporane-ous to this measured section with an average of 385±147 mm yr−1

with a standard deviation of 71 mm yr−1. Precipitation reconstructionsusing the CIA-K climofunction described by Sheldon et al. (2002) fromtwo contemporaneous B horizons from other localities (SweetwaterCanyon and Timber Hill, MT) within 3 km of the site sampled in thisstudy yield an estimate of 809±182 mm yr−1 (Retallack, 2007). Thepreviously published precipitation estimates are compiled in Tables 5and 6 of the Supplementary Data. Hamer (2009) reconstructed MAPusing the CIA-Kmethod for another contemporaneous site in SouthwestMontana and reported an average of 640±182 mm yr−1 precipitation.Though the twomethods ofMAP reconstruction give estimates that varysignificantly, it is likely that this area received an average precipitationsimilar to that predicted by the CIA-K due to the abundance of tracefossils and lack of pedogenic carbonate in most of the paleosols (e.g.,Sheldon and Tabor, 2009). In North America, the boundary betweencalcareous and non-calcareous soil is 750 mm yr−1 (Retallack, 2000).Given that only six out of thirty-five paleosols in this study containpedogenic carbonate, the local climate was probably relatively wet withprecipitation closer to the CIA-K estimates. The six levels containingcarbonates could represent brief dry periods punctuating an otherwisewet climate. If the climate was as wet as the CIA-K reconstructionsestimate, it is unlikely that the plants were water stressed, so a large

Fig. 5. Plot of lateral variability of δ13Corg for five paleosol beds. Maximum variability along a single transect is 3.97‰.

94 J.M. Cotton et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 337–338 (2012) 88–98

isotopic shift in C3 biomass would not be expected. Thus, the isotopicend member for a pure C3 ecosystem for our site in the late Mioceneshould not be significantly heavier than the average isotopic composi-tion of C3 plants after correction for shifts in δ13Ca.

For this reconstruction, an isotopic composition of −24.7‰ wasassigned as the threshold for pure C3 systems. This pure C3 thresholdwas chosen using two different literature compilations of modernplant δ13C and its relationship to precipitation and focusing onprecipitation similar to that reconstructed for our sample site.Diefendorf et al. (2010) published 159 isotopic compositions fromplants growing in 1000 mm yr−1 precipitation or less. Adjusting forthe shift in δ13Ca to the Miocene atmospheric value of −6‰ (Passeyet al., 2009; Tipple et al., 2010), the average isotopic composition ofthese plants is −24.6‰. Kohn (2010) published 360 isotopic composi-tions of plants growing in 1000 mm yr−1 precipitation or less. Adjustingfor the shift in atmospheric CO2, the average isotopic composition ofthese plants is −24.7‰. Because these two studies both arrived at thesame average isotopic composition of C3 plants growing in sub-humid toarid regions, this gives us confidence that our assumption of−24.7‰ asa C3 end member for paleovegetation reconstruction is valid. This endmember takes into consideration any water stress influencing the C3plants that could be interpreted as a C4 signature. This value of−24.7‰was used as a threshold for identification of C4 vegetation. Any isotopiccomposition heavier than thiswas assumed to contain some C4 biomass.Of the 34 paleosols analyzed, 26 had isotopic compositions heavier than

Fig. 6. Comparison between C4 reconstructions from isotopic method and phytolithmethod. Both methods of reconstruction generally agree on low but measurableamounts of C4 biomass for this site.

−24.7‰, indicating a presence of C4 biomass in almost all paleosolsidentified in this section.

The percentage of C4 vegetation from soil organic matter can beestimated using a two-component mixing line with one end memberrepresenting the isotopic composition of a pure C3 system (−24.7‰)and the other endmember representing a pure C4 ecosystem (−11‰).For example, a modern soil from southwestern Montana analyzed inthis study has an average δ13Corg of−23.4‰. Assuming a pre-industrialvalue of −6.5‰ for δ13Ca of the atmosphere (Friedli et al., 1986), ourpure C3 end member would be −25.2‰ (0.5‰ lighter than the lateMiocene value) and our pure C4 endmember would be−11.5‰. Usingthis mixing model, the soil organic material would be comprised ofapproximately 13%C4 biomass. If C3 and C4 organicmaterial are degradedat approximately the same rate, the vegetation growing above this soilwould be about 13% C4. This reconstruction is well in line with theestimates for abundance of C4 vegetation of 10–20% for this areapublished by Epstein et al. (2002). Similarly, the percentage of C4vegetation can be estimated using the late Miocene paleosols with thetwo-component mixing model end members of –24.7‰ for C3 and−11‰ for C4 (readjusted to the Miocene atmosphere of −6‰). Usingthis method, the amount of C4 vegetation varies from 0 to 18%throughout the section. Individual estimates for each paleosol are listedin Table 1.

4.2. Phytolith and δ13Corg comparison

Phytolith assemblages and conservative organic isotopic estimatesof C4 biomass are generally in agreement. Using −24.7‰ as a C3 endmember, the percentages of C4 biomass are calculated to be 0%, 7%and 5% for stratigraphic levels 2, 19, and 38, respectively (Table 1).These levels correspond with the same samples analyzed forphytolith analysis discussed above. Fig. 6 shows that for two out ofthe three samples that yielded enough phytoliths for quantitativereconstruction, the paleovegetation reconstruction obtained usingthe isotopic data estimates C4 biomass on the low end of what isreconstructed using phytoliths. For beds 19 and 38, both vegetationreconstructions using isotopes and phytoliths produce amounts of C4

biomass within error estimates, showing that these two methodsgenerally produce similar results. However, for Bed 2 the twomethods of reconstruction do not match, with the phytolithestimation of up to 17.5% C4, while the isotope record indicates noC4 biomass. There are several possible explanations for this discrep-ancy. These differences could be caused by changes in C4 biomass atthis location through time. As shown in the δ13Corg lateral profile data(Fig. 5), there is significant variation in isotopic composition oforganic material along a profile, which translates to spatial variabilitywithin vegetation depending on local topography and proximity tostreams. Phytoliths accumulate in the A horizons of soils through time

95J.M. Cotton et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 337–338 (2012) 88–98

and represent a time-averaged signature, but as the profile builds, theorganic carbon at each level preserves the isotopic composition of theinput plant material, which changes through time as a result ofecosystem-scale changes (e.g., succession, fire, etc.). Thus, the δ13Corgmay represent mainly the “last” vegetation at the site, whereas thephytoliths should record all of the vegetation at a site. Alternatively,if the vegetation shifts, a shift can be identified in the isotopiccomposition of carbon in the soil profile as a function of depth, butnot in observations of a single A horizon (Monger et al., 2009), so thediscrepancy between phytolith and isotope results for Bed 2 couldrepresent different time periods in the development of a single soil.

The difference between the estimates of C4 vegetation betweenδ13Corg and phytoliths may also be due to the differential degradationof organic matter. Wynn and Bird (2007) show that in a mixed C3/C4area, the C4 organic material is degraded more quickly than the C3organic material, suggesting that estimates of C4 vegetation usingδ13Corg of preserved organic material are likely minimum estimates.The differential degradation would cause the organic material pre-served in the soil to be isotopically lighter than the average vegetationinput and would lead to lower estimates of C4 vegetation producedfrom isotopic data compared to the estimates produced from thephytolith data, This expected trend is observed in all three paleosolswith isotopic and phytolith data (Fig. 6), with the δ13Corg under-estimating the amount of C4 vegetation compared to the phytoliths.

4.3. Reliability of isotopic data in establishing C4 photosynthesis

Though the abundance of organic carbon for this section is low(0.07±0.01%; Fig. 7, Supplemental Data), it is unlikely that thiscarbon is detrital because it is preserved primarily in black carbonrootlets that appear to be in situ within the paleosol profiles. Therootlets spread out downward, commonly in dense root mats, fromthe surfaces of the paleosols. Kuenzi and Fields (1971) interpreted themajority of the Sixmile Creek formation as a low energy fluvialenvironment. Because this section is comprised of mostly fine grainedmaterial (Fig. 2) and only contains two paleochannels, the largestbeing only 1 m deep, it is likely that this section also represents a lowenergy depositional environment. If any detrital carbon was depos-ited in these sediments and was preserved, it is likely that detritalcarbon would have come from the same catchment and shouldpreserve evidence of the same regional ecosystem biomass signal.

Fig. 7. Plot of fSOC vs. δ13C for typical paleosol (bed 15) found in this section against theexpected Rayleigh distillation curve for modern soils (Wynn et al., 2005; Wynn, 2007).fSOC plots for the remaining seven profiles can be found in Supplementary Fig. 1.

There is large variability in the isotopic composition of carbonwith depth throughout individual paleosol profiles. In many modernsoils, Rayleigh distillation is observed with a trend of increasingisotopic composition with depth, due to the microbial decompositionof organic matter. Wynn (2007) observed a 2 to 6‰ positive shift inδ13Corg with increasing depth in the soil profile, depending on thetexture of the soil and the fraction of organic matter remaining. Fig. 7shows an example of a typical paleosol profile from this section. Alsoplotted on the graph is the expected generalized trend of Rayleighdistillation for organic carbon in soils described by Wynn (2007). Forthis plot, fsoc represents the normalized remaining fraction (C) of theoriginal carbon concentration at the surface (Cs) according to theequation below.

f SOC ¼ CCs

ð6Þ

It is clear that this profile data does not exhibit the Rayleighdistillation described by Wynn (2007). The remaining profiles (Supple-mentary Fig. 1) are similar and do not exhibit Rayleigh distillation either.Monger et al. (2009) found that changes in vegetation can be recorded inthe isotopic composition of pedogenic carbonates with depth through-out a single soil profile. The variability in the isotopic composition ofthese soil profiles is not an artifact of microbial decomposition, but morelikely a signature of vegetation changes through time similar to thoseobserved by Monger et al. (2009). Organic matter preserved at differentdepths in the soil profile will behave similarly, recording the isotopiccomposition of the material at the time of deposition. Changes in theamount of C4 vegetation through time could lead to the observed patternin isotopic composition with depth.

Large lateral variation is also observed in δ13Corg of SOM (Fig. 5).Along with the variation with depth in soil profiles, this variationsuggests that single spot sampling in many cases may not be sufficientto reconstruct accurately vegetation for a given time period. The onlyway to determine if spot sampling is representative is to analyzemultiple samples along strike to assess the variability of past vegetation.The variabilitywithin δ13Corgwas roughly the same through time as it isacross a 55 m lateral transect, indicating substantial ecosystem hetero-geneity. Given that the lateral variation in δ13Corg virtually the sameas the vertical variation in δ13Corg, it is likely that all the variability inthis site could be caused by local heterogeneity in the vegetation.Behrensmeyer et al. (2007) also found significant variation in δ13C ofcarbonates from paleosols in Pakistan dated to 6.74 to 4.85 Ma,supporting these results that ecosystem variation is preserved on themeter scale.

4.4. Comparison to pedogenic carbonate methods

Assumptions about the isotopic fractionation of carbon duringcarbonate formation can lead to large error in % C4 reconstruction. Thetemperature of soil formation must also be taken into considerationwhen calculating the isotopic fractionation of carbon during theprecipitation of carbonate because that fractionation is temperaturedependent and decreases at higher temperatures (Romanek et al.,1992). Carbon isotopic compositions are enriched 14‰ to 17‰relative to biomass during the precipitation of the carbonate nodules(Cerling, 1984; Koch, 1998). The range in pedogenic carbon isotopeenrichment (3‰) equates to a range of 21% C4 plants. The Fox andKoch (2003) study assumes that the Δ13C between organic carbonand carbonate is 15.5‰. For the 20 organic carbon isotope data pointsreported in their study, the average Δ13C is 16.2‰ with a standarddeviation of 1.9‰. Fox and Koch (2003) report four organic carbonisotope data points obtained from roughly the same time period asthis study. Their δ13Corg value at 9.1 Ma is −23.7‰, which translatesto 12.8% C4 plants. Their δ13C value of carbonate from the samepaleosol is−6.0‰, which gives a Δ13C value of 17.8‰. Calculating the

96 J.M. Cotton et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 337–338 (2012) 88–98

percentage of C4 plants using the carbonate isotope data using a Δ13Cvalue of −15.5‰ produces a value of 28.6% C4 plants, significantlyoverestimating the percentage determined by preserved organicmatter. At 9.55 Ma, their δ13Corg is −20.5‰, which corresponds to35.7% C4 plants. The δ13C value of carbonate from the same paleosolis −7.6‰, giving a Δ13C of 13.4‰. Using the assumption ofΔ13C=15.5‰, this gives a C4 value of 20.7%, underestimating thepercentage compared to the percentage determined by organiccarbon isotopes.

It is clear that the Δ13C used to convert δ13C of carbonate torespired organic carbon δ13Corg has an effect on estimates of C4biomass, which may have caused the Fox and Koch (2003) study tooverestimate the percent C4 composition in their reconstructions ofGreat Plains grasslands. While some data points show lower Δ13Cthan the assumed 15.5‰, the majority of the 20 data points have alarger Δ13C than was assumed in the calculations, which wouldincrease the predicted amount of C4 vegetation. Recalculating theirpercent C4 using their average Δ13C of 16.2‰ for the time period of10.2 to 8.9 Ma, the average percent C4 biomass contribution to SOM isreduced from 22.1% to 17.4%, which is more comparable to the resultsfrom this study. These differences may also be in part an artifact of theuse of carbonate versus organic matter to determine the proportion ofC4 photosynthesis, although there may be real regional differencesbetween the Great Plains and Montana in the late Miocene as thereare today (Epstein et al., 2002),

The pedogenic carbonates collected from the section in Montanacannot be used to quantify C4 biomass because the carbon and oxygenisotopic composition of both of the nodules are indistinguishablefrom the surrounding cementing material, indicating post burialalteration (Fig. 8). The Δ13C value for a soil can also be indicative ofalteration (Sheldon and Tabor, 2009).

The Δ13C for the three beds (Beds 13, 37 and 39) are 18.2‰, 20.3‰and 22.4‰, much greater than the 14–17‰ range for fractionationduring the formation of carbonate (Cerling and Quade, 1993). HighΔ13C values can arise from a few different scenarios, namely lowproductivity ecosystems that allow for greater diffusion of isotopi-cally heavier atmospheric CO2 into the soil, limestone parent materialwhich upon weathering will contribute isotopically heavier CO2 tothe soil air, and groundwater recrystallization after burial which willreset the isotopic compositions of the pedogenic carbonates (Sheldonand Tabor, 2009). The paleosols in this section were parented onfluvially derived sediments of igneous origin, which excludes parentmaterial as a cause for the high Δ13C values. The presence of multiple

Fig. 8. Plot of the isotopic composition of pedogenic carbonates nodules and carbonatematrix from three beds (13, 37 and 39) at Montana site. The carbonate nodules haveisotopic compositions that closely match the diagenetic carbonate matrix. The standarderror for these measurements are 0.05‰ for oxygen and 0.03‰ for carbon and the errorbars are smaller than the plot symbols.

large (likely mammal) burrows as well as the sub-humid precipita-tion regime estimated from nearby sites (as discussed in Section 4.1)suggest that the area was not a low productivity ecosystem and thuswould not cause Δ13C values to be high. The remaining explanation ofgroundwater recrystallization is likely the reason that these soils havehigh Δ13C values. This explanation agrees with observations that theoutcrop is cemented mostly by a carbonate-rich cement as well asisotopic data showing that the nodules are isotopically indistinguish-able from the cementing material. Thus, the carbonate nodules in thissection are unreliable for paleovegetation reconstructions. Using thesecarbonates to reconstruct C4 biomass would yield unrealistically highvalues of 40–50% C4 plants, which does not agree with the re-constructions based on both δ13C of organic matter and phytoliths.

4.5. Comparison to contemporaneous records of C4 photosynthesis

There are several studies that present estimates of C4 biomassduring the same time period as this work. Fig. 9 compares the resultsof this study to three other studies using pedogenic carbonates formother regions of the world. Isotopic data from these previous studieswas converted to % C4 biomass using the same method as Fox andKoch (2003). Carbonates from Pakistan (Quade and Cerling, 1995;Behrensmeyer et al., 2007) show little, if any signal of C4 photosyn-thesis until 8–7 Ma ago. The three samples from the same time periodas this study in Argentina (Latorre et al., 1997) translate to between 5and 10% C4 plants. Though climatic reconstructions of the Great Plainsof North America (Retallack, 2007) estimate similar MAT and MAP tosouthwestern Montana during this time period, the C4 vegetationreconstructions from the Great Plains (Fox and Koch, 2003) aregreater than all but two data points from this study in Montana. Theresults from Montana fall in between the majority of the previousdata for 10.2 to 8.9 Ma, with a larger C4 signature than in Pakistan buta smaller signature than the Great Plains. If the estimates of C4

photosynthesis from Latorre et al. (1997) are an overestimation dueto the problems with pedogenic carbonate discussed above, then theC4 signal in Argentina would be essentially nonexistent. Both phytolithand δ13Corg show that there was a small but measurable proportion ofC4 vegetation present in Montana before the localities in Pakistan andArgentina. The presence of C4 biomass in only North America duringthis time period suggests that the pathway may have evolved in NorthAmerica. Edwards et al. (2010) supports this suggestion, showing theearliest evidence for C4 photosynthesis in both paleosol carbonate and

Fig. 9. Data from Montana (this study) compared to Pakistan (Quade and Cerling,1995), the Great Plains (Fox and Koch, 2003), and Argentina (Latorre et al., 1997) forthe same time period (10.5–8.5 Ma).

97J.M. Cotton et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 337–338 (2012) 88–98

tooth enamel is found in the early Miocene in the Great Plains, butfurther studies are needed to assess C4 origination and its spread.

Some of the variation between this dataset and the data obtainedfrom pedogenic carbonates may be attributed to the problemsmentioned above as a result of relying on pedogenic carbonates aswell as some assumptions necessary to use the proxy. While theorganic matter preserved in paleosols represents the bulk composi-tion of the plants living in the soil and it is assumed that carbonatesrepresent the same value, pedogenic carbonates only represent thecarbonate growth season, which has been identified in some modernsoils as warm, dry periods during the year (Breecker et al., 2009).Some plants have also been shown to exhibit seasonal changes inisotopic composition due to water stress from seasonal changes inprecipitation. Plants tend to produce biomass more enriched in 13Cduring drier periods (Ehleringer, 1993), and it has been documentedthat some plants exhibit annual variation in carbon isotopic composi-tion of biomass based on water stress conditions (Ogaya and Penuelas,2008). Pedogenic carbonates precipitate in the driest months (Breeckeret al., 2009), which means that they may only record plant signatureswhile the plants are producing their heaviest biomass. This may skewresults towards heavier isotopic compositions and inferred higherpercentages of C4 plants.

5. Conclusions

Paleosols that preserve organic matter in the form of black carbonrootlets provide direct evidence of proportions of C3/C4 biomass thatgrew in the soil. Using δ13Corg as a proxy for C3/C4 composition is freeof any potential mixing bias from isotopically heavier atmosphericCO2 because it is in effect sampling the plants that made up theecosystem, rather than sampling the soil gas conditions as in systemsusing pedogenic carbonates. Using preserved organic carbon to recon-struct paleovegetation also offers the advantage of a broader range ofsites and climatic regimes that can be sampled, instead of just paleosolsthat contain pedogenic carbonates indicative of relatively arid condi-tions. This method has been compared to phytolith records, whichproduced paleovegetation reconstructions with very similar amounts ofC4 biomass for this section.

This dataset shows that a small but measurable amount of C4biomass was present in Montana by the late Miocene. Sedimentolog-ical features of the section (large burrows and the lack of forestindicator paleosol types such as Alfisols) as well as the phytolithreconstructions with large percentages of C3 grass morphotypes andlower percentages of C4 morphototypes indicate that this area waslikely a C3 grassland. The modern vegetation for this area in Montanais a mixed C3 grassland and shrubland (USGS, 2006), with a similaramount of C4 grasses to the late Miocene (Epstein et al., 2002). Fossilpollen analyses from the nearby Snake River Plain in Idaho also showthe presence of sagebrush by the middle to late Miocene (Davis andEllis, 2010). Therefore, this paleovegetation reconstruction as well asthat of Davis and Ellis (2010) shows that the modern ecosystem ofsouthwestern Montana was likely already in place by the mid to lateMiocene.

The isotopic data from this study also indicate significant ecologicalvariation both vertically through individual paleosol profiles, andlaterally throughout individual stratigraphic beds. The largest variabil-ity in δ13Corg along a single transect was 4.0‰, which translates toroughly a 30% difference in C4 biomass throughout the ecosystem,depending on where the sample is collected. This variation shows thatlike modern ecosystems, paleovegetation varies greatly spatially. Sincethe climate is thought to be relatively stable during this short timeperiod, the ecosystem patchiness is most likely variation in vegetationin one location through time, probably caused by the migration of riverchannels or changes in the local fire regime. If C4 photosynthesisevolved to fill a certain ecological niche, and that niche migratedthrough time, then it is possible that the δ13Corg of the paleosols could

represent lateral migration of 100% C4 vegetation and that the soils arerecording a time-averaged signal with one small portion of time whenthe vegetation is predominately C4. However, it is more likely that a fewC4 plants were dispersed in a C3 dominated system, much like themodern vegetation of southwestern Montana. These data demonstratethat more than one sample must be analyzed at each bed and alongstrike to present an accurate assessment of the ecosystem and C4abundance for a particular time period.

Acknowledgements

The authors would like to acknowledge Lora Wingate for per-forming isotopic analyses, Michael Hren for discussion about the dataduring early stages of the manuscript and Caroline Strömberg for thephytolith analyses and photographs. The authors would also like tothank Editor Finn Surlyk and an anonymous reviewer for suggestionsthat improved the manuscript. JMC would like to thank Lauren Millerand Alex Dutchak for field assistance. The authors would like toacknowledge support from NSF Grant 1024535 to NDS.

Appendix A. Supplementary data

Supplementary data to this article can be found online at doi:10.1016/j.palaeo.2012.03.035.

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