Stable isotope constraints on the elevation history of the Sierra Nevada Mountains, California

588 For permission to copy, contact [email protected]© 2008 Geological Society of America

ABSTRACT

Research on the uplift history of the Sierra Nevada mountain range has yielded seemingly confl icting results. Some studies argue for substantial uplift within the past 3–5 m.y.; others suggest that high elevations may have existed since the Cretaceous. The rain shadow across the Sierra Nevada is associated with a strong isotopic gradient, with lower δ18O values in precipitation on the leeward side of the range. Reconstruction of the δ18O value of meteoric water as a moni-tor of paleoelevation has focused mainly on the leeward side of the Sierras, but inter-pretation of the results of these studies may be complicated by shifts in global climate and regional moisture sources. We address these concerns by analyzing the δ18O value of tooth enamel bioapatite from contempo-raneous mammalian fossils on either side of the present Sierra range. By sampling across the range, δ18O differences induced by a rain shadow can be isolated from other compli-cating factors. Our results indicate that the Sierra rain shadow has existed since at least 16 Ma, which is an important constraint on models for the tectonic evolution of the west-ern United States. Unfortunately, temporal resolution for localities is too coarse to dif-ferentiate between glacial and interglacial localities during the past 2 m.y., so we can-not evaluate if there was a latest Cenozoic pulse of uplift or elevation loss.

Keywords: Sierra Nevada, stable isotopes, bio-apatite, oxygen, enamel.

INTRODUCTION

The elevation history of the Sierra Nevada mountain range is a subject that has spawned substantial debate. The modern average elevation of the range is ~2800 m, increasing from 2100 to 2700 m in the north to 4000–4400 m in the south (Wakabayashi and Sawyer, 2001). Various lines of evidence suggest that the Sierras experienced a phase of Late Cenozoic uplift, albeit as part of a longer, complex history of uplift. Yet other lines of evidence indicate that substantial topog-raphy existed in the Sierras prior to 15 Ma; some authors argue for substantial elevation in the Paleogene or late Cretaceous (Lindgren, 1911; House et al., 1998; Chamberlain and Poage, 2000; Wakabayashi and Sawyer, 2001; Horton et al., 2004; Stock et al., 2004; Mulch et al., 2006).

We offer new constraints on the uplift history of the Sierras through oxygen isotope analysis of fossil mammals. The present Sierra Nevada range creates an imposing barrier to moisture fl ow from west to east that generates a conspicu-ous “isotopic rain shadow” (Smith et al., 1979, 2002; Kendall and Coplen, 2001). Fossil teeth have been used to reconstruct ancient isotopic rain shadows in other areas (Kohn et al., 2002; Fricke, 2003). For the Sierras, there have been studies using authigenic minerals on either the leeward (Chamberlain et al., 1999; Chamberlain and Poage, 2000; Horton et al., 2004; Abruzz-ese et al., 2005; Takeuchi and Larson, 2005) or windward side (Mulch et al., 2006) of the range. Yet because these studies do not reconstruct the composition of meteoric water on both sides of the range, it is diffi cult to tease apart the oro-graphic signal from other factors that infl uence isotopic value, especially climate change.

Some recent isotopic studies have offered quantitative estimates of the paleoelevation (Chamberlain and Poage, 2000; Rowley et al.,

2001; Mulch et al., 2006). Such approaches require assumptions about temperature and air fl ow patterns that are diffi cult to verify. Con-sequently, we do not attempt to place quantita-tive constraints on paleoelevation. Rather, we use isotopic differences in contemporaneous samples of the same genus across the range to determine if a signifi cant topographic barrier has existed between California and Nevada over the past 16 m.y. Elevations as low as half a kilometer are thought to form a topographic bar-rier with a signifi cant isotopic gradient (Fricke, 2003). Therefore, even low paleoelevations in the Sierras should be detectable.

History of the Sierra Nevada Mountain Range

The Sierra Nevada range began as a volcanic arc during the Cretaceous with emplacement of the Sierra Nevada batholith 125–80 Ma. These rocks crystallized 11–15 km below the surface and are exposed today, testifying that a great deal of uplift and exhumation has occurred since 85 Ma (Ague and Brimhall, 1988; Small and Anderson, 1995; Wakabayashi and Sawyer, 2001). Based on evidence from sedimentation rates, the range may have experienced rapid uplift between the Late Cretaceous and Eocene (85–50 Ma) (Wakabayashi and Sawyer, 2001). Elevation gain during this time could have been due to rebound after cessation of volcanism, or related to Laramide deformation across the western United States (Chase et al., 1998; Wak-abayashi and Sawyer, 2001).

Exhumation tapered off by ca. 50 Ma (Huber, 1981; Unruh, 1991; House et al., 1998, 2001; Wakaba yashi and Sawyer, 2001). Approxi-mately 35–33 Ma, extension of the Basin and Range began to the east of the northern Sier-ras. This extension reached the southern Sierras

Stable isotope constraints on the elevation history of the Sierra Nevada Mountains, California

Brooke E. Crowley†

Department of Anthropology, University of California, Santa Cruz, California 95064, USA

Paul L. KochDepartment of Earth and Planetary Sciences, University of California, Santa Cruz, California 95064, USA

Edward B. DavisMuseum of Vertebrate Zoology, University of California, Berkeley, California 94720, USA

†E-mail: [email protected]

GSA Bulletin; May/June 2008; v. 120; no. 5/6; p. 588–598; doi: 10.1130/B26254.1; 5 fi gures; 2 tables; Data Repository Item 2008029.

Isotopic elevation history of the Sierra Nevada

Geological Society of America Bulletin, May/June 2008 589

by 14 Ma (Dilles and Gans, 1995). Extensive volcanism associated with extension lasted until 6–5 Ma, when motion changed between the North American plate and both the Sierra Nevada microplate and Pacifi c plate (Atwater and Stock, 1998; Argus and Gordon, 2001). Some volcanism continues to persist along the eastern border of the range.

Debate over Late Cenozoic Uplift

Given the abundant evidence for tectonic activity involving the Sierras reviewed above, high elevations in the region prior to 10 Ma seem entirely plausible. Investigations supporting this conjecture rely on: (1) the oxygen or hydrogen isotope values of minerals as a proxy for meteoric water, which varies with rainout over mountain ranges (Chamberlain and Poage, 2000; Mulch et al., 2006); (2) cosmogenic isotope studies to quantify erosion rates (Stock et al., 2004, 2005); (3) U-Th/He geothermometry and He-diffusion dating to place limits on the evolution and mor-phology of elevation profi les (House et al., 1998, 2001); and (4) paleobotanical studies of leaf phys-iognomy to determine free-air enthalpy, which co-varies with elevation (Wolfe et al., 1997).

Most studies arguing for recent elevation gain in the Sierras suggest it occurred within the past 10 m.y., largely between 5 and 3 Ma (Huber, 1981; Unruh, 1991; Ducea and Saleeby, 1998; Jones et al., 2004). These studies fall into three categories: (1) those that use the inclina-tion of Late Cenozoic strata to infer incision rates (Huber, 1981; Unruh, 1991; Wakabayashi and Sawyer, 2001); (2) studies document-ing changes in stream gradients and incision rates through time (Dalrymple, 1964; Huber, 1981; Wakabayashi and Sawyer, 2001), and (3) paleobotanical studies using the presence of temperate plant fossils in the high Sierras as indicators of local paleoelevation (Axelrod, 1962). The cause for possible recent elevation gains is not well understood. The foundering of a dense eclogitic root under the southeast-ern Sierra Nevada has been proposed as one option (Ducea and Saleeby, 1998; Jones et al., 2004). Changes in motion on plate boundar-ies, or isostatic rebound in response to high erosion rates under glacial climates, are other suggested causes (England and Molnar, 1990; Molnar and England, 1990; Unruh, 1991; Small and Anderson, 1995; Atwater and Stock, 1998; Wakabayashi and Sawyer, 2001).

There are variations on these end-member scenarios. For example, while Huber (1981) allows that uplift may have begun ca. 25 Ma, he proposes that most of the current topography arose in the past 3 m.y. Conversely, some papers note that recent uplift occurred but that it was

minor compared with earlier periods of eleva-tion gain (Stock et al., 2004, 2005).

Finally, some studies point to Late Cenozoic elevation loss, not gain (Wernicke et al., 1996; Wolfe et al., 1997; House et al., 1998, 2001; Hor-ton et al., 2004). Yet, as noted by Wakabayashi and Sawyer (2001), if net elevation loss has occurred, then stream incision rates should be exceeded by ridge-top erosion. This pattern is not seen in the Sierras (Small and Anderson, 1995), where most studies have revealed an increase in stream incision since ca. 10 Ma (Huber, 1981; Unruh, 1991; Stock et al., 2004, 2005).

Stable Oxygen Isotopes in Precipitation and Isotopic Rain Shadows

As a moist air mass moves inland from its source, water enriched in 18O will condense and precipitate more readily than water enriched in the lighter isotope, 16O (Araguas Araguas et al., 1996). This distillation process is temperature dependent, and in temperate latitudes mean annual temperature is correlated with the mean annual oxygen isotope (δ18O)1 value of meteoric water (Dansgaard, 1954, 1964; Rozanski et al., 1993). In a similar fashion, as an air mass moves over a topographic barrier, adiabatic cooling causes condensation and loss of moisture, and 18O-enriched water will again condense more readily than 16O-enriched water. As a conse-quence of this “elevation effect,” the relatively dry air masses on the lee side of mountain ranges have water that is 16O-enriched relative to wet-ter air masses on the windward side (Smith et al., 1979; Araguas Araguas, 1996). The higher the range, the more pronounced the isotopic rain shadow. Thus, the δ18O value of meteoric water (rain and snow) can be used to track the rain shadow of a mountain range, which is, in turn, a proxy for the elevation of the topographic barrier separating the region from the source of moisture (Chamberlain and Poage, 2000).

Today, the Sierras form a large orographic barrier with a substantial rain shadow. Based on storm tracks and precipitation measurements, only 10% of precipitation derived from the Pacifi c Ocean is estimated to reach the eastern side of the Sierras (Smith et al., 1979; Friedman et al., 2002b). The Sierras also create a strong isotopic rain shadow (Friedman et al., 2002b). California’s Coast and Transverse Ranges also have signifi cant relief, but based on stream data, they appear to have negligible isotopic rain shadows compared with the Sierras (Kendall and Coplen, 2001) (Figs. 1 and 2).

The storms reaching the lee side of the Sierra come from several different regions and have different isotope values (Smith et al., 1979; Friedman et al., 2002a). Most winter storms begin in the Pacifi c Ocean and move eastward over the Sierras. Summer storms, in contrast, advance northward from the Gulf of California or eastward from the Pacifi c Ocean. Most of the precipitation falling directly on the Sierras, as well as immediately east of the range, accumu-lates as snow during January and February and is derived from Pacifi c sources that move from west to east across the range.

Relative humidity can also affect the iso-topic value of surface water (rivers, lakes, and streams). In areas of high humidity, small streams and ponds may accurately refl ect the isotopic value of precipitation. Conversely, sur-face water in arid areas may undergo signifi cant evaporative 18O-enrichment (Gat, 1996). With increased catchment area, water contributing to larger streams, ponds, and lakes will be a mix-ture of precipitation from different elevations and different microclimates. This homogeniza-tion of surface water may lead to isotopic differ-ences between local surface water and precipi-tation (Gat, 1996; Kendall and Coplen, 2001). Because stream water tends to be sourced from higher elevations, local topography can exac-erbate this difference (Dutton, 2003). Thus, depending on stream size and local topography,

Figure 1. Modern δ18O values for streams and rivers across California, after Kend-all and Coplen (2001).

1δ18O = ((Rsample

/Rstandard

)-1)*1000, where R = 18O/16O and the standard is Vienna Standard Mean Ocean Water (V-SMOW).

Crowley et al.

590 Geological Society of America Bulletin, May/June 2008

surface water may refl ect the isotopic value of local or regional precipitation.

Stream water δ18O values show a gradient of ~4‰ from the Pacifi c Coast eastward to the Sierra foothills, then drop very sharply by another ~6‰ just west of the crest of the Sier-ras (Fig. 1) (Kendall and Coplen, 2001). There is a negligible isotopic gradient of 2‰–3‰ east of the Sierras spread across the state of Nevada (Fig. 1). The δ18O of precipitation is variable, but the average difference across the Sierras (6‰–7‰) and the overall range of data (10‰–12‰) are similar to those of stream water (Fig. 2). Unlike stream water, however, precipitation does not show a regular decrease from the coast to the Sierra foothills. We suspect the gradient in stream water δ18O values across California west of the Sierras is smoothed as a result of transport of 18O-depleted waters from high elevations westward to lower elevations.

Stable Isotopes in Mammals

We use the oxygen isotope value of mamma-lian tooth enamel to reconstruct the isotopic rain shadow of the Sierra Nevada Mountains over the past 16 m.y. We use enamel hydroxyapa-tite [Ca

10(PO

4, CO

3)

6)(OH, CO

3)

2], because it is

much more crystalline than bone or dentin apa-tite and retains in vivo isotopic values with high fi delity. Though alteration can occur and may be evaluated through a variety of tests (Wang and Cerling, 1994; Koch et al., 1997; Zazzo et al.,

2004a, 2004b). Our discussion of controls on the δ18O value of vertebrate hydroxyapatite fol-lows the reviews by Koch (1998) and Kohn and Cerling (2002). For large mammals, the chief control on the δ18O value of body water, from which bioapatite crystallizes, is the δ18O value of ingested water. For a large mammal that is obli-gated to drink, most ingested water will be taken in by drinking surface water. Because some of the oxygen in body water comes from O

2 via

aerobic respiration, and because O2 is relatively

invariant spatially, the relationship between δ18O value in body water or bioapatite and δ18O in ingested water and/or surface water will not be 1:1. In general, differences in enamel or body water δ18O values are 70%–90% of differences in the δ18O value of the ingested water and/or surface water. The estimate for equids is 71% (Delgado Huertas et al., 1995).

Physiological differences in water use effi -ciency among species may also affect the δ18O value of body water. Major changes in ambient temperature have some impact as well, most likely by altering the evaporative stress on ani-mals (Hoppe et al., 2004a). Finally, leaf water can make a large contribution to the water bud-get of some species. Bryant and Froelich (1995) estimated that 20%–30% of ingested oxygen comes from plant water, even in obligate drink-ers. Depending on levels of humidity, water in leaves can be 18O-enriched relative to meteoric water by a substantial amount (Sternberg, 1989). Since these processes are more pronounced in

dry climates, we expect these environmental and physiological processes will make bioapa-tite δ18O values higher and more variable in hot and arid regions.

We attempted to minimize the impact of physiological and ecological factors by ana-lyzing similar species or genera. Equids were the preferred target of our study because they are large mammals with large, robust teeth. Additionally, they are obligate drinkers, their remains are relatively abundant in fossil deposits, and they existed across the western United States for over 20 m.y. (MacFadden et al., 1999; Hoppe et al., 2004b). Different teeth mineralize and erupt at different times in the fi rst year of a horse’s life. Because nursing may potentially affect the isotopic value of a mam-mal’s tissues, care should be taken to sample teeth that form post-weaning (Bryant, 1995), although some studies of wild animals have suggested the “nursing effect” is small to non-existent (Kohn et al., 1998, 2002; Gadbury et al., 2000). In horses, adult premolars (P2–P4) and the adult third molar (M3) mineralize after weaning (Bryant, 1995; Hoppe et al., 2004b). Unfortunately, discriminating among P3–P4 and M1–M2 is diffi cult when using isolated teeth, which are the most common and least valuable specimens for morphological study. We attempted to use only P2 and M3, but to get a robust sample for some localities, we could not always follow this stricture. Additionally, because cheek teeth mineralize over a period of time, there are isotopic variations along an individual tooth and among teeth within an individual (Bryant, 1995; Higgins and Mac-Fadden, 2004; Hoppe et al., 2004a).

METHODS

Sampling Protocol and Analysis

Fossil specimens from California and Nevada were provided by the Museum of Paleontology at the University of California, Berkeley (UCMP), and modern Equus teeth were provided by the Nevada Bureau of Land Management. Samples were analyzed at the University of California Santa Cruz (UCSC) Stable Isotope Laboratory (Table DR1; see footnote 2). Fossil specimens were placed into North American Land Mam-mal Ages (NALMA). When possible, more precise ages were determined using UCMP’s online MioMap database (Carrasco et al., 2005)

-18

-16

-14

-12

-10

-8

-6

-4

-2

0

0 200 400 600 800 1000 1200

Distance from coast (km)

δ18O

‰ (

SM

OW

)

~4-6ä ~6-7‰

Figure 2. The δ18O value of modern precipitation with increasing distance from the Pacifi c Coast (precipitation source) across the Sierras. Oxygen data come from Benson (1994), Friedman et al. (1992, 2002b), and the International Atomic Energy Agency (IAEA) GNIP2001 yearly database. Distance from the coast was approxi-mated using a coordinate distance calculator (http://boulter.com/gps/distance). Shaded region is the approximate location of the Sierra crest.

2GSA Data Repository Item 2008029, Table DR1, containing specimen information, age estimates, and isotopic results, is available at www.geosociety.org/pubs/ft2008.htm. Requests may also be sent to [email protected].



or the Faunmap database (Faunmap Working Group, 1994). Otherwise, chronological ages for NALMAs are as cited in Woodburne (2004). Our specimens range in age from Hemingford-ian (18.9–16 Ma) to modern (Fig. 3).

Enamel powder was collected by drilling perpendicular to growth increments along the length of each tooth. Hoppe et al. (2004b) cal-culated that 3–4 vertical cm should be sampled to integrate the time of tooth crown formation. To minimize damage to teeth, however, we gen-erally sampled little more than 1 cm. Thus our samples are likely averaging months of growth (Higgins and MacFadden, 2004; Hoppe et al., 2004b). Given seasonal and inter-individual variations, multiple individuals should be sam-pled to obtain a robust estimate of the mean δ18O value for each locality (Clementz and Koch, 2001; Fricke, 2003). Unfortunately, access to specimens from some localities and time peri-

ods was very limited, and some localities are represented by single specimens (Table DR1; see footnote 2). To bolster sample sizes at some localities, we analyzed samples from several other large, obligate-drinking taxa (proboscid-eans and rhinocerotids).

To isolate the carbonate fraction for isotopic analysis, ~10 mg of powdered enamel were pretreated following the protocol of Koch et al. (1997). Briefl y, 0.5 ml of 2%–3% sodium hypochlorite (NaOCl) was added to each pow-dered sample and left for 24 h to oxidize organic material. Samples were rinsed 5× with ultra-pure water, then reacted for 24 h with 0.5 ml of 1 M acetic acid buffered to pH 5 with calcium acetate to remove non-lattice bound carbonates. Samples were again rinsed 5× with ultrapure water, and freeze-dried. Samples were heated under vacuum at 65 °C for 1 h prior to analysis to remove any remaining water.

Approximately 1.5 mg of each sample were analyzed on a Micromass Optima gas source mass spectrometer integrated with an Isocarb automated carbonate device in the Departments of Earth and Ocean Sciences, University of Cali-fornia, Santa Cruz. Samples were dissolved in 100% H

3PO

4 at 90 °C, with concurrent cryo-

genic distillation of CO2 and H

2O and automated

CO2 admittance to the mass spectrometer for

analysis. Reaction time was set at 12 min and blanks were run between samples. Standards used in this study were Carrera Marble (CM) and National Bureau of Standards (NBS)-19. Data analyzed on different days were normalized for minor instrumental differences using a fi xed value for CM. The mean and standard deviation for NBS-19 analyzed with samples was −2.21 ± 0.04‰ (n = 32), very close to the known value of −2.20‰. The average difference in δ18O value between replicates was 0.34‰ for duplicates (n = 56) and 0.39‰ for triplicates (n = 16). Three specimens had δ18O differences of >1‰ after replication (Table DR1). These specimens were considered unreliable and were excluded.

Statistical Analysis

Statistical analyses were performed using SigmaStat 2.03 and Systat 10. We performed general linear models (GLM), Student’s t-tests (for comparisons of means for two populations), and one- and two-factor analysis of variance (ANOVA) (for comparisons of means of three or more populations). If signifi cant differences in mean were detected by ANOVA, we applied pairwise comparisons (post-hoc Tukey tests) to identify the source of signifi cant variation.

For statistical comparisons, data were clus-tered into two main groups—localities west of the Sierras and localities east of the Sierras. A one-factor ANOVA indicates that this assign-ment of animals to a side of the range is a reason-able categorical proxy for longitude (p <0.001). One-factor ANOVA indicates no signifi cant dif-ference among taxa at localities with multiple taxa (p <0.05). Taxa were thus considered to be statistically indistinguishable for subsequent statistical tests.

The crest of the Sierras is thought to have shifted westward with the inception of Basin and Range extension and the collapse of the eastern portions of the range. It is possible that some localities that are east of the crest today might have been west of the crest in the past (V5691, V5693, and V5694, Inyo, California, Table DR1) (Jones et al., 2004). These localities were omitted from statistical tests.

For analysis of gradients across the range, localities were clustered by NALMA. Because data are not evenly distributed on either side of

Sierra N

evada

Coast R

angesRancholabreanIrvingtonianBlancanHemphillianClarendonianBarstovianHemingfordian

Modern

–82–00–146 146–333333–532532–736736–933933–11381138–13491349–15601560–17811781–20182018–22802280–25742574–29122912–33203320–4378

Meters

Nevada DEM 90 from Nevada Bureau of Mines and Geology

0 440 Kilometers

Transverse Ranges

55 110 330220

Figure 3. Specimen localities organized by North American Land Mam-mal Ages (NALMA). See Table DR1 for locality details.

Crowley et al.


the range through time, four NALMAs were chosen that had suffi cient numbers of samples. Student’s t-tests were used to compare δ18O val-ues east and west of the range for these temporal bins (Table 1).

RESULTS

There is a signifi cant difference in enamel δ18O values across the Sierras today and dur-ing the Hemphillian (9–4.9 Ma, Late Miocene to Early Pliocene), Clarendonian (12.5–9 Ma, Late Miocene), and Barstovian (16–12.5 Ma, Middle Miocene) NALMAs (Student’s t-test, p <0.05) (Table 1). Even for time intervals lacking suffi cient data to perform statistical tests, it is evident that δ18O values differ between localities on the east and west sides of the range (Fig. 4). However, there is extensive overlap in the δ18O values for animals from either side of the range during the Rancholabrean (RLB). Consequently, there is no signifi cant difference across the Sier-ras for the samples from the past two million years (Student’s t-test, p = 0.106) (Table 1).

A general linear model (GLM) reveals a signifi cant interaction between “side” and age (p <0.005) (Table 2). However, this interac-tion is likely due to the insignifi cant difference across the Sierras in the Rancholabrean (Fig. 4). Because these oxygen values overlap across the range, a second set of statistical tests was per-formed where fossil data younger than 2 Ma were excluded. Without values from more recent time intervals, there is no signifi cant interac-tion term between side and age. A subsequent two-factor ANOVA including the modern data indicates consistent signifi cant isotopic offset both in modern times and from 18 to ca. 2 Ma between taxa on the east and west sides of the Sierras (p = 0.05) (Table 2, Fig. 4).

This ANOVA indicates that δ18O values have statistically signifi cant variations with age as well as side. Temporal variation is expected, because meteoric water δ18O values are sensi-tive to the climatic fl uctuations that occurred over the past 20 million years (Zachos et al., 2001). Indeed, it was the possibility of such

variation that led us to examine δ18O gradients across the range, rather than just on the lee side as in most other analyses. Latitude has little effect on enamel δ18O values. This is not sur-prising, because the latitudinal isotopic gradient in meteoric water across this region is not steep (MacFadden et al., 1999; Kendall and Coplen, 2001). Moreover, the effect of latitude is small compared with the importance of the rain shadow effect across the range (GLM results for

12.5–9 Ma yield p = 0.619 and p = 0.002 for latitude and side, respectively).

DISCUSSION

Today, there is a ~6‰–7‰ decrease in δ18O of precipitation and streams across the crest of the Sierras. Poage and Chamber-lain (2001) calculate a global average lapse rate of −0.28‰/100m. Thus this ~6‰–7‰

Figure 4. Bivariate plot of δ18O value versus age. West* denotes collection localities that are currently east of the Sierras but were possibly west of the Sierras 3 Ma. Open symbols signify animals that lived west of the range. Closed symbols denote animals that lived east of the range. Circles denote equids, while squares are used to signify ungulates other than horses. For viewing purposes, modern specimens are set 0.5 Ma above the Rancholabrean (RLB) specimens. (A) Each point represents one individual. (B) Average δ18O values for each Land Mammal Age. Error bars represent ± one standard deviation. Data from this study are sup-plemented with data from Feranec (2004), Hoppe (2006), and Connin et al., (1998).

TABLE 1. STATISTICAL RESULTS COMPARING ANIMALS FROM BOTH SIDES

OF THE SIERRAN CREST NALMA df p-value Modern 8 0.029 Rancholabrean 55 0.106 Hemphillian 34 0.016 Clarendonian 98 <0.001 Barstovian 60 <0.001 Note: Tests are considered significant if p <0.05. NALMA—North American Land Mammal Age; df—degrees of freedom. Hemphillian comparisons were done using a Mann-Whitney rank sum test. All other NALMAs were tested using a Student's t-test.

Modern

Irvingtonian

Blancan

Hemphillian

Clarendonian

Barstovian

Hemingfordian

RLB0

2

4

6

8

10

12

14

16

18

121620242832

Ag

e (M

a)

δ18O‰

0

2

4

6

8

10

12

14

16

18

121620242832

West

East

West*

East Equids

West Equids

A

B



decrease in δ18O across the range corresponds to 2100–2500 m of elevation between the foot-hills and crest, in general agreement with the 2800 m average elevation of the range.

Based on our results, there has been a signifi -cant isotopic difference between animals living on the east and west sides of the Sierras since at least 18 Ma (Fig. 4). In examining tooth enamel from modern animals on either side of the range, we fi nd that the modern δ18O shift across the Sierras is similar to that prior to 2 Ma. Avail-able data suggest that enamel δ18O values from modern wild horses in Nevada are similar to fossil values. No published data presently exist for modern California horses. However, enamel δ18O values from Catalina Island Bison bison fall within the range of δ18O values recorded for animals west of the Sierras through time (Fig. 4) (Hoppe, 2006).

The difference in δ18O values between mod-ern ungulates east and west of the Sierras var-ies from 4‰ to 12‰. The most negative value comes from a horse that lived just east of the Sierras. The difference between this animal and the bison living on Santa Catalina Island in southern California is on the high end of the difference in δ18O values for precipitation and stream water between the Pacifi c Coast and northwestern Nevada (Figs. 1 and 2). Recall, as well, that we would expect bioapatite δ18O values from equids to underestimate the differ-ence in ingested water and/or surface water δ18O values by ~30% (Delgado Huertas et al., 1995). Bioapatite may be showing a greater-than-expected difference because of high values on Santa Catalina Island, which is considerably far-ther south than northern Nevada. The spread in δ18O values among modern horses from Nevada likely refl ects differential consumption of water sourced from snowmelt immediately east of the Sierras versus water that has been 18O-enriched by evaporation in the deserts farther to the east. Therefore, while these data provide general sup-port to our approach, they illustrate that a variety of factors may complicate our interpretations of paleoelevation from the δ18O value of mamma-

lian bioapatite. Many of these complications also affect interpretations based on other authi-genic minerals (soil or lacustrine carbonates and weathering products).

For example, although other workers have determined that high elevations have existed in the Sierras for at least the past 20 m.y. (Cham-berlain and Poage, 2000; Horton et al., 2004; Abruzzese et al., 2005; Mulch et al., 2006), they have interpreted a rise in δ18O since the late Miocene as evidence for a decrease in Sierran elevation (Horton et al., 2004). However, when sampling only east of the Sierras, it is impossi-ble to tease apart the relative effects of regional climate change and elevation on δ18O values. In our study, we did not fi nd any evidence for a decrease in elevation. Except for specimens younger than 2 Ma, fossils east and west of the Sierras consistently differ by ~6‰, and the total isotopic range in each time interval is roughly similar to that expected based on modern pre-cipitation and stream data (Fig. 4).

While we contend that this difference is cre-ated by an isotopic rain shadow resulting from a substantial orographic barrier dating back to at least 16 Ma, there are several factors that might temper our confi dence in this conclusion. The fi rst concerns the reliability of δ18O records from mammalian tooth enamel as proxies for the δ18O value of surface water. These factors relate to animal physiology and behavior, climatic and hydrologic biases, and diagenesis. A second suite of issues concerns factors other than high elevation in the Sierras that might explain a dif-ference in enamel δ18O values between western California and Nevada, such as shifts in the source of moisture or tectonic changes to the landscape other than Sierra Nevada uplift.

The Reliability of Enamel δ18O Values as a Proxy for Surface-Water δ18O Values

While some factors, especially those related to changes in humidity, can affect the net frac-tionation of oxygen isotopes between precipita-tion and tooth apatite, others can affect the reli-

ability of mammalian tooth enamel as a monitor of meteoric water δ18O values by increasing the variability in isotope values among individuals at a locality without affecting the mean δ18O value. We consider the following factors: (1) changes in humidity and evaporation, (2) seasonal and interannual variations in ingested water, (3) met-abolic and ontogenetic differences, (4) migra-tion, and (5) diagenesis.

The Impact of Changes in Humidity and Evaporation

Shifts in humidity and evaporation can impact surface-water δ18O values. If the surface water an animal ingests has experienced substan-tial evaporative water loss, it will have higher δ18O values than meteoric water. For example, in a study of water sources available to modern horses, Hoppe et al. (2004a) found that mean δ18O values of seasonal playas in eastern Ore-gon and New Mexico are comparable to local precipitation during the wet season, but they are 18O-enriched by as much as 22‰ relative to pre-cipitation in the dry season. Similarly, as plants lose water by evapotranspiration, their leaf water becomes 18O-enriched relative to surface water, with enrichments as high as 20‰ under arid cli-mates (Förstel, 1978; Sternberg, 1989). Finally, as humidity decreases, an animal loses more water by evaporation off the respiratory track or transcutaneously, and as a result, the δ18O value of its body water rises (Kohn, 1996). The cumu-lative effect of these three factors causes the difference between enamel and meteoric water δ18O values to increase as humidity decreases.

If conditions were humid on both sides of the range, a 6‰ gradient in precipitation across the range would be faithfully monitored in stream, pond, and leaf water. Actual climatic conditions rarely conform to this idealized scenario. Today, most of California is seasonally arid, and sig-nifi cant evaporation occurs in both Nevada and southeastern California (Friedman et al., 2002a). Still, humidity is lower east of the range in the desert created by the rain shadow.

We can get a sense of the impact of evapo-rative processes by comparing observed bio-apatite δ18O values for modern feral horses in the Great Basin to those expected for animals ingesting meteoric water. The mean δ18O value for modern feral horses near Reno is 24‰ (n = 3), whereas the mean annual δ18O value of precipitation and stream water is −12‰ to −14‰ (Fig. 2) (Friedman et al., 1992, 2002b; Kendall and Coplen, 2001). Using the measured bioapatite carbonate values for feral horses in Nevada, the bioapatite phosphate-to-water fractionation of Delgado Huertas et al. (1995), and the bioapatite carbonate-to-phosphate frac-tionation of Bryant et al. (1996), we estimate

TABLE 2. STATISTICAL TESTS FOR SIDE VERSUS AGE Test df MS F-ratio p-value GLM all dataSide 1 54.37 10.43 0.001 Age 5 22.38 4.29 0.039 Side*Age 5 148.93 28.56 0.0005 GLM without fossils younger than 2 MaSide 1 80.48 18.07 0.000 Age 5 19.30 4.33 0.038 Side*Age 5 0.03 0.007 0.935 Two-factor ANOVA without fossils younger than 2 MaSide 1 1114.76 251.28 0.000 Age 5 36.02 8.12 0.005 Note: GLM—general linear model; df—degrees of freedom; MS—mean squared; ANOVA—analysis of variance. Tests are considered significant for p <0.05. Asterisk signifies the interaction term between the “Side” and “Age” variables.

Crowley et al.


that horses ingested water with a δ18O value of −10.5‰, which is 18O-enriched by 3‰ relative to precipitation (Fig. 1). This is similar to the estimated 18O-enrichment in ingested water for modern feral horses in arid New Mexico and Oregon (Hoppe et al., 2004a).

The timing of development of arid conditions in the Great Basin is not highly constrained. Paleobotanical data suggest that wet summer conditions existed as far south as southern Nevada for most of the Cenozoic (Axelrod, 1962). Horton et al. (2004) used δ18O, δD, and trace-element data to argue that evapora-tive effects in the northern and western Great Basin were minor for most intervals prior to the mid-to-late Miocene. As noted above, they discovered a rise in δ18O and δD values in the Plio-Pleistocene, which they interpreted as indi-cating a decrease in elevation of the Sierras. Unfortunately, few of their independent proxies for aridity extend into the Plio-Pleistocene, so it is possible that increasing δ18O and δD values signal the increasing aridity in the Great Basin, not decreasing elevation in the Sierras.

Regardless of when it occurred, the transi-tion from more mesic to more xeric condi-tions in the rain shadow of the Sierras should

decrease the enamel δ18O gradient across the range, because evaporation coupled with less precipitation would increase the concentration of 18O in animals living east of the range. Our data do not reveal such a decrease in the δ18O gradient across the range (Figs. 4 and 5). Either the impacts of changing aridity gradients have been small, or they were offset by uplift as arid-ity increased. If arid conditions developed in the Plio-Pleistocene, however, we would not be able to detect the change, because we have sparse sampling east of the Sierras in the Plio-cene, and we see substantial overlap in enamel δ18O values across the range in the Pleistocene (Fig. 4). Ongoing climate modeling and verifi -cation studies will hopefully address some of these questions by identifying the timing and magnitude of past humidity fl uctuations in Cali-fornia and the Great Basin (Christina Ravelo, 2007, personal commun.).

Seasonal and Interannual Variations in Ingested Water

The δ18O value of meteoric water is highly correlated with annual temperature (Rozanski et al., 1993). This correlation also holds for bio-genic apatite, where δ18O values decrease during

colder conditions and increase under warmer conditions (MacFadden et al., 1999). Drought conditions and temperature shifts can create large seasonal and interannual variations. Even if two individuals lived concurrently at the same locality, if they were born in different seasons or different years, their teeth will record differ-ent δ18O values related to seasonal changes in temperature or meteoric water δ18O values (Fricke et al., 1998; Higgins and MacFadden, 2004). Such seasonal variability can be seen in data from modern horses from western Nevada and bison from Santa Catalina (Hoppe, 2006) (Fig. 4). Although we attempted to homogenize possible variations by sampling perpendicular to growth lines, seasonal or annual differences might explain some of the within-taxon vari-ability at collection localities (Table DR1). Such variability would presumably obscure, or add noise, to any intrinsic isotopic signal but would not create a consistent bias in mean values.

Metabolic and Ontogenetic DifferencesAlthough variations in the δ18O values of

ingested water are responsible for most variabil-ity in the δ18O value of mammalian bioapatite, metabolic differences do affect isotopic values

Hemphillian (9-6 Ma)

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Modern Meteoric Water

Figure 5. Change in δ18O value with increasing distance from the Pacifi c Coast for four NALMAs. The modern sur-face- water curve is provided for comparison and continuity across time bins. Latitude and longitude were approximated using www.topozone.com (Maps a la Carte, 2006) and the coor-dinate distance calculator (see Fig. 2) for stream and river water δ18O data from Kendall and Coplen (2001). Shaded area indicates possible location of the Sierran crest over the past 5 m.y. where the edge closest to the coast is the approximate location of the modern range.



as well (Bryant and Froelich, 1995; Kohn, 1996). We have noted that production of meta-bolic water, which incorporates atmospheric O

2

into the body water pool, causes bioapatite δ18O values to vary less than ingested water δ18O val-ues. We attempted to minimize the impact of such metabolic differences by sampling a single higher taxon (equids), assuming that they would have roughly similar physiology and ecology. When we included other taxa, however, we dis-covered that they were not signifi cantly different from equids; therefore, we presume the effects of metabolic and ecological differences are rela-tively minor among these taxa.

The main way that ontogenetic differences would affect our data is through analysis of teeth that mineralize at different points in an animal’s life. As noted above, potential nursing effects might infl uence δ18O values from teeth that form before weaning (M1 and M2) (Bryant, 1995), although recent work has suggested that these effects are minor (Kohn et al., 1998, 2002; Gad-bury et al., 2000). We tried to sample only P2 and M3 (which form after weaning), but there were several cases where we had to sample other cheek teeth. Furthermore, since P3, P4, M1, and M2 are indistinguishable as isolated teeth, it is possible that some additional specimens were inadvertently M1 and M2. Regardless, since δ18O values of cheek teeth that were possibly M1 or M2 are generally within the same range as P2 and M3 from the same locality, nursing effects are likely negligible (Table DR1).

With a few exceptions, a consistent intra-local-ity spread in δ18O values of 1‰–3‰ appears even when all of the teeth are M3 and P2 (Table DR1). The magnitude of variation at these sites is smaller or similar to that seen in modern terres-trial mammal populations and other fossil equid populations (Chillón et al., 1994; Clementz and Koch, 2001; Hoppe et al., 2004a). Based on the consistency of our data, ontogenetic differences do not appear to systematically bias the δ18O val-ues of bioapatite in our specimens.

MigrationNumerous studies have shown that natural

variations in the isotopic value of animal tissues can be used to study migration (van der Merwe et al., 1990; Vogel et al., 1990; Koch et al., 1995; Hobson, 1999). The progressive westward col-lapse of the Sierran crest, as well as limited high elevation sites, makes it diffi cult to examine the possibility of past migratory behavior. Migration across the range would tend to erase isotopic dif-ferences caused by a rain shadow; therefore, the persistence of a δ18O gradient across the range through time, as well as the absence of outliers, strongly suggests that migration effects are minor compared to the rain shadow signal. Analysis of

strontium isotopes, which are known to vary with bedrock geology in specimens at different localities, would help identify the presence or absence of migration (Koch et al., 1995).

DiagenesisWhen dealing with fossil chemistry, dia-

genesis must always be considered. Although enamel bioapatite is more resistant to alteration than bone or dentine bioapatite, enamel can still undergo signifi cant alteration (Chillón et al., 1994; Koch et al., 1997; Zazzo et al., 2004a, 2004b). The δ18O value of the phosphate fraction in bioapatite has been considered more resistant to diagenetic alteration than the carbonate frac-tion. However, Zazzo et al. (2004a, 2004b) have shown that the phosphate fraction of hydroxy-apatite is more susceptible than carbonate to microbial alteration, whereas carbonate is more susceptible to inorganic alteration. Depending on the type of alteration, diagenesis can affect some fossils and not others, increasing variabil-ity among the individuals at a locality (Chillón et al., 1994). Alternatively, diagenesis can homog-enize isotopic values among fossils at a locality, reducing variability (Zazzo et al., 2004b).

Several of the specimens we analyzed have δ18O values that are 3‰–4‰ different from other individuals at their locality (Table DR1). For example, despite replicate runs, one Neo-hipparion from Fish Lake Valley, Nevada, has a δ18O value of 15.2‰, which is almost 5‰ lower than the average for the seven other Neo-hiparrion sp. at this site (19.3‰). The range for this locality is 6.5‰. Similarly, Equus sp. from the Coso Mountains (localities V-5691–V-5694), California, have a range of 5.5‰ (δ18O = 18.1‰–23.8‰), and Merychippus sp. from Bar-stow, California (Locality 1399), have a range of almost 5‰ (δ18O = 23.2‰–28.4‰). This high range in δ18O values could be the result of selec-tive diagenesis. However, other studies have found that bulk enamel δ18O values at a single locality can vary by up to 6.5‰ (Clementz and Koch, 2001; Hoppe et al., 2004a). It has also been shown that multiple samples from a single tooth can vary by up to 7‰ (e.g., Sharp and Cer-ling, 1998).

On the other hand, δ18O values for Heming-fordian Merychippus sp. and Parahippus sp. at Massacre Lake, Nevada, are remarkably similar. With the exception of one Parahippus sp., δ18O values range from 19‰ to 21‰ (n = 9) (Table DR1). Because most other localities show an intra-taxon spread of 3‰–4‰, it is pos-sible that these Hemingfordian specimens were all altered and now refl ect the isotopic value of pore fl uids. However, our modern horse samples from Nevada only show an intra-locality range of ~1.5‰, so perhaps these fossil specimens

have not been altered. Regardless, fossil δ18O values from these eastern sites do not overlap with those from any western site (Fig. 4).

Following the methodology of Koch et al. (1997), we pretreated our samples with acetic acid to remove exogenous carbonates. Unfortu-nately, this pretreatment cannot remove apatite that has recrystallized or exchanged carbonate with the diagenetic environment. To assess if diagenetic alteration has occurred, it is necessary to compare the δ18O values of the phosphate and carbonate fractions (Zazzo et al., 2004a, 2004b). We plan to analyze the phosphate fraction of a suite of our specimens, but those analyses are beyond the scope of this paper.

Factors that Complicate Data Interpretation

Several factors other than elevation can affect the δ18O value of surface water. These include: (1) distance from the source of moisture, (2) shifts in seasonal precipitation, (3) other mountain ranges, and (4) Late Cenozoic climate change.

Distance from the Source of MoistureStream water δ18O values decrease between

the coast of California and the Sierran foothills and again across the Sierran range (Fig. 1). 18O-depleted water that falls either as liquid or snow at higher elevations recharges rivers and groundwater (Rowley et al., 2001). This water then moves to lower elevations, continuously mixing with less 18O-depleted surface waters as it travels. Transects across the Sierras during the Barstovian, Clarendonian, and Hemphillian also show a decrease in surface-water δ18O val-ues (inferred from enamel δ18O) from the Pacifi c Coast up to the Sierran foothills (Fig. 5). We argue that the large isotopic gradients from the coast to Nevada during the past are most consis-tent with a substantial topographic barrier in the Sierras back to 16 Ma. This conclusion implic-itly assumes that distance or the “continental” effect (Dansgaard, 1964) would be negligible. Could an isotopic gradient as steep as the one we observe as early as the Miocene exist if there was not substantial topography in the Sierras?

To evaluate this idea, we examined the shifts in the mean annual meteoric water δ18O value with increasing distance from a precipitation source in three transects to determine how much distance is required to shift δ18O by several per mil. These transects cross areas with low topo-graphic relief, including temperate and tropi-cal regions. Two of these transects, one across the Amazon Basin and one from the northern Atlantic Ocean to Germany, run roughly paral-lel to known storm tracks and isotherms. The third transect, from the Gulf of Mexico north to

Crowley et al.


Kansas, crosses isotherms. Inspection of these transects reveals that distance effects are rela-tively small, even when crossing isotherms. For example, in Brazil, it takes 3000 km to reduce the δ18O value of precipitation by 3‰. In Europe (northern France through Germany), δ18O shifts less than 2‰ over 1000 km.

Fewer than 800 km separate the California coast from the eastern edge of Nevada. Due to Basin and Range extension, this distance was smaller several Ma. Although surface water shifts more gradually than precipitation across the modern Sierras, the overall shifts in the δ18O value of precipitation and meteoric water are similar. Thus, decreases in the δ18O value of ~6‰ across the range as early as the Bar-stovian suggest that a factor other than distance caused this isotopic offset. This sharp decrease in δ18O values from west to east is similar to that observed in modern surface-water data, which results from orographic effects.

Changes in the Source or Seasonality of Precipitation

Friedman et al. (2002a) identify fi ve modern storm trajectories for sites in the Great Basin east of the Sierras. These storm tracks are from the Gulf of Alaska and North Pacifi c, Central Pacifi c, Tropical Pacifi c, Gulf of Mexico, and west across the North American continent. However, they argue that at present most mois-ture moves from west to east, and ~75% of the precipitation in Nevada comes from the Central Pacifi c and the Gulf of California during the winter, when strong winds push air masses over the crest of the Sierras (Friedman et al., 2002a, 2002b; Smith et al., 2002). There is also a sea-sonal bias in surface-water recharge, with 18O-depleted snow responsible for most recharge (Kendall and Coplen, 2001; Friedman et al., 2002b). Bias toward winter precipitation in the annual water budget may be increased further because summer precipitation is quickly cycled through plants and returned to the atmosphere via evapotranspiration.

The likelihood is small that storm intensity and seasonality have remained constant through time. As noted above, conditions were likely warmer and wetter in the Great Basin prior to 5 Ma. More mesic conditions could have been achieved by: (1) an increase in winter pre-cipitation, still largely derived from the Pacifi c, (2) increased summer precipitation, also sourced from the Pacifi c, or (3) increased summer precip-itation from the Gulfs of California or Mexico.

The fi rst two options would not affect our interpretation of paleoelevation; most moisture would move across the range, as it does today. However, because moisture originates south of the Sierras in the third alternative, it would

not travel across the range. Because southern water sources have higher δ18O values, if they made a greater contribution in the past, the δ18O gradient across the range in surface water would be smaller, causing us to underestimate paleoelevation. Evidence for arid conditions in the Miocene in the Mojave Desert and south-ern Basin and Range (Horton and Chamber-lain, 2006) suggest that the moisture for this precipitation is not from the south. In any case, because we sampled both sides of the range through time, our approach should at least sup-ply a minimum estimate of paleoelevation rela-tive to the modern.

Effects of Other California RangesWe briefl y consider how uplift of the Califor-

nia Coast and Transverse Ranges could affect our interpretation of the δ18O gradients across the Sierra. The timing of uplift for both of these ranges is unclear. Some researchers argue that the Coast Ranges were formed as a result of changes in the plate motion between the Pacifi c and North American plates from 5 to 3.2 Ma (Pollitz, 1986; McIntosh et al., 1991; Atwater and Stock, 1998). Others hold that there is no evidence for change in plate boundary motion since 8 Ma, thereby implying that the Coast Ranges are at least this old (Nicholson et al., 1994). With respect to the Transverse Ranges, Atwater and Stock (1998) argue that rotation and uplift has occurred since 8 Ma and continues today, whereas Nicholson et al. (1994) maintain that rotation may have begun as early as 20 Ma. At fewer than 1500 m, the Coast Ranges are low relative to the Sierras. The Transverse Ranges, on the other hand, have elevations that exceed 3500 m. Yet modern surface-water data indi-cate a shift of less than 2‰ across both of these ranges (Fig. 1).

Regardless, there is evidence that the Trans-verse Ranges affect airfl ow patterns (Smith et al., 1979). Although most moisture-laden air fl ows over the Sierras, some refracts around the southern end of the Sierras rather than passing over the crest. It is possible that the Transverse Ranges are responsible for this defl ection. As a result, large storms can bring precipitation around the Sierran range, creating slightly heavier surface-water δ18O values than expected in areas directly east of the range (Fig. 1).

We lack the fossil data necessary to recon-struct δ18O gradients across these smaller ranges during the Miocene and Pliocene (Fig. 3). The fact that Hemingfordian δ18O values from San Bernardino County correspond well with the other western data, however, suggests that the Transverse and Coastal Ranges had only a minor effect on δ18O values west of the Sierras at that time (Fig. 4).

Late Cenozoic Climate ChangeA clear difference in enamel δ18O values is

apparent between our localities on the east and west sides of the Sierra Nevada from roughly 16–2 Ma. Our most recent fossil specimens, however, do not show this east-west dichotomy, and their δ18O values are not statistically distin-guishable (Fig. 4). Glacial climate cycles may have contributed to this overlap. Changes in the δ18O value of seawater during glacial cycles will cause the δ18O value of meteoric water to differ between interglacial and glacial periods. More important, the δ18O values of precipitation and surface water are linked to surface air tem-peratures. As much as 50% of long-term change in meteoric water δ18O values can be directly related to surface air-temperature changes ( Araguas Araguas et al., 1996).

Our inability to resolve climatic variables and specimen ages to differentiate between glacial and interglacial episodes might explain some of the overlap in δ18O values across the Sier-ras in the Plio-Pleistocene specimens (Fig. 4). Huber (1981) identifi es two major Sierran gla-cial episodes between 3.1 and 2.8 Ma and 2.7 to 0.7 Ma. Others date the inception of glaciation at ca. 2.5 Ma (Stock et al., 2005). These inferred periods of glaciation coincide with proposed Late Cenozoic uplift (e.g., Huber, 1981; Unruh, 1991). This overlap in factors that affect enamel δ18O values prevents us from distinguishing between climatic and orographic isotopic sig-nals during this time.

SUMMARY

Stable isotope data suggest that the Sierra Nevada Mountains have had substantial relief (similar to modern) since at least 16 Ma. Our fi ndings do not support hypotheses attributing most of the current topography to uplift between 10 and 2 Ma. However, some workers who argue for a Late Cenozoic phase of uplift propose an increase of only 1–2 km (e.g., Jones et al., 2004; Stock et al., 2004). Because of the high variabil-ity in our data from this time period, we cannot resolve uplift (or downdrop) of this magnitude. Regardless, recent uplift must have been rela-tively small compared with pre-Miocene uplift, because the isotopic gradient does not seem to have increased between 2 and 0 Ma.

Some studies have suggested that higher ele-vations, in the form of a plateau, once existed to the east of the Sierra Nevada (Forest et al., 1995; Wolfe et al., 1998; House et al., 2001; Horton et al., 2004; Abruzzese et al., 2005; Mulch et al., 2006). This claim is supported by evidence from magmatic intrusion and associ-ated deformation in Nevada, which argues for an elevation increase of roughly 2 km during



this period (Horton et al., 2004). The idea of high topography extending east of the Sierras and across the Basin and Range is compelling. However, because our study tests only for the existence of a rain shadow, we are only able to detect the presence of an orographic barrier, not its width.

Additional data are needed to remove the remaining uncertainty in isotopic stud-ies of paleoelevation. The recently developed “clumped isotope” thermometer (Ghosh et al., 2006a) would allow us to deconvolute changes in temperature from those in surface-water δ18O values in studies of soil carbonates. The approach has already been used to study the paleoeleva-tion of the Altiplano (Ghosh et al., 2006b). By comparing isotopic data from sources that might have experienced strong evaporation to those more likely to track meteoric water, it may be possible to assess regional moisture levels. By combining these three sources of information, it may be possible to discriminate between high, cold, and dry plateaus, such as the hypothesized “Nevadaplano” (DeCelles, 2004) and low, hot, and dry basins that sit in rain shadows, such as modern day Nevada.

ACKNOWLEDGMENTS

We thank Pat Holroyd for access to collections of the University of California Museum of Paleontol-ogy, Kirsten Pasero and John Neil from the Bureau of Land Management’s Wild Horse and Burro Program, and Robert Feranec, Kathryn Hoppe, Mark Clementz, Rebecca Zisook, Kena Fox-Dobbs, Greg Stemler, Page Chamberlain, and Gabe Bowen. This research was funded by National Science Foundation grant EAR-0309383 to PLK.

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Stable isotope constraints on the elevation history of the Sierra Nevada Mountains, California

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