Geological Society of America Bulletin - esci.umn.edu · Geological Society of America Bulletin doi: 10.1130/B30402.1 Geological Society of America Bulletin 2012;124, no. 3-4;431-443

Geological Society of America Bulletin

doi: 10.1130/B30402.1 2012;124, no. 3-4;431-443Geological Society of America Bulletin

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

-dominated grasslands4isotopes and paleoclimate during the evolution of Cthe Neogene in the southern Great Plains, southwest Kansas, USA: Oxygen 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:

Oxygen isotopes and paleoclimate during the evolution of C

4-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. 431–443; doi: 10.1130/B30402.1; 5 fi gures; 4 tables; Data Repository item 2011261.

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

ABSTRACT

Neogene strata in the Meade Basin (south-west Kansas) preserve numerous superposed calcareous paleosols in sections that range in age from the Clarendonian North American Land Mammal Age (NALMA; 12.0–9.0 Ma, early late Miocene) to the early Irvingtonian NALMA (ca. 2.5–ca. 1 Ma, early Pleistocene). The carbon isotope compositions (δ13C rela-tive to Vienna Peedee belemnite [VPDB]) of pedogenic carbonates from these sections record the protracted regional increase in the abundance of grasses using the C4 pho-tosynthetic pathway over this interval, with the fi rst appearance of a modern-like grass-land ecosystem having >70% C4 biomass around 1.3 Ma. We use the stable oxygen isotope composition (δ18O relative to Vienna standard mean ocean water [VSMOW]) of 194 paleosol carbonates from 19 measured sections to reconstruct the climatic condi-tions in the Meade Basin during the rise of C4 grasses to ecological dominance in the region. Pedogenic carbonate δ18O values are sensitive to soil temperature and the δ18O of soil water, and hence they are a paleoclimate proxy. Carbonate δ13C values do not exhibit consistent trends in relation to δ18O values, indicating no consistent relationship between short-term climatic conditions in terms of temperature or aridity and the abundance of C4 biomass. Mean carbonate δ18O values within biostratigraphic intervals decrease

from the Clarendonian (25.3‰ ± 0.72‰) to the early and middle Blancan (21.8‰ ± 0.87‰ and 22.1‰ ± 0.69‰, respectively), and they also exhibit a decreasing trend in the late Blancan–early Irvingtonian, from ~25‰ to ~21‰. The increase in δ18O values between the end of the middle Blancan and the begin-ning of the late Blancan–early Irvingtonian sections could refl ect the onset of North-ern Hemisphere glaciation and a change in the isotope composition of the hydrosphere due to increased ice volume. The long-term trend in δ18O values suggests that C4 biomass increased coincident with some combination of decreasing temperature, increasing pro-portion of winter precipitation recharge of soil water, and/or increasing soil moisture. Thus, the ecological dominance of C4 grasses in the region today does not seem to be linked to warmer temperatures or increased aridity.

INTRODUCTION

The carbon isotope composition (δ13C) of pedogenic carbonates in Neogene sands and silts near the town of Meade in southwest Kan-sas (Fig. 1) records the rise to ecological domi-nance in the southern Great Plains of plants using the C

4 photosynthetic pathway (warm-

growing-season grasses that dominate tropical and subtropical grasslands, most sedges, a few dicots; Still et al., 2003; Sage, 2004; Edwards and Smith, 2010; Edwards et al., 2010) relative to plants that use the C

3 pathway (trees, shrubs,

and cool-growing-season grasses that dominate high-latitude and high-altitude grasslands; Still et al., 2003; Edwards and Smith, 2010; Edwards et al., 2010). Here, we discuss the basic stratig-raphy, sampling strategy, analytical methods, and δ13C values of pedogenic carbonates in the area of Meade, Kansas, which are interpreted in detail by Fox et al. (2011) in a companion paper in this issue. In brief, the sections sam-pled range in age from early late Miocene to early Pleistocene based on detailed small mam-mal biostratigraphy (Martin et al., 2000, 2008; Honey et al., 2005), magnetostratigraphy for some sections (Lindsay et al., 1975; Martin et al., 2008), and several interbedded ashes in the Pleistocene sections (Izett et al., 1981; Izett and Honey, 1995; Lanphere et al., 2002). The carbon isotope record indicates that the evolution of the modern, C

4-dominated grassland ecosystem of

that region of the southern Great Plains was pro-tracted and took place in several stages (Fig. 2; Fox and Koch, 2003, 2004; Martin et al., 2008; Fox et al., 2011).

The Miocene sections in the Meade area suggest a local increase in the abundance of C

4 grasses from the early late Miocene (Clar-

endonian North American Land Mammal Age [NALMA], 12.0–9.0 Ma, 17.3% ± 6.1%, all means reported ±1 standard deviation) to the late late Miocene (Hemphillian NALMA, 9.0–4.9 Ma, 25.8% ± 6.1%). However, the estimated percent C

4 biomass for both late Miocene sec-

tions in southwest Kansas is similar to estimates for similar-aged carbonates elsewhere in the



Fox et al.

432 Geological Society of America Bulletin, March/April 2012

Great Plains, which does not suggest a regional increase in the abundance of C

4 biomass (Clar-

endonian: 22.1% ± 5.1%; Hemphillian: 23.9% ± 5.4%; Fox and Koch, 2003, 2004; Fox et al., 2011). Moreover, if the two late Miocene sec-tions are combined, the mean percent C

4 bio-

mass (21.2% ± 6.3%) is statistically indistin-guishable from both the mean of the combined Clarendonian and Hemphillian data from other sections in the Great Plains (23.6% ± 5.1%) and the mean of the total Miocene data set from the Great Plains (23.5% ± 3.0%; Fox and Koch, 2003, 2004; Fox et al., 2011).

Between the end of the Hemphillian section and the beginning of the early Blancan NALMA (early Pliocene, 4.9–3.0 Ma), the abundance of C

4 biomass in the Meade area increased, and

the mean percent C4 biomass for the early

Blancan (37.6% ± 6.1%) is statistically greater than the means for both the Hemphillian and

the combined Clarendonian-Hemphillian data. Mean abundance of C

4 biomass in the region

does not change from the early to the middle Blancan (3.0–ca. 2.5 Ma; 38.3% ± 7.4%). During the late Blancan to early Irvingtonian NALMA (ca. 2.5–ca. 1.0 Ma), the abundance of C

4 biomass steadily increases up section

from initial values close to the mean for the early and middle Blancan to within the mod-ern range of C

4 abundance in the region today

(78% ± 10.8%) by ca. 1.3 Ma (~17.5 m in the late Blancan–early Irvingtonian section in Fig. 2). During the late Blancan–early Irving-tonian interval, local extinction, immigration, and in situ evolutionary changes among small mammals led to the appearance of small mam-mal faunas of modern taxonomic composition (Martin et al., 2008). Thus, the fi rst appear-ance in the area of an ecosystem like the mod-ern, C

4-dominated grassland appears to have

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.

occurred by 1.3 Ma in the Meade area (Fig. 2; Fox et al., 2011), but the increase to ecologi-cal dominance of C

4 grasses in the region was

protracted over the late Neogene and did not follow a simple monotonic pattern.

In this paper, we examine the paleoclimatic context of the long-term increase in the abun-dance of C

4 grasses and the appearance of the

modern ecosystem in the region using the δ18O values of the pedogenic carbonates from the Meade area. Paleosol carbonate δ18O values are sensitive to changes in temperature and aridity and so provide a means to determine the role of climatic change in the rise to ecological domi-nance of C

4 grasses in southwest Kansas.

Figure 2. Composite stratigraphic section for pedogenic carbonate samples around Meade, Kansas, with measured δ13C and δ18O values (left, open and fi lled symbols, respectively) and estimated percent C4 biomass (right) in each biostratigraphic interval. Lines through δ18O values in each interval are reduced major axis (RMA) regressions (see Table 2 for parameters). 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 Fox et al., 2011), adjusted for estimated δ13C of atmospheric CO2 for the interval (see Fig. 3B in Fox et al., 2011); short dashed line indicates arid C3 end member based on enrichment of Passey et al. (2002) and estimated δ13C of C3 bio-mass for each interval (see Fig. 3C and 3D in Fox et al., 2011). 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 percen-tiles (tips of left and right whiskers, respec-tively) of δ13C values of Miocene carbonates from across the Great Plains from Fox and Koch (2003, 2004). Percent C4 estimated biomass in each interval was calculated by mass balance based on measured δ13C val-ues and estimated end-member values (see explanation of calculation and Fig. 3 in Fox et al., 2011); heavy dashed line and light-gray box indicate mean modern abundance of C4 biomass in the region ±1 standard deviation (see Table DR2 [see footnote 1]). VSMOW—Vienna standard mean ocean water; VPDB—Vienna Peedee belemnite.



Paleoclimate and C4 biomass in the Neogene of the Great Plains

Geological Society of America Bulletin, March/April 2012 433



Fox et al.


OXYGEN ISOTOPES IN PEDOGENIC CARBONATES

Pedogenic carbonate forms in oxygen isotope equilibrium with soil water (Cerling and Quade, 1993), and the δ18O value of soil carbonate is a function of the temperature-dependent fraction-ation of oxygen isotopes during carbonate pre-cipitation and soil water oxygen isotope com-position (O’Neil et al., 1969; Kim and O’Neil, 1997). Soil temperature at shallow depth in a soil varies seasonally and is generally warmer than air temperature throughout the year (Brady, 1990; Flerchinger and Pierson, 1997). At greater depth, soil temperatures are generally warmer than air temperature during the cool season and cooler than air temperature during the warm sea-son (Brady, 1990). Consequently, mean annual soil temperature at depth is generally within a few degrees Celsius of mean annual air tempera-

ture (e.g., Cerling and Quade, 1993; Breecker et al., 2009). Soil water is recharged by local mete-oric water. Shallow soil water oxygen isotope composition refl ects seasonal and storm-event variation in meteoric water δ18O values but is generally enriched in 18O relative to meteoric water by evaporation from the soil surface prior to infi ltration deeper into the soil; evaporative enrichment is most extreme during warm and/or dry seasons, increasing the variability in shallow soil water δ18O values (Hsieh et al., 1998; Liu et al., 1995; Mathieu and Bariac, 1996; Gazis and Feng, 2004). Soil water at depth (i.e., below ~50 cm) is a temporally integrated mixture of direct meteoric recharge through macropores, evaporatively enriched shallow soil water infi l-trating by various routes, and antecedent water derived from earlier episodes of deep-water recharge that is generally not subject to direct evaporative enrichment in 18O (Mathieu and

Bariac, 1996; Gazis and Feng, 2004). Although deep soil water δ18O values can vary seasonally or episodically, deep soil water typically has δ18O values that are intermediate between the seasonal extremes of meteoric water because it is a mixture of water derived from local mete-oric water (Ferretti et al., 2003; Gazis and Feng, 2004). The precipitation amount weighted mean annual δ18O value of local meteoric water is a lower bound on the δ18O of deep soil water and represents soil water that has not experienced evaporative enrichment.

Pedogenic carbonate δ18O values vary inversely with soil temperature and directly with soil water δ18O (O’Neil et al., 1969; Kim and O’Neil, 1997; Fig. 3A). Thus, for a constant soil temperature, a secular decrease in soil water δ18O value will lead to pedogenic carbonates with lower δ18O values. Conversely, for constant soil water δ18O values, a secular decrease in

δ18Occ = 10‰

δ18Occ = 20‰

δ18Occ = 30‰

early Blancan δ18Occ = 21.8 ± 0.84‰Clarendonian δ18Occ = 25.3 ± 0.72‰

-25

-20

-15

-10

-5

0

5

5 10 15 20 25 30δ18

O s

oil

wat

er (

‰, V

SM

OW

)Temperature (°C)

B-20

-15

-10

-5

0

5

10

15

20

25

30

0 5 10 15 20 25

Met

eoric

wat

erC

arbo

nate

δ18O

(‰, V

SM

OW

)

Temperature (°C)

A

δ18Osw = -20‰

δ18Osw = -10‰

δ18Osw = -15‰

δ18Osw = mw

Figure 3. Relationships among temperature, δ18O of water, and δ18O of carbonate. (A) Thin dashed lines are δ18O values of carbonate calculated for different temperatures at the indicated δ18O values of soil water using the temperature-dependent fractionation of Kim and O’Neil (1997). Upper two thin dashed lines correspond to δ18Osw values of −5‰ and 0‰. For constant temperature, carbonate δ18O decreases with decreasing soil water δ18O; for constant soil water δ18O, carbonate δ18O increases with decreasing temperature. Open gray circles are amount weighted mean annual δ18O of meteoric water and mean annual temperature for 133 stations between 20°N and 60°N and S latitude in the International Atomic Energy Association’s Global Network of Isotopes in Precipitation (GNIP) through 1997 (http://www-naweb.iaea.org/napc/ih/GNIP/IHS_GNIP.html). Black star indicates amount weighted mean annual δ18Omw value for Lake Scott State Park, Kansas, from Vachon (1996) and mean annual temperature for 1895–2008 for Scott City, Kansas (National Climatic Data Center Summary of the Day, www.ncdc.noaa.gov/oa/ncdc.html). Thick solid line is least squares lin-ear regression through the GNIP data (δ18O = 0.43T −13.0; R2 = 0.70). Thick dashed line is carbonate δ18O value calculated for differ-ent temperatures and water δ18O values that are dependent on temperature according to the regression for the GNIP data; carbonate δ18O values increase with temperature due to the temperature dependence of δ18Omw values. (B) Dashed lines indicate possible com-binations of soil temperature and soil water δ18O values for carbonate δ18O values of 10‰, 20‰, and 30‰. Solid lines are solutions for mean δ18O values of Clarendonian and early Blancan carbonates (gray envelopes indicate solutions for ±1 standard deviation).





soil temperature will lead to pedogenic carbon-ates with higher δ18O values. However, amount weighted mean annual δ18O values of modern meteoric waters at middle to high latitudes have a strong positive correlation with local air temperatures (Fig. 3A; Dansgaard, 1964; Rozanski et al., 1993), although the slope and intercept of this relationship were likely differ-ent in the past (Boyle, 1997). Additionally, in continental interiors, the seasonal variation in meteoric water δ18O values strongly correlates with air temperatures (Rozanski et al., 1993). Thus, on seasonal to geological time scales, the δ18O value of pedogenic carbonate should actually vary directly with temperature through the relationship between temperature and the δ18O value of meteoric water (Fig. 3A). A given measured δ18O value of pedogenic carbonate is consistent with formation over a range of com-binations of soil temperature and water compo-sition (Fig. 3B). However, because the effects on carbonate δ18O values of changes over geo-logical time in both evaporative enrichment of soil water and soil temperature are impossible to deconvolve without additional information, we cannot make quantitative estimates of either meteoric water composition or temperature from our paleosol carbonate data. We do exam-ine the sensitivity of carbonate δ18O values to changes in temperature, seasonal distribution of precipitation, and moisture availability, but we only interpret changes in carbonate δ18O val-ues in terms of cooler/wetter and warmer/drier. Thus, a decrease (increase) through time in car-bonate δ18O values could refl ect (1) a decrease (increase) in mean annual temperature, and hence a decrease (increase) in the δ18O value of meteoric water and soil water; (2) a decrease (increase) in aridity and net evaporative fl ux, and hence a reduction (increase) in evaporative enrichment of soil water in 18O; (3) a relative increase (decrease) in the contribution of iso-topically light winter precipitation to soil water recharge, and hence a decrease (increase) in the δ18O value of deep soil water; or (4) a combina-tion of some or all of these changes.

The general relationship between aridity and δ18O value of paleosol carbonate (through the evaporative enrichment of soil water in 18O) provides a means by which we can constrain the infl uence of aridity on the carbon isotope composition of paleosol carbonates (see dis-cussion and references in Fox et al., 2011). An increase in aridity can increase the δ13C value of paleosol carbonate without any change in the relative abundance or even presence of C

4

biomass, given a decrease in apparent carbon

isotope enrichment in C3 plants in response to

water stress. For stratigraphic sections in which paleosol carbonate δ13C values increase up sec-tion but remain within the range of δ13C values for C

3 ecosystems, as discussed in Fox et al.

(2011) for some individual measured sections, aridifi cation predicts carbonate δ18O and δ13C values that are strongly correlated and a linear regression with δ18O value as the independent (or x) variable that is statistically signifi cant, has a positive slope, and explains a large fraction of the variance in δ13C values. The same statistical relationships between δ18O and δ13C values for sections in which δ13C values are high enough to require C

4 biomass and δ13C values increase

up section could indicate either a change in fractionation by a constant proportion of C

3 bio-

mass or an increase in C4 biomass in response to

increased aridity (which increases photorespira-tion rate and therefore favors C

4 photosynthesis;

Sage, 2004). Under the common assumption that the Neogene increase in the abundance of C

4 grasses was a response to increasing aridity

(e.g., Axelrod, 1985), we also expect a long-term relationship between δ13C and δ18O values of pedogenic carbonates that should also yield a linear regression with δ18O value as the indepen-dent variable that is statistically signifi cant, has a positive slope, and explains a large fraction of the variance in δ13C values.

Timing of Carbonate Precipitation

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, Breecker et al. (2009) argued that Holocene paleosol carbonates that developed on fl ood-plain gravels in the Chihuahuan Desert, New Mexico, indicate that soil carbonate isotopic records are seasonally biased to warm and dry conditions when net evaporation of soil water drives carbonate precipitation. Aspects of the climate regime in the Great Plains are broadly similar to those in New Mexico where Breecker et al. (2009) studied young soil carbonates, in particular, the coincidence of peak rainfall and the warm season (for Meade, Kansas, on average, 66% of annual precipitation falls dur-ing the warm months of May–September; see Fig. DR11). Thus, it is likely that at least modern carbonates in the Meade area may also be biased to the warm months, if not also the Neogene car-bonates we sampled.

To explore the consequences of different assumptions about the timing of carbonate pre-

cipitation (i.e., mean annual conditions, inte-grated monthly conditions, or warm season biased), in the discussion here we use long-term meteorological data from 1895 to 2008 for Scott City, Kansas (~160 km north-northwest of Meade) and measured meteoric water oxygen isotope composition (δ18O

mw) for nearby Lake

Scott State Park from Vachon (1996; Fig. DR1, parts B and D–F [see footnote 1]) to estimate carbonate δ18O values based on the temperature-dependent calcite-water fractionation factor of Kim and O’Neil (1997). Long-term mean val-ues and seasonal variation for both temperature and rainfall at Scott City are similar to those for Meade (see Fig. DR1, parts A–D [see foot-note 1]). Although mean annual temperature is 1.5 °C higher, mean annual rainfall is quite similar (489 ± 142 mm for Scott City and 545 ± 132 mm for Meade), and 67% of annual rainfall at Scott City occurs during the warm months of May–September. However, δ18O values of mete-oric water are not available for Meade.

We used six models of temperature and water composition to estimate mean annual carbonate δ18O values calculated various ways for Meade today based on the weather station and δ18O

mw

data from Scott City (Table 1): (1) mean annual temperature and mean annual δ18O

mw; (2) mean

annual temperature and weighted mean annual δ18O

mw value based on the monthly amount of

rainfall; (3–4) mean monthly temperature and monthly δ18O

mw value; and (5–6) mean warm

season (May–September) temperature for all carbonate precipitation and monthly δ18O

mw

value. In each model, air temperature and δ18O

mw values were used as the soil temperature

and soil water δ18O value. In models 1 and 2, assumed soil temperature was used to calculate the temperature-dependent calcite-water frac-tionation factor (α) of Kim and O’Neil (1997):

1000 ln α = 18.03(103T–1 – 32.42), (1)

where T is in Kelvin. Equation 1 was then used to calculate an estimated mean annual carbonate δ18O value (δ18O

CO3) based on the assumed soil

water composition (δ18Osw) and α:

δ18OCO3

= (1000 + δ18OSW

) × α −1000. (2)

In models 3–6, we fi rst calculated a monthly carbonate δ18O value based on the temperature-dependent fractionation factor α (Eq. 1) and the model of soil water composition. For these models, predicted annual carbonate was then calculated either as the mean of the monthly carbonate δ18O values calculated from Equation

1GSA Data Repository item 2011261, Stable isotope data from pedogenic carbonates in measured sections, is available at http://www.geosociety.org/pubs/ft2011.htm or by request to [email protected].



Fox et al.


2 for monthly data (3 and 5) or as the weighted mean monthly carbonate δ18O values calculated from Equation 2 for monthly data and weighted by the monthly precipitation amount (4 and 6).

Models 1–4 represent different versions of the standard interpretation of soil carbonates as representing annual conditions. Models 1 and 2 do not account for seasonal variation in temperature or the amount and δ18O value of meteoric (i.e., soil) water. Model 3 assumes that carbonate forms incrementally throughout the year in equal amounts per month. Carbon-ate formation throughout the year is reasonable for the record in Meade, which includes frost-intolerant tortoises up through the Irvingtonian (e.g., Butler Spring fauna; Schultz, 1969), but it is not consistent with modern winter monthly temperatures near or below 0 °C for both Scott City and Meade. Model 3 also does not take into account seasonal variation in the amount of meteoric water. By weighting the mean annual contribution of carbonate for each month by the amount of precipitation in each month, model 4 is the most realistic of the models based on annual conditions and is one of the two pre-ferred models we use for interpretation. Models 5 and 6 both assume that carbonate formation is biased to the warm months of the year, in accord with the results of Breecker at al. (2009) for the Chihuahuan Desert. Of the two, model 6 is more realistic because it accounts for monthly varia-tion in the amount and δ18O of meteoric water, and so it is the other preferred model for inter-pretation of the Neogene data.

Estimated carbonate δ18O values assume insignifi cant contribution of water derived from snow to soil water, as is the case for grassland soils in northeastern Colorado (Ferretti et al., 2003). Because winter precipitation is mostly snow and has low δ18O values, inclusion of snow melt into soil water would lower the mean annual δ18O value of carbonate regardless of the model assumptions. If we use total meteoric

water (precipitation + snow) for Scott Lake and the assumptions of model 4, estimated carbonate δ18O decreases to 20.2‰; for the assumptions of model 6, estimated carbonate δ18O decreases to 17.3‰ (Table 1). By using meteoric water δ18O values for soil water, the estimated carbonate values also assume no evaporative enrichment of soil water in 18O. Allowing soil water to be enriched in 18O relative to meteoric water would lead to estimated carbonate with higher δ18O values, but no data are available for the area to constrain the amount of evaporative enrichment.

METHODS

Both fi eld and laboratory methods are described in Fox et al. (2011). In brief, 194 fi eld samples of stage II (nodular) and stage III (coalesced nodules to caliche) carbonates (Gile et al., 1966; Machette, 1985) were collected from 19 measured sections in the Meade Basin fi eld area (Fig. 1). Individual measured sections ranged in length from 3 to 38 m. Petrographic thin sections of 55 fi eld samples indicated no secondary (e.g., crack-fi lling) sparry calcite in any of the samples. Field samples were cut on a lapidary wheel, rinsed, dried, and then sampled with a 0.5 mm diamond bur in a handheld rotary drill. As discussed in Fox et al. (2011), replicate samples of 108 fi eld samples exhibited low vari-ability for both δ13C and δ18O values, so those analyses are reported here as midpoint values, and the remaining 86 fi eld samples were not analyzed in replicate. In total, we measured the carbon and oxygen isotope composition of 194 fi eld samples in 302 unique stable isotope analy-ses. Sample powders were roasted in vacuo at 400 °C for at least 1 h to combust any organic matter and to remove water. The majority of samples were analyzed at the University of Min-nesota 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 CO2

was measured using a Finnigan MAT 252 gas-source isotope-ratio mass spectrometer. Preci-sion 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 normalized to the compo-sition of NBS-19. Samples from three sections (Aries, Borchers Badlands north of Kansas Hwy 23, Alien Canyon) were analyzed at the Univer-sity of California–Santa Cruz, where samples were reacted with phosphoric acid at 90 °C in a Micromass Isocarb automatic carbonate extrac-tion system, and the isotopic composition of the resulting CO

2 was measured on either a Micro-

mass Prism or Optima isotope-ratio mass spec-trometer. These samples were normalized to the composition of Carrara marble laboratory stan-dard, and the normalization was checked against the corrected 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 labo-ratory and NBS standards at both laboratories. Five fi eld samples from the Borchers section were analyzed in both laboratories (discussed in Fox et al., 2011); reproducibility of δ18O values between laboratories for these fi ve samples was 0.0‰, 0.1‰, 0.2‰, 0.5‰, and 0.5‰.

To analyze stratigraphic trends in the δ18O values within biostratigraphic intervals and to examine the relationship between δ13C values (discussed in Fox et al., 2011) and δ18O val-ues (this study), we use reduced major axis

TABLE 1. PARAMETERS FOR DIFFERENT MODELS OF CARBONATE δ18O VALUES

Model Soil temperature Soil water δ18O δ18OCO3 represents Predicted δ18OCO3

(‰)1 Mean annual Mean annual Mean annual 20.42 Mean annual Monthly precipitation amount weighted mean annual Mean annual 22.23 Mean monthly Monthly Mean of monthly δ18OCO3 values 20.54 Mean monthly Monthly Mean of precipitation amount weighted monthly δ18OCO3 values 21.25 Mean warm season Monthly Warm season–biased mean monthly δ18OCO3 values 18.36 Mean warm season Monthly Warm season–biased, precipitation amount weighted monthly

δ18OCO3 values20.1

4(total mw)

Mean monthly Monthly Mean of precipitation amount weighted monthly δ18OCO3 values 20.2

6(total mw)

Mean warm season Monthly Warm season–biased, precipitation amount weighted monthly δ18OCO3 values

17.3

Note: Based on meteorological data for Scott City, Kansas (www.ncdc.noaa.gov/oa/ncdc.html) and measured meteoric water δ18O values for Lake Scott State Park, Kansas (Vachon, 1996). Soil temperature is based on air temperature from the weather station data. Soil-water δ18O is based on meteoric water δ18O values. Predicted carbonate δ18O values were calculated using temperature-dependent calcite-water fractionation factors of Kim and O’Neil (1997). Models 2 and 4–6 use only rainfall amounts in calculating amount weighted means; models 4 (total mw) and 6 (total mw) also include snowfall amounts.





(RMA) linear regressions calculated based on the linear method of Sokal and Rohlf (1994) using RMA for Java 1.21 (Bohonak and van der Linde, 2004). Slopes of the RMA regressions for which the 99% confi dence interval (CI) does not include 0 are considered statistically distinct from 0, implying either a statistically signifi cant stratigraphic trend in δ18O values or statistically signifi cant relationship between δ13C and δ18O values. As with regression analyses in Fox et al. (2011), results for the more familiar and com-mon ordinary least squares linear regressions are also presented in the supplemental material (see footnote 1). All other statistical tests (t-test, Mann-Whitney U-test, analysis of variance [ANOVA] with post hoc Scheffé test for mul-tiple comparisons) were done using SPSS 13.0 for Macintosh.

RESULTS

All δ13C and δ18O values of carbonates from Meade are presented in Table DR1 (see foot-note 1).

The Clarendonian (early late Miocene, 12–9 Ma) section in the Ogallala Formation below the Borchers Badlands has a mean δ18O value of 25.3‰ ± 0.72‰, and the δ18O values range from 24.3‰ to 27.0‰ (Fig. 2; Table 2). The slope of the RMA regression of the Claren-donian data with meter level as the independent variable is statistically signifi cant and indicates that δ18O values decrease by 0.22‰ per meter up section (Table 3).

The mean δ18O value for the Hemphillian (late late Miocene, 9–4.9 Ma) High Banks section of the Ogallala Formation (24.8‰ ± 0.13‰) is slightly lower than for the Claren-donian Ogallala section below Borchers (Fig. 2; Table 2). The Hemphillian data exhibit low vari-ability, and both the standard deviation and the range (24.4‰–25.0‰) are lower than for the Clarendonian data. The δ18O values at High Banks do not exhibit a statistically signifi cant stratigraphic trend (Table 3).

The δ18O values in the early Blancan (early Pliocene, 4.9–3.0 Ma) section are lower than and do not overlap with the late Miocene sec-tions (Fig. 2; Table 2). The mean δ18O value in the early Blancan is 21.8‰ ± 0.87‰, and the range is from 19.9‰ to 23.7‰. RMA regres-sion indicates that the δ18O values have a statis-tically signifi cant trend to higher values through the section, although the proportion of variance explained is negligible (R2 < 0.01; Table 3).

The δ18O values for the middle Blancan (mid-Pliocene, 3.0–2.5 Ma) sections are similar to those for the early Blancan, as was the case with the δ13C values (Fox et al., 2011; Fig. 2; Table 2). The middle Blancan mean δ18O value (22.1‰ ± 0.69‰) is almost identical to that for the early Blancan, and the range for the middle Blancan (20.8‰–23.7‰) overlaps that for the early Blancan. Middle Blancan δ18O values also increase signifi cantly up section (Table 3), but, as with the early Blancan data, meter level explains a negligible fraction of the variance in δ18O values (R2 = 0.01).

TABLE 2. SUMMARY STATISTICS OF δ18O VALUES FOR EACH BIOSTRATIGRAPHIC INTERVAL

δ18O(VSMOW)

Mean SD N Max Min Range

Late Blancan–Early Irvingtonian All data 23.3 2.48 39 31.0 20.9 10.1Without outliers 22.7 1.37 36 25.5 20.9 4.6

Middle Blancan 22.1 0.69 27 23.7 20.8 2.9Early Blancan 21.8 0.87 106 23.7 19.9 3.8Hemphillian 24.8 0.18 10 25.0 24.4 0.6Clarendonian 25.3 0.72 12 27.0 24.3 2.7Miocene combined 25.1 0.59 22 27.0 24.3 2.7All 22.5 1.71 194 31.0 19.9 11.1

Note: Data for the Late Blanca–Early Irvingtonian are reported both with and without three outliers that have δ18O values >29‰.

TABLE 3. REGRESSION STATISTICS FOR REDUCED MAJOR AXIS (RMA) REGRESSIONS WITH METER LEVEL OF SAMPLES IN COMPOSITE SECTIONS FOR EACH BIOSTRATIGRAPHIC INTERVAL AS X

VARIABLE AND δ18O VALUE OF PALEOSOL CARBONATES AS Y VARIABLE

Slope S.e. 99% CI Intercept R 2

Late Blancan–early Irvingtonian All data –0.37 0.046 –0.49, –0.24 28.5 0.43Without outliers –0.20 0.019 –0.25, –0.15 28.5 0.69

Middle Blancan 0.45 0.089 0.20, 0.69 20.2 0.01Early Blancan 0.06 0.005 0.04, 0.07 19.6 <0.01Hemphillian –0.08 0.027 –0.17, 0.01 25.0 0.01Clarendonian –0.21 0.052 –0.37, 0.07 26.8 0.39

Note: Regressions and confi dence intervals 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; CI—confi dence intervals. Bold indicates regressions for which the 99% confi dence interval for the slope does not include zero.

The late Blancan–early Irvingtonian (late Pliocene–early Pleistocene, 2.5–ca. 1.0 Ma) sections have three samples with unusually high δ18O values (i.e., >29‰; Fig. 2). The fi eld sam-ples are outwardly similar to the other nodular carbonates from that part of the section, the sed-iments from which they were collected are not different than the base of that section or the base of the section below the Borchers fauna, and the δ13C values of these samples are not outliers in the section (Fig. 2). The high δ18O values pre-sumably represent either a brief, rather extreme change in local soil hydrology or a diagenetic event localized to only that part of the section that did not alter the δ13C values. The former seems more likely, but this would require con-siderable evaporative enrichment of soil water in 18O, and we otherwise have no evidence for any environmental changes of such an extreme nature. However, Breecker et al. (2009) reported values this high for shallow carbonates from dif-ferent habitats in the Chihuahuan Desert, so it is plausible that evaporative enrichment of soil water could lead to such high carbonate δ18O val-ues. Including the three samples with high δ18O values, the mean δ18O value is 23.3‰ ± 2.48‰; without these three samples, the mean and stan-dard deviation are considerably lower (22.7‰ ± 1.37‰; Table 2). The δ18O values decrease through the section by almost 8‰ considering all of the data and ~4‰ without the three outli-ers with unusually high δ18O values (>29‰). If the outliers are included, the δ18O values have a statistically signifi cant negative trend relative to meter level (−0.37‰ per meter; Table 3). If the three outlier values are excluded, the slope of the RMA regression is lower but still statis-tically signifi cant (Table 3), and stratigraphy explains an even larger fraction of the variance in δ18O values in this section (R2 = 0.69). If the Aries and Borchers north section are correlated on the Cerro Toledo B ash, RMA regressions for the δ18O values are unaffected and remain statistically signifi cant whether the outliers with δ18O values >29‰ are included (δ18O = –0.37 × meters + 28.5, R2 = 0.43, 99% CI for the slope: −0.49 to −0.24) or excluded (δ18O = –0.20 × meters + 25.7, R2 = 0.69, 99% CI for the slope: −0.26 to −0.15). Thus, the stratigraphic trend in the δ18O values is also insensitive to the correla-tion of the Aries and Borchers sections.

DISCUSSION

The staged increase in the abundance of C4

biomass from the late Miocene to the early Pleistocene in the Meade Basin (Fig. 2; Fox et al., 2011) seems paradoxical given the long-term record of global cooling during this inter-val from δ18O values of benthic foraminifera



Fox et al.


(Zachos et al., 2001), assuming that C4 abun-

dance in a region responds at least in part to warmer conditions, as is often assumed, and that local or regional conditions broadly follow aver-age global climatic conditions. To examine the role of local climate change in the increase in C

4

biomass in the Meade Basin, we fi rst examine the long-term climate record of paleosol carbon-ate δ18O values, then compare these to the record of increasing C

4 biomass from δ13C values, and

fi nally compare the measured δ18O values to the estimated δ18O values of modern carbonate in the region based on the meteorological and isotopic data from Scott City, Kansas (Fig. DR1 [see footnote 1]). As discussed previously, the oxygen isotope composition of paleosol car-bonate is sensitive to air temperature and soil water composition, which refl ects meteoric water composition modifi ed by evaporation to a greater or lesser degree. In general, lower (higher) carbonate δ18O values indicate lower (higher) temperatures through lower (higher) δ18O values of meteoric water, or a greater fraction of winter (summer) season recharge of soil water, or wetter (drier) conditions and a decrease (an increase) in evaporative enrich-ment of soil water in 18O, or a combination of some or all of these.

Long-Term Trend in Pedogenic Carbonate δ18O Values in the Meade Basin Record

One-way ANOVAs of the δ18O values within intervals indicate statistically signifi cant differ-ences in mean values among intervals (results of statistical tests other than RMA regressions are discussed in the text and presented in full in the supplemental materials [see footnote 1]), whether or not the unusually high δ18O values (>29‰) in the late Blancan–early Irvingtonian are included or the Clarendonian and Hemp-hillian are combined (computed with the late Blancan outliers having δ18O >29‰ excluded). To examine the long-term trend in values in the Meade Basin, we used both the post hoc pair-wise Scheffé test for multiple comparisons (Table DR3 [see footnote 1]) and pair-wise Mann-Whitney U-tests (Table DR4 [see foot-note 1]) to identify the intervals in the Meade Basin that have statistically signifi cant differ-ences in mean δ18O values.

Post hoc Scheffé tests indicate that Clarendo-nian and Hemphillian data are not statistically distinguishable, but the Mann-Whitney U test indicates that the Hemphillian δ18O values are statistically lower than the Clarendonian values (though by only 0.5‰). Both tests indicate that the Miocene δ18O values, whether grouped or treated separately, are statistically higher than those in the later intervals. These results are

in good accord with the pattern of global cool-ing from the middle Miocene to the Pleisto-cene (Zachos et al., 2001), which would yield meteoric water with lower δ18O values in the Great Plains (if regional temperature followed the global pattern recorded by benthic fora-minifera), and soil carbonates with lower δ18O values, all else being equal. Both the early and middle Blancan sections have statistically sig-nifi cant stratigraphic trends toward higher δ18O values (Fig. 2; Table 3), but neither regression explains much of the variance in the δ18O val-ues, and the means for the two intervals are not statistically different (Tables DR3 and DR4 [see footnote 1]). These intervals correspond gener-ally to the Pliocene warm period (Brierly et al., 2009). The similarity of these sections despite the statistical evidence for stratigraphic trends suggests that climatic conditions as recorded by δ18O values were stable in the Meade Basin dur-ing the Pliocene warm period.

If the outliers >29‰ are excluded from the late Blancan–early Irvingtonian interval, the mean δ18O value for the Middle Blancan is not statistically different from that for the late Blancan–early Irvingtonian (mean values are statistically different according to the post hoc Scheffé test if the outliers are included, refl ect-ing the infl uence of the unusually high δ18O values). However, the base of the late Blancan–early Irvingtonian section (Fig. 2) has higher δ18O values (22.9‰–25.4‰) than the top of the middle Blancan section (22.8‰), and the late Blancan–early Irvingtonian data have a strong, statistically signifi cant, stratigraphic trend toward lower values; this section reaches val-ues less than the mean for the middle Blancan (~21‰ vs. 22.7‰) and close to the minimum δ18O value for the middle Blancan (20.8‰). Thus, δ18O values in the Meade Basin increase after the end of the Pliocene warm period, which could potentially refl ect local warming or increased aridity. We do not have precise age control on either the top of the Middle Blancan section or the base of the late Blancan–early Irvingtonian section. However, the interval between the top of the middle Blancan and below the Huckleberry Ridge Ash (2.06 Ma) in the late Blancan–early Irvingtonian corre-sponds to the onset of Northern Hemisphere glaciation beginning at ca. 2.7 Ma (Haug et al., 2005). This interval also corresponds to the age of the fi rst advance of the Laurentide ice sheet into the midcontinent, based on a cosmogenic radionuclide exposure age of 2.47 ± 0.19 Ma for the Whippoorwill Formation, a periglacial soil buried beneath the Atlanta Formation, the lowest glacial till in northern Missouri, ~800 km east and ~1.5° north of the Meade Basin (Rovey and Balco, 2009).

Given the decrease in global temperatures associated with the onset of Northern Hemi-sphere glaciation, the increase in δ18O values between the top of the middle Blancan section (~22‰) and the base of the late Blancan–early Irvingtonian section (~24‰; Fig. 3; Figs. DR5 and DR6 [see footnote 1]) seems paradoxi-cal. Alkenone unsaturation indices and dia-tom δ18O values from Ocean Drilling Program (ODP) Site 882 suggest that late summer–early autumn sea-surface temperatures in the sub-arctic Pacifi c increased at 2.7 Ma, providing a source for increased moisture delivery to boreal North America to feed continental ice-sheet growth (Haug et al., 2005). If Arctic cooling and the onset of glaciation were associated with a defl ection of storm tracks to the north, then arid-ity in the midcontinent could have increased. Additionally, the onset of Northern Hemisphere glaciation would have left water vapor sources enriched in 18O. Thus, the increase in δ18O values between the top of the middle Blancan section and the base of the late Blancan–early Irvingtonian likely refl ects a combination of ice-volume increase and midcontinental aridi-fi cation associated with the onset of glaciation rather than warming. We interpret the steady decrease in δ18O values during the late Blancan–early Irvingtonian as refl ecting some combina-tion of cooling and/or increasing contribution of winter precipitation to soil water after estab-lishment of the Northern Hemisphere glaciation and, presumably, defl ection of storm tracks to the south of the Laurentide ice sheet front.

To examine the role of local climatic variation in the increase of C

4 biomass during the study

interval, we used RMA linear regressions with δ18O values in each interval as the independent variables and δ13C values and percent C

4 bio-

mass as the dependent variables (Fig. 4; Table 4; Table DR5 [see footnote 1]). We also used RMA linear regression between δ18O and δ13C values over the entire study interval to exam-ine the role of long-term climatic change in the overall increase in C

4 biomass (Fig. 5; Table 4;

Table DR5 [see footnote 1]). If increased mean annual or warm season aridity or temperature were an important control on the abundance of C

4 biomass on short and/or long time scales,

we would expect a strong, positive correlation between δ13C and δ18O values and a strong statistical dependence of δ13C values on δ18O values to explain a substantial portion of the variance in δ13C values. Results are essentially identical for regressions with percent C

4 instead

of δ13C values (regressions for percent C4 not

fi gured, but they are presented in Table 4).In general, the regression analyses do not

support a strong or consistent role for climate as recorded by carbonate δ18O values in the





increase in C4 biomass in the Meade Basin. The

regression for the Clarendonian data is statisti-cally signifi cant and explains a relatively large fraction of the variance in δ13C values, but δ13C and δ18O values are negatively correlated as δ13C values increase and δ18O values decrease signif-icantly up section in the Clarendonian section (Table 3). Thus, the Clarendonian data suggest that C

4 biomass increased locally in response to

some combination of cooling, a relative increase in winter precipitation, increased precipitation, and/or decreased evaporative enrichment of soil water in 18O. The slope for the Hemphillian regression cannot be distinguished from 0, and the regression explains a negligible amount of

••

••••••• •••

-10

-8

-6

-4

-2

0

17 19 21 23 25 27 29 31 33

δ13C

(‰

, VP

DB

)

δ18O (‰, VSMOW)

A

δ13C = -1.27 δ18O + 24.4R2 = 0.62

••••••••••-8

-6

-4

-2

0

δ13C

(‰

, VP

DB

)

B

δ13C = -1.7 δ18O + 35.6R2 = 0.01

••• •••

••••

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••••••

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-8

-6

-4

-2

0

δ13C

(‰

, VP

DB

)

Cδ13C = 1.04 δ18O - 27.6R2 = 0.11

δ13C

(‰

, VP

DB

)

-8

-6

-4

-2

0

••••

•

••

••

••• •••••

••

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Dδ13C = 1.60 δ18O - 40.2R2 = 0.22

••••••• •

•

•••

•••••••• ••••••

••••••••••

-8

-6

-4

-2

0

2

δ13C

(‰

, VP

DB

)

Eδ13C = -0.88 δ18O + 17.5R2 = 0.37 (solid line)

variance in δ13C values. For both the early Blan-can and the middle Blancan sections, δ18O and δ13C values are positively correlated, and the regressions are statistically signifi cant, consis-tent with higher temperatures or aridity during the growing season or greater summer recharge of soil water corresponding to higher abun-dance of C

4 biomass. However, in both cases,

δ18O values have a lower range than δ13C val-ues, and neither regression explains much of the variance in δ13C values (11% for the early Blancan and 22% for the middle Blancan). However, δ13C and δ18O values in these sections have stratigraphic trends with signifi cant slopes, low explanatory power, and opposite signs (i.e., δ13C trends both have negative slopes [Fox et al., 2011]; δ18O trends have positive slopes), and the δ18O values have statistically indistinguish-able means. Thus, the regression results for these sections do not indicate an obvious rela-tionship between climatic conditions and the abundance of C

4 biomass. The regression for the

late Blancan–early Irvingtonian section is also statistically signifi cant (whether or not outlier δ18O values are included) and explains more of the variance in δ13C values (37% if outlier δ18O values are excluded), but in contrast to the early and middle Blancan results, the slope of the regression is negative as δ13C values increase signifi cantly and δ18O values decrease signifi -cantly up section.

For all of the data considered together, the regressions are signifi cant for both δ13C values and percent C

4 (i.e., 99% confi dence intervals

Figure 4. Reduced major axis (RMA) linear regressions of individual δ13C values for each biostratigraphic interval on corresponding δ18O values. (A) Clarendonian (early late Miocene, 12.0–9.0 Ma). (B) Hemphillian (late late Miocene, 9.0–4.9 Ma). (C) Early Blancan (early Pliocene, 4.9–3.0 Ma). (D) Middle Blancan (mid-Pliocene, 3.0–2.5 Ma). (E) Late Blancan–early Irvingtonian (late Pliocene–early Pleistocene, 2.5–ca. 1.0 Ma). In E, the solid line and equation represent the regression excluding the three outli-ers with δ18O >29‰ in the Borchers sec-tion north of Kansas Highway 23, and the dashed line is the regression including the outliers (δ13C = –0.48 δ18O + 8.6, R2 = 0.19).

TABLE 4. REGRESSION STATISTICS FOR REDUCED MAJOR AXIS (RMA) REGRESSIONS WITH δ18O VALUES OF PALEOSOL CARBONATES AS INDEPENDENT VARIABLE AND δ13C VALUES OF

PALEOSOL CARBONATES OR PERCENT C4 BIOMASS AS DEPENDENT VARIABLE

δ13C values(‰)

Slope S.e. 99% CI Intercept R2

Late Blancan–early Irvingtonian All data –0.48 0.071 –0.67, –0.28 8.6 0.19Without outliers –0.88 0.120 –1.21, –0.55 17.5 0.37

Middle Blancan 1.60 0.281 0.81, 2.38 –40.2 0.22Early Blancan 1.04 0.096 0.79, 1.29 –27.6 0.11Hemphillian –1.70 0.597 –3.70, 0.31 35.6 0.01Clarendonian –1.27 0.248 –2.05, –0.48 24.4 0.62

All data All data –0.94 0.067 –1.11, –0.76 16.4 0.02Without outliers –1.15 0.081 –1.38, –0.94 21.0 0.07

Interval means All data –1.21 0.546 –4.39, 1.98 23.0 0.39Without outliers –1.18 0.457 –3.85, 1.49 22.2 0.55

Percent C4Late Blancan–Early Irvingtonian All data –3.34 0.483 –4.32, –2.36 132.4 0.22

Without outliers –6.13 0.807 –8.35, –3.91 194.2 0.41Middle Blancan 10.74 1.89 5.48, 15.99 –198.6 0.22Early Blancan 7.01 0.645 5.31, 8.70 –115.1 0.12Hemphillian –11.47 4.048 –25.05, 2.11 310.3 0.004Clarendonian –8.48 1.668 –13.77, –3.19 232.1 0.61All data All data –6.72 0.479 –7.67, –5.78 190.5 0.02

Without outliers –8.23 0.575 –9.78, –6.94 223.4 0.08Interval means All data –9.00 3.985 –32.28, 14.27 245. 0.41

Without outliers –8.88 3.369 –28.56, 10.80 242.0 0.57Note: For the late Blancan–early Irvingtonian data, the fi rst row is the regression for all data, and the second

row is the regression without the three outlier δ18O values >29‰. Regressions and confi dence intervals 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; CI—confi dence intervals. Bold indicates regressions for which 99% confi dence interval does not include 0.



Fox et al.


do not include 0), but the slopes of the regres-sions for all data are negative and generally steep, and none of the regressions explains more than 8% of the variance in δ13C values or per-cent C

4 biomass (Fig. 5A). Although all of these

regressions are weak, they do suggest that the percentage of C

4 biomass increased over the

long-term even as climate became cooler and/or less arid. Given the variability in both δ13C and δ18O values within sections, the inverse rela-tionship between the abundance of C

4 biomass

and carbonate δ18O value is more obvious for the mean δ13C and δ18O values within intervals (Fig. 5B). The regression for interval means has a negative slope, but the 99% confi dence inter-val is wide and includes 0, although the regres-sion is necessarily supported by far fewer data. Nevertheless, this regression explains a rela-tively large proportion of the variance in δ13C values (R2 = 0.55).

The regressions for all of the data (as well as those for the late Blancan–early Irvingto-nian) imply that the abundance of C

4 biomass

increased to modern levels in the Meade Basin in conjunction with some combination of lower temperatures (hence lower δ18O

mw values, coun-

teracting the infl uence of lower temperature on oxygen isotope fractionation during carbonate precipitation), an increase in the proportion of winter recharge of soil water, and increased moisture availability through increased precipi-tation and/or decreased evaporation of soil water. These possible explanations for the trend in δ18O values are evident from the allowable solu-tions to the calcite paleotemperature equation of Kim and O’Neil (1997) for the mean Clarendo-nian and early Blancan δ18O values (Fig. 3B). Thus, the Meade Basin record suggests that the abundance of C

4 biomass in the southern Great

Plains increased as climate became cooler and/or wetter, contrary to common expectations for the response of C

4 biomass to climate. These

implications apply whether paleosol carbonate represents mean annual conditions or if carbon-ate formation is warm-season biased, and the modern abundance of C

4 biomass in the region

appears to follow the same trend. If paleosol carbonate refl ects mean annual conditions, as is assumed in model 4 for estimated modern carbonate δ18O value (Table 1), the estimated δ18O value of modern carbonate is lower than the mean for any interval in the past (Fig. 5B)

and equal to or less than 35 of the 194 measured values. If paleosol carbonate is warm-season biased in terms of the temperature of formation (model 6, Table 1), the estimated modern δ18O value of soil carbonate is even lower (Fig. 5B) and equal to or less than only six measured val-ues, and the temporal trend in δ18O values from the Clarendonian to modern is even larger than under model 4.

The association of increased C4 abundance

and colder climate is obvious given (1) the generally coincident timing of increasing C

4

abundance over the Neogene documented here (and in numerous other studies for other regions globally) and the long-term record of global cooling since the middle Miocene from benthic foraminifera (Zachos et al., 2001), (2) the fi rst appearance of modern abundance of C

4 biomass

after the onset of Northern Hemisphere glacia-tion and the fi rst penetration of the Laurentide ice sheet into central North America (Haug et al., 2005; Rovey and Balco, 2009), and (3) the presence of frost-tolerant reptiles in the Meade Basin until into the Irvingtonian (Schultz, 1969), but not afterward or today, despite the continued increase in the abundance of C

4 biomass to the

Figure 5. (A) Reduced major axis (RMA) linear regression of δ13C values for all intervals on corresponding δ18O values. Solid black line is for all data; dashed black line excludes the three samples with δ18O >29‰ in the Borchers section north of Kansas Highway 23 (late Blancan–early Irvingtonian). (B) RMA linear regression of mean δ13C values on mean δ18O values for each interval. Late Blancan–early Irvingtonian means do not include the three samples with δ18O >29‰. Modern δ13C values were calculated from the mean modern abun-dance of C4 biomass in the region (78% ± 10.9%; Fox et al., 2011) assuming an enrichment between organic matter and carbonate of +15‰. Gray bars indicate standard deviation of modern organic matter δ13C values. Modern δ18O values are for assumptions used in models 4 and 6 (see Table 1 and discussion in text).





modern in the region at present (Fig. 5B; Fox et al., 2011). Because at midlatitudes δ18O

mw is

closely tied to temperature (i.e., Fig. 3A), cool-ing alone could lead to a decrease over time in carbonate δ18O because the direct relationship of carbonate and soil water δ18O has a greater effect than the indirect relationship between carbonate δ18O and soil temperature. However, under the assumptions of model 4 and the rela-tionship between temperature and δ18O

mw in

Figure 3A (i.e., δ18Omw

changes 0.43‰/°C), the decrease in carbonate δ18O from the Clarendon-ian to the estimated modern value would require a decrease in mean annual temperature since the Clarendonian of 18 °C (if all months changed equally), implying mean annual and maximum monthly temperatures for the Clarendonian of 30.2° and 42.8 °C, respectively, which are implausibly high. Assuming no seasonality of temperature during the Clarendonian and allow-ing all long-term cooling to take place during the winter (which effectively converts model 4 into model 2), model 4 would require an even higher Clarendonian mean annual temperature and a decrease in winter month temperature to the modern in excess of almost 35 °C to explain the difference between mean Clarendonian δ18O and the modern estimate for model 4. For cooling alone, model 6 requires monthly tem-peratures during the Clarendonian to be almost 22 °C higher than today, yielding a mean warm season temperature of 42 °C.

For the assumptions of models 4 and 6 and no change in temperature, an increase in the proportion of cool season precipitation from an assumed 0% for the Clarendonian to the mod-ern proportion of 34% (which makes the models equivalent) only yields a decrease in carbonate δ18O value of ~2‰, which is not suffi cient to explain the 4‰–5‰ trend from the Clarendon-ian to the estimated modern value. If we com-bine the same increase in the proportion of cool season precipitation with a modest 5 °C cooling, and allow δ18O

mw to follow temperature as in

Figure 3A, we still only fi nd a 3.3‰ decrease in carbonate δ18O value. Thus, under either model of the timing of carbonate formation, plausible changes in temperature since the Clarendonian also require an increase in moisture availability to explain the trend in measured and estimated carbonate δ18O values (Fig. 5B).

If mean monthly (hence mean annual) tem-perature decreased 5 °C from the Clarendonian to the present (and if δ18O

mw follows temperature

as in Fig. 3A), under the assumptions of model 4, a decrease from the Clarendonian to the pres-ent in evaporative enrichment of soil water in 18O by 2.9‰ in every month or 4.5‰ in the warm season can account for the trend in δ18O values. For the same temperature decrease under model

6, a decrease in evaporative enrichment of 4‰ in all months or 6‰ in the warm months can account for the trend. Under either model, the necessary decrease in evaporative enrichment is smaller if the Clarendonian was even warmer or if the proportion of winter precipitation recharge to soil water has also increased. Additionally, as discussed already, the models of modern carbonate use measured meteoric water δ18O values at Scott Lake, Kansas, as soil water and do not adjust for evaporative enrichment today. Thus, the changes in evaporative enrichment are maximal estimates for the Clarendonian and would decrease in subsequent intervals. Seasonal enrichments of soil water in 18O of 5‰–10‰ have been observed in Hawaii and in mixed C

3:C

4 grasslands in Colorado (Hsieh et

al., 1998; Ferretti et al., 2003), so the changes suggested here are not implausible.

None of the values discussed here is intended to be an actual estimate of changes in environ-mental parameters from the Clarendonian to the present. Instead, the intent is a sensitivity analy-sis to determine the scale of change in the rel-evant parameters (temperature, δ18O

mw, seasonal

distribution of precipitation, moisture availabil-ity) necessary to explain the observed trend in δ18O values. These analyses indicate that some combination of modest cooling and decreasing δ18O

mw, increased percentage of winter precipi-

tation, and decreased evaporative enrichment of soil water in 18O can explain the changes in carbonate δ18O values in the Meade Basin. The critical parameter that would allow quan-titative estimation of the other parameters, and therefore various aspects of the climate record in the region, is soil temperature, which could be measured using the clumped isotope or Δ

47

paleothermometer on these same samples (Eiler, 2007; Passey et al., 2010). Such analyses would allow for a direct comparison with modern con-ditions. Direct reconstruction of mean annual precipitation from physical characteristics of paleosols (i.e., depth to carbonate horizon; Sheldon and Tabor, 2009) could be useful, but most calcareous layers in the Meade Basin have sharp and erosional tops, indicating that few if any sections in the Meade area have preserved A horizons (necessary for accurate measure-ment of depth of carbonate below the ancient land surface). Moreover, the suitability of depth to carbonate as a proxy for mean annual pre-cipitation has been questioned (Royer, 1999; but see Retallack, 2000). Regardless, the increased abundance of C

4 biomass in the region from the

Clarendonian to the present does not appear to correspond to warmer or more arid conditions but, based on the carbonate δ18O record, in fact corresponds to cooler and/or less arid condi-tions. This conclusion is independent of whether

soil carbonate isotopic records generally record mean annual conditions or are biased to the warmer, drier season (Breecker et al., 2009).

The long-term relationship between the abun-dance of C

4 biomass and the climatic implica-

tions of carbonate δ18O values is contrary to common assumptions about the role of tem-perature and moisture in the evolution of eco-logical dominance by C

4 grasses in grassland

ecosystems. However, Edwards and Still (2008) have shown using a phylogenetic approach that C

4 grass species in Hawaii are not associated

with warm temperatures necessarily because of C

4 photosynthesis but because they are mem-

bers of a warm climate clade that includes both C

3 and C

4 species. Edwards and Still (2008)

do not rule out aridity as a factor in C4 abun-

dance on Hawaii, but other studies have ques-tioned the importance of aridity in the evolu-tion of C

4 photosynthesis (Keeley and Rundel,

2003). Edwards and Smith (2010), using a more comprehensive phylogeny of grasses, showed that C

4 grass clades generally have basal sister

taxa that are warm-adapted C3 species, imply-

ing that the climatic tolerance and geographic distribution of C

4 grasses are likely not a conse-

quence of photosynthetic pathway. According to Edwards and Smith (2010), evolutionary tran-sitions from C

3 to C

4 photosynthesis in grasses

are generally associated with shifts into drier (but not warmer) habitats, but the climatic shifts associated with the multiple origins of C

4 pho-

tosynthesis in grasses are, on average, within what would be considered mesophytic (not arid) conditions. The average reconstructed mean annual precipitation for C

3 nodes that are ances-

tral to C4 clades is 1839 mm/yr, and the aver-

age for the basal node of descendant C4 clades

is still 1229 mm/yr (compare to 545 ± 132 mm mean annual rainfall and 1016 ± 289 mm mean annual total precipitation for Meade; Meade sta-tion NCDC Summary of the Day, www.ncdc.noaa.gov/oa/ncdc.html). Moreover, Edwards and Smith (2010) identifi ed no transitions to C

4

photosynthesis that were associated with mean annual precipitation (MAP) <500 mm/yr.

The climatic preferences of specifi c grass clades, which are not necessarily related func-tionally to photosynthetic pathway (Edwards and Still, 2008; Edwards and Smith, 2010; Edwards et al., 2010), may provide an explana-tion for the increase of C

4 biomass in southwest

Kansas in association with some combination of cooler and wetter conditions. The C

3 clade

Pooideae, which includes the most abundant C3

species in the Great Plains (Sims, 1988; Fred-lund and Tieszen, 1997), has the strongest pref-erence for both cool and dry conditions among major grass clades. One of the two clades of C

4

grasses that includes an ecologically dominant



Fox et al.


species in the Great Plains, Chloridoideae, has similar preference for dry conditions, but the other ecologically dominant C

4 clade, Panicoi-

deae, prefers wetter conditions than either Pooi-deae or Chloridoideae (Sims, 1988; Fredlund and Tieszen, 1997; Taub, 2000; Edwards and Smith, 2010). Thus, it is possible that greater moisture availability, despite cooler conditions, allowed species of Panicoideae to become eco-logically dominant at the expense of both Pooi-deae and Chloridoideae independent of pho-tosynthetic pathway. This idea could be tested with a combination of phylogenetic analysis, niche modeling, and biogeography for species in the Great Plains.

CONCLUSIONS

In parallel to the trend of increasing C4 abun-

dance recorded by the δ13C values of carbon-ates in southwest Kansas (Fig. 2; Fox et al., 2011), the associated δ18O values decrease from the Clarendonian to the Irvingtonian, and values predicted for modern carbonate in the region are lower yet. Based on the oxy-gen isotope composition of soil carbonates, the increase in C

4 biomass in southwest Kan-

sas was not associated with a long-term trend toward greater aridity or warming, but instead occurred in association with some combina-tion of cooling, increased available moisture, and/or an increase in the proportion of soil water derived from low-δ18O winter precipita-tion (i.e., direct cool season rainfall and spring infi ltration of meltwater from snow). Thus, the association of high C

4 species abundance or

biomass with warm/arid conditions in modern ecosystems (Teeri and Stowe, 1976; Hatter-sley, 1983; von Fischer et al., 2008) may not be refl ective of the climatic trends associated with the increase in C

4 biomass over geologi-

cal time, at least in the southern Great Plains. Other factors that remain poorly constrained in the Neogene geological record, such as fi re frequency, herbivore pressure, and seed and seedling predation, may be of equal or greater importance in the history of C

4 biomass (Jan-

zen, 1971; Edwards and Crawley, 1999; Bond and Keeley, 2005; Bond, 2008; Osborne, 2008; Edwards et al., 2010), and these fac-tors probably have complex interdependencies with each other and with climate change. Our results, and those of Edwards and Still (2008) and Edwards and Smith (2010), indicate that some ideas about the climatic controls on the evolution of C

4 grass abundance may need to

be reevaluated, whether in the context of phy-logenetic analysis, as suggested by Edwards et al. (2010), or additional high-resolution paleo-ecological records.

ACKNOWLEDGMENTS

This research would not have been possible without access to private land granted by numerous landown-ers in the Meade area. Jenn Campbell, Mark Clem-entz, Neil Kelly, Sam Matson, and Seth Newsome provided excellent help with fi eld work for this project over several seasons. We thank Sam Matson for com-ments on a draft of this paper; Bill Gilhooly, Andrea Lini, Nate Sheldon, and one anonymous reviewer for thoughtful reviews on an even longer previous version of this manuscript that included the data and discus-sion in Fox et al. (2011), and Bill Gilhooly and Andrea Lini for additional reviews on the fi nal versions of this paper and Fox et al. (2011). This research was funded by grants from the National Science Foundation 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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