Climatic and ecologic changes during Miocene surface uplift in the Southern Patagonian Andes

Earth and Planetary Science Letters 230 (2005) 125 – 142 www.elsevier.com/locate/epsl

Climatic and ecologic changes during Miocene surface uplift in the Southern Patagonian Andes Peter M. Blisniuka,*, Libby A. Sternb, C. Page Chamberlainc, Bruce Idlemand, Peter K. Zeitlerd a

Institut fu¨r Geowissenschaften, Universita¨t Potsdam, Postfach 60 15 53, D-14415 Potsdam, Germany b Department of Earth Sciences, University of Texas, Austin, Texas, USA c Department of Geological and Environmental Sciences, Stanford University, Stanford, CA, USA d Department of Earth and Environmental Sciences, Lehigh University, Bethlehem PA 18015, United States Received 17 May 2004; received in revised form 22 October 2004; accepted 15 November 2004 Available online 7 January 2005 Editor: V. Courtillot

Abstract The up to ~4 km high southern Patagonian Andes form a pronounced topographic barrier to atmospheric circulation in the southern hemisphere westerlies, and cause one of the most drastic orographic rain shadows on earth. Geologic data imply that this climatic pattern has been established or significantly enhanced during Miocene surface uplift of this Andean segment. We report evidence for important climatic and ecologic changes in the eastern foreland of the Patagonian Andes that appear to be the result of this uplift. To provide constraints on Miocene plant ecosystems and precipitation in the eastern (leeward) foreland of the Patagonian Andes, we determined carbon and oxygen isotope values of pedogenic carbonate nodules from a ~500 m thick section of the continental Santa Cruz Formation. The age of these deposits was constrained by Ar/Ar dating of intercalated tuffs, which range from ~22 to 14 Ma. At ~16.5 Ma, the y13C values increase by ~3x, the y18O values decrease by N2x, and the scatter in the oxygen isotope data increases significantly. We interpret these changes as the consequence of N1 km surface uplift in this Andean segment (from the y18O values), and increased aridity to its east (from the y13C values and the increased scatter in the y18O values). Sediments overlying the Santa Cruz Formation are very limited in extent and volume, and dominated by coarse conglomerates related to Pleistocene and older glaciations. It thus seems that, by ~14 Ma, deposition in the eastern foreland of the Southern Patagonian Andes had essentially ceased as the result of rain shadow formation. D 2004 Elsevier B.V. All rights reserved. Keywords: stable isotopes; pedogenic carbonate; rain shadow; paleoclimate; paleoelevation; Miocene; Andes; Patagonia

* Corresponding author. Tel.: +49 331 977 5408; fax: +49 331 977 5060. E-mail address: [email protected] (P.M. Blisniuk). 0012-821X/$ - see front matter D 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2004.11.015

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1. Introduction and geologic setting Tectonically controlled surface uplift in mountain belts can cause important local and regional-scale climatic and ecologic changes, which, in turn, affect the distribution and rates of erosion, sediment transport, and deposition (e.g., [1]). However, few previous studies have examined these interactions in detail, and consequently our quantitative understanding of the links between these processes is limited. The Patagonian Andes south of the point of subduction of the Chile Ridge (Figs. 1 and 2) are ideally suited for a study of these interactions. Today, these up to ~3700 m high mountains form a pronounced topographic barrier to atmospheric circulation in the southern hemisphere westerlies, and cause one of the most drastic orographic rain shadows on earth (Fig. 3); rainfall in the humid western foreland is N3000 mm/yr, contrasting with only ~300 mm/yr to the east of the mountains [2]. Geologic

data imply that this climatic pattern has been established or significantly enhanced during Cenozoic uplift of the southern Patagonian Andes: on the basis of plate geometry reconstructions, Cenozoic uplift in this region has been related to two episodes of oceanic ridge subduction beneath the South America plate, (1) the subduction of the Aluk-Farallon spreading center during the Paleogene, and (2) the subduction of several segments of the Chile Ridge since ~15 Ma (Fig. 2) [3,4]. The main factor controlling uplift of the southern Patagonian Andes, however, appears to have been a strong increase in the convergence rate along with a decrease in convergence obliquity between the Nazca and South America plates at 26–28 Ma [5–7]. Apatite fission track data imply that increased denudation started at ca. 30–23 Ma in the region near the Pacific coast and subsequently migrated ~200 km eastward to the region of the present-day topographic axis of this Cordilleran segment until ~12 Ma, most likely as the result of subduction erosion [8]. This

46˚ ATLANTIC OCEAN

ice

47˚ PACIFIC OCEAN

48˚

ice

49˚

76˚

74˚

72˚

70˚

68˚

Cretaceous and Cenozoic granitoid intrusives

Holocene alluvium

Taitao Ophiolite

Pleistocene alluvium

Mesozoic rocks

Pleistocene glacial deposits

Paleozoic rocks

Plio- Pleistocene gravels

thrust fault

Miocene to recent basalts

studied section

Miocene sediments

lake

Paleogene sedimentary and volcanic rocks

66˚

Fig. 1. Generalized geologic map of the Southern Patagonian Andes (after [8,64]). The location of the map is indicated by the black box on the inset in the lower left. The location of studied sediment section is indicated by the circled asterisk symbol.

P.M. Blisniuk et al. / Earth and Planetary Science Letters 230 (2005) 125–142 18 Ma

14 Ma

10 Ma

6 Ma

127

present 6.7

9.0 45˚

50˚

2.0

90˚

85˚

80˚

75˚

70˚

65˚

55˚

Fig. 2. Sketch map showing the geometry of the present-day subduction zone and segments of the active Chile spreading ridge (black lines) between the Nazca Plate in the East and the Antarctic Plate in the West. Grey lines indicate reconstructed approximate positions of Chile Ridge segments at ~18 Ma, ~14 Ma, ~10 Ma, and ~6 Ma; arrows indicate the present-day spreading rate at the Chile Ridge and convergence rates along the subduction zone (after [3,4]). Note that the convergence rates are much lower in the south compared to north of the point of ridge collision. The location of the studied sediment section is indicated by the circled asterisk symbol.

evidence for increasing uplift and denudation since the Oligocene is in good agreement with the stratigraphic sequence in the eastern foreland basin, which includes Eocene and Miocene to Pleistocene plateau basalts that have been related to slab windows associated with the Cenozoic ridge subductions, and thick Oligocene to recent syn-orogenic clastic (molasse) deposits (Fig. 1) [4,9,10]. The lower part of the molasse deposits is represented by the Oligocene to early Miocene(?) Centinela Formation, a series of near-shore marine conglomerate, sandstone, and shale beds deposited after a transgression from the Atlantic [4,11]. These marine beds are conformably followed by early to middle Miocene fluvial deposits of the Santa Cruz Formation, which have been related to the main phase of Cenozoic deformation and surface uplift in the Cordillera to the west [4,9]. Based on K–Ar ages of tuffs from several scattered outcrops of the Santa Cruz Formation, the age range of these deposits has been estimated at ~19 to ~15 Ma [12,13]. A ~14 Ma minimum age is implied by the 13.9F0.3 Ma age of the oldest overlying plateau basalt [10]. Sediments post-dating the Santa Cruz Formation are relatively limited in extent and volume, and dominated by coarse con-

glomerates related to Pleistocene and older glaciations [14,15]. The geologic data implying important early to middle Miocene surface uplift in the southern Patagonian Andes are independently supported by paleontologic evidence. Mammal faunas from the richly fossiliferous Santa Cruz Formation, which is the basis for recognition of the Santacrucian Land Mammal Age of South America, are taxonomically the most diverse in terms of known mammalian genera of all Tertiary faunas on this continent [16,17]. A comparison of Santacrucian faunas with those of younger land mammal ages indicates a radical change from a subtropical humid climate to a cooler and much more arid climate in the eastern foreland [18– 20]. This climate change was apparently accompanied by a transition from balanced subtropical woodlands and grasslands to predominantly grasslands, since most of the climate-sensitive mammal taxa suggesting subtropical woodlands, including primates, became extinct, rare, or absent in southern South America during the transition from the Santacrucian to younger land mammal ages [18–20]. Plausible causes for these drastic climatic and ecologic changes include both global and local

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-6

SURFACE WATER SAMPLES

-10

precipitation non-evaporated spring/stream evaporated spring/stream

-12

δ18OSMOW (‰)

-8

-14 -16

West

East

2000

Elevation [m]

4000

Precipitation [mm/yr]

4000 6000

studied sediment section

precipitation

Maximum elevation

3000 2000

Mean elevation

Minimum elevation

76˚

75˚

74˚

73˚

72˚

71˚

70˚

69˚

Fig. 3. Diagram illustrating the present-day rain shadow effect across the Southern Patagonian Andes. Top: oxygen isotope values of present-day surface waters collected along a transect of the Southern Patagonian Andes. Note that, compared to samples from the windward western side of the mountains, surface waters from the leeward eastern side have significantly lower y18O values. Also note that deuterium excess values demonstrate that a relatively high proportion of surface waters east of the mountains has experienced substantial evaporation, which causes a y18O increase. As discussed in Stern and Blisniuk [39], east of the mountains waters classified as non-evaporated may have experienced some evaporation; in that part of the transect the lower values can be considered to be from the least evaporated waters and thus the best proxy for precipitation. Bottom: maximum, mean, and minimum elevation along an east–west oriented and ~70 km wide swath between 47820VS and 48800VS, which contains water sampling locations and the studied sediment section. Stippled bold line indicates mean annual rate of precipitation along the swath [2]. Modified after Stern and Blisniuk [39].

factors. The decreasing temperatures implied by the fossil record [19] may be related to gradual global cooling since ~14.5 Ma, which was associated with rapid Antarctic ice sheet growth and major biogeographic changes on all continents [21–23]. The increase in aridity in the eastern foreland has been attributed to this cooling, and also to a progressive elevation increase in the Andean Cordillera farther west, which presumably established a pronounced orographic rain shadow by the middle Miocene [13,18,19]. To better resolve the nature, timing, and causes of these environmental changes, we carried out a detailed study of Santa Cruz Formation deposits, using (1) carbon isotope analysis of pedogenic (soil-formed) carbonate nodules in these deposits, to reconstruct local plant ecosystems, (2) oxygen isotope analysis of these nodules, to track elevation changes in the

Andean segment immediately west of the studied section, and (3) 40Ar/39Ar dating of feldspars from intercalated tuff beds, to establish a chronologic framework for these data.

2. Background for stable isotope studies Carbon and oxygen isotope values of pedogenic carbonate, which generally forms in regions with less than 750–1000 mm annual precipitation, are valuable indicators of environmental and climatic conditions. The y13C value of pedogenic carbonate is largely controlled by the average y13C of vegetation growing in the soil, which in turn is largely governed by the relative proportion of C3 plants to C4 plants during pedogenesis [24]. C3 plants, which include trees, most shrubs and herbs, and cool season grasses, are typical

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of relatively humid and temperate climates, whereas C4 plants are predominantly summer grasses, which are adapted to high light and water-stressed conditions, and typical of more arid climates [24]. The y13C values of pedogenic carbonate formed in soils associated with pure C3 plant ecosystems can range from 22x to 8x, and are typically between 12x and 10x [24,25]. Values near the lower limit of 22x are typically associated with soils formed under a closed forest canopy [24], whereas values near the upper limit of 8x are indicative of semi-arid environments where C3 plants can be under significant moisture stress [25,26]. In soils associated with pure C4 plant ecosystems, pedogenic carbonate typically has y13C values of 0x to +2x, and y13C values between 8x and 0x are generally attributed to mixed C3/C4 ecosystems [24,27]. The y18O value of pedogenic minerals is controlled by the isotopic composition of the soil water present during pedogenesis (and, to a lesser degree, on soil temperature), which in turn largely depends on the y18O value of precipitation, and on the amount of soil water evaporation [28]. Due to a strong systematic decrease in the y18O values of precipitation with elevation [29], y18O values of pedogenic minerals are thus potentially useful for reconstructions of paleo-topography. The main effect of topography on precipitation is related to the progressive condensation and removal of precipitation from an uplifting airmass. As a moisture-bearing airmass ascends it undergoes pseudo-adiabatic cooling and condensation, which partitions more of the heavy isotope into the condensed phase. After passing over a mountain belt an airmass thus not only retains less water than on the windward side, but is also depleted in 18O. Accordingly, the y18O value of pedogenic minerals can, in appropriate settings, be used to reliably constrain paleo-topography [30–37]. The southern Patagonian Andes are ideally suited for the use of oxygen isotopes to reconstruct paleotopography because they are a narrow mountain belt located immediately east of the Pacific ocean, and in the center of the southern hemisphere westerlies. Therefore, precipitation is derived from a single dominant moisture source (the Pacific Ocean), and topography has a relatively simple influence on the patterns of atmospheric circulation. In fact, a

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comparison of GCM simulations with and without the high topography of the Andes indicates that summer precipitation in the Southern Andes is purely orographic [38]. This is in good agreement with data from present-day precipitation and surface waters sampled along a transect of the southern Patagonian Andes. On the windward western side of the mountains, an eastward increase of the maximum windward elevation by ~2 km is accompanied by a N4x decrease of the y18O values of meteoric waters (Fig. 3) [39], consistent with an isotopic lapse rate estimate of 0.28x/100 m derived from a global compilation of precipitation and surface water samples [34]. In the rain shadow on the leeward side of the mountains, however, the y18O values of precipitation and surface waters show no significant trend with local elevation [39]. This implies that the y18O values of precipitation in that region are largely controlled by the progressive condensation and rain-out of moisture on the windward western side of the southern Patagonian Andes, and that the oxygen isotope composition of paleosol carbonates on their leeward (eastern) side should reflect changes in the surface elevation of this cordilleran segment according to the isotopic lapse rate estimated at 0.28x/100 m [34,39]. We also note that both the westerly winds and the South American plate have not undergone significant latitudinal motion since the Miocene [40,41]. Accordingly, the westerlies have been the primary source of precipitation in the area throughout the time interval relevant for our study, making interpretations of changes in the isotopic composition of precipitation through the geologic record relatively straightforward.

3. Description of studied section The studied section of Santa Cruz Formation deposits is located near Lago Posadas (Fig. 1). The choice of this section was based on the unique quality and extent of exposure at that location, which significantly exceeds that at any other site at which the Santa Cruz Formation was either previously described, or observed by us during extensive reconnaissance in the eastern foreland between 478S and 508S. At this location a ~507.5 m thick section of

sediment thickness [m]

B

D

C 13

δ C (PDB) [‰]

14.24 ± 0.78 Ma

500

500

130

A

18

δ O (SMOW) [‰] 500

500

400

400

300

300

200

200

100

100

136.2 ± 94.4 m/Myr 400

0 ± 24.5

84.9 ± 43.4 m/Myr

300

averag

ed 114.

300

15.51 ± 0.41 Ma

271.2 ± 448.5 m/Myr 200

16.45 ± 0.25 Ma

200 16.71 ± 0.63 Ma

85.0 ± 27.6 m/Myr 100

100

18.15 ± 0.31 Ma 11.6 ± 2.2 m/Myr

22.36 ± 0.73 Ma

0 24 23 22 21

20 19

18 17 16

0 ~30 m covered interval

15

sediment age [Myr]

OLIGOCENE

-15

-14 -13 -12 -11

-10

-9

-8

-7

-6

-5

-4

-3

-2

-1

0 13

14

15

16

17

18

19

20

21

22

average from several nodules

average from several nodules

single sample

single sample

5-pt running average

5-pt running average

23

dated tuff conglomerate sandstone

volcanic ash/tuff siltstone

Fig. 4. Isotopic and geochronologic data from the studied sediment section. (A) Plot of stratigraphic position versus age of dated tuffs from the Santa Cruz Formation (1r age uncertainties indicated by error bars). A minimum age for the top of the section (which may have experienced some erosion) is provided by a 14.2F0.8 Ma age for a tuff sampled 3 m beneath the top of the Santa Cruz Formation at a locality ~10 km to the SE of the studied section, where the Santa Cruz Formation is overlain by basalts dated as ~12.1F0.7 Ma [42]. Numbers between points of age control are sediment accumulation rates for corresponding time intervals. Bold line to right of graph indicates average sediment accumulation rate between ~18 and ~14 Ma. Sediment accumulation rates are not corrected for compaction. (B) Stratigraphic log of the studied sediment section, with positions and ages of dated tuffs indicated to the right of the section. (C) Carbon isotope data of pedogenic carbonate nodules in paleosols contained in the studied section. Paleosols from which several individual nodules were analyzed are shown by filled symbols, with error bars indicating one standard deviation. For symbols without error bars, one standard deviation is smaller than width of symbol. Paleosols from which only a single nodule was analyzed are indicated by open symbols (see Tables 2 and 3). Note solid dashed line highlighting a y13C value of 8x, with higher values strongly indicative of the presence of a significant proportion of C4 vegetation. (D) Oxygen isotope data of pedogenic carbonate nodules in paleosols contained in the studied section. Use of open and filled symbols as for (C).

P.M. Blisniuk et al. / Earth and Planetary Science Letters 230 (2005) 125–142

m/Myr

400

P.M. Blisniuk et al. / Earth and Planetary Science Letters 230 (2005) 125–142

subhorizontally bedded Santa Cruz Formation deposits is exposed, following on older Oligocene deposits of the Centinela Formation after a ~30 m covered interval (Fig. 4B). The top of the exposed section is the surface of a relatively flat-topped mountain which may be a relict land surface, but we cannot exclude the possibility that the top portion of the Santa Cruz Fm. deposits has been eroded. In a locality ~10 km to the SE of the studied section, however, the thickness of (less well exposed) Santa Cruz Fm. deposits, between underlying deposits of the Centinela Fm. deposits and overlying Miocene basalts, is V600 m, implying that the studied section is at least almost complete. The Santa Cruz Formation is a sequence of fluvial deposits dominated by alternating sand-, silt-, and claystone beds. Individual beds are laterally continuous for tens to hundreds of meters, and between a few tens of centimeters and several meters thick. The base of sandstone beds is sometimes a slight erosional unconformity, but commonly observed desiccation cracks and well-preserved soil structures near the tops of underlying beds attests to a predominantly aggradational nature of these deposits. The presence of trough cross bedding, lateral accretion structures, fining upward cycles and clayrich sandstone beds is suggestive of deposition by a meandering stream system with frequent flooding events, whose overbank deposits are represented by the siltstone and claystone beds. The lowermost ~110 m of the section is dominated by silt- and claystone beds, whereas the younger part of the section contains more, thicker, and coarser sandstone beds. In the uppermost ~100 m of the sequence, conglomeratic sandstones and coarse conglomerates are present, including a 13 m thick debris flow deposit ~100 m below the top of the studied section (Fig. 4B). At several other locations in the study area, up to 5 m thick coarse conglomerate beds were observed at or near the top of the formation. In addition to clastic deposits, the studied section also contains numerous tuff and ash beds that are up to ~2 m, and typically a few centimeters to tens of centimeters thick. Numerous paleosol horizons were observed in all lithologies, but mostly contained in the silt- and claystone units, and are particularly common in the uppermost 100 m of the studied section.

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4. Geochronologic data 4.1. Methods To provide age control and constrain sedimentation rates for the studied deposits, we dated 5 tuffs contained in the studied section. Additionally, we dated a tuff from 3 m beneath the top of the Santa Cruz Fm. at a locality ca. 10 km farther SE, where these rocks are overlain by basalts dated as ~12.1F0.7 Ma [42]. Most of the samples were disintegrated in an ultrasonic bath with deionized water. Several of the more highly indurated tuffs were lightly crushed in an agate mortar. The samples were then sieved, feldspars were handpicked from the 500–1,000 Am size fraction, and the separates were cleaned with 10% acetic acid and acetone. The purified samples were irradiated for 2 h in position 5C of the McMaster University research reactor. CaF2 and K2SO4 were included in the irradiation package to monitor neutron-induced interferences from Ca and K, respectively. For each sample, we obtained 2–4 ages by fusing single grains or, in most cases, groups of 2–6 grains using a CO2 laser. The evolved Ar was purified using Zr– Al getters operated at 400 8C, and Ar analyses were performed with a VG3600 noble gas mass spectrometer equipped with an electron multiplier operated in the analog mode. The mass spectrometer sensitivity was ~1.510–17 mol/mV 40Ar. Both the extraction line and mass spectrometer were operated under computer control using LabSpec, a custom LabVIEW program developed at Lehigh University. The individual and weighted average ages obtained from the analyses of each tuff are presented in Table 1 along with information on standards and constants. Despite relatively low apparent K/Ca values suggesting that the crystals were calcic to intermediate plagioclase, radiogenic 40Ar yields were sufficiently high to yield consistent ages with reasonably low uncertainties. For sample LP 181.2, the MSWD (mean square of the weighted deviates) for the analyzed grains was unacceptably high, so the calculated uncertainty on the weighted mean was multiplied by the square root of the MSWD to obtain the reported uncertainty.

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Table 1 40 Ar/39Ar laser total fusion ages of tuffs in the Santa Cruz Formation 37

36

K/Ca

40

Ar* (%)

Age (Ma)b,c

F1r d

LP-Iz504.5 m e J=0.0004720F0.5% 6 55.040 0.08977 6 111.471 0.10950 6 113.094 0.11702

26.802 26.221 26.294

0.1385 0.3280 0.3367

0.016 0.016 0.016

29.88 15.07 14.01 Wtd. Ave.=

14.26 14.55 13.73 14.24

1.13 1.39 1.70 0.78

LP 331.5 331.5 m e J=0.0004720F0.5% 1 57.472 0.06706 1 47.310 0.09924

22.304 20.645

0.1416 0.1045

0.019 0.020

30.58 38.49 Wtd. Ave.=

15.18 15.70 15.51

0.67 0.51 0.41

LP 251.7 251.7 m e J=0.0004728F0.5% 1 43.361 0.03086 2 60.683 0.04077 2 74.869 0.05068

7.198 7.341 7.670

0.0838 0.1430 0.1889

0.059 0.058 0.056

44.29 31.37 26.29 Wtd. Ave.=

16.40 16.26 16.81 16.45

0.33 0.45 0.51 0.25

LP 181.2 181.2 m e J=0.0004720F0.5% 1 94.605 0.07740 1 200.573 0.15551 1 129.869 0.09601 1 119.095 0.09686

12.649 11.563 17.184 18.691

0.2541 0.6151 0.3857 0.3431

0.034 0.037 0.025 0.023

21.75 9.83 13.35 16.21 Wtd. Ave.=

17.62 16.87 14.91 16.61 16.71

0.49 1.00 0.69 0.82 0.63

LP 58.8 58.8 m e J=0.0004720F0.5% 4 79.607 0.06391 4 33.869 0.02947 4 72.296 0.09048

16.703 11.164 8.351

0.2012 0.0463 0.175

0.025 0.038 0.051

27.13 62.49 29.44 Wtd. Ave.=

18.55 18.1 18.16 18.15

1.14 0.4 0.5 0.31

LP 10 10.0 m e J=0.0004720F0.5% 6 150.826 0.12456 6 161.441 0.14244 6 165.432 0.14794

23.354 28.122 29.065

0.4302 0.4666 0.4806

0.018 0.015 0.014

17.03 16.08 15.65 Wtd. Ave.=

22.15 22.47 22.44 22.36

1.31 1.17 1.29 0.73

Sample (grains)

40

Ar/39Ara

38

Ar/39Ara

Ar/39Ara

Ar/39Ara

a

Corrected for mass spectrometer background, extraction line blank, mass discrimination, and radioactive decay of 37Ar and 39Ar. Mass discrimination (1 amu)=1.01047F0.2%. Average blanks: 40Ar=3.010 16 mol, 39Ar=3.010 18 mol, 37Ar=3.910 19 mol, 36Ar=9.610 19 mol. b Corrected for (1) above plus atmospheric argon and neutron-induced interferences. Interference corrections: (36Ar/37Ar)Ca=0.0002972, (39Ar/37Ar)Ca=0.0008094, (40Ar/39Ar)K=0.0180. c Ages calculated using the decay constants of Steiger and J7ger [62]. Neutron flux monitored with GA1550 biotite (98.79F0.54 Ma; [63]). d Uncertainties for weighted average ages include a 0.5% J-factor contribution. e Stratigraphic level above base of studied section. Sample LP-I was collected 3 m below a basalt flow overlying the Santa Cruz Formation in an outcrop ~10 km to the SE of the studied section; its age is interpreted to represent a minimum age for the stratigraphic level 504.5 m above the base (3 m below the top) of the studied section, where erosion may have occurred (see text).

4.2. Results The ages of the dated samples range from at ~22.4F0.7 Ma for a tuff from 10 m above the base, to 15.5F0.4 Ma for a tuff from 331.5 m above the base

of the studied section. In the younger part of the section we could not find datable tuffs. However, a reliable minimum age for the top of our section, which may have experienced some erosion, is represented by the 14.2F0.8 Ma age for the tuff from 3 m beneath the

P.M. Blisniuk et al. / Earth and Planetary Science Letters 230 (2005) 125–142

top of the Santa Cruz Formation at a locality ~10 km SE of the studied section, where these rocks are overlain by basalts dated as ~12.1F0.7 Ma [42]. The sediment accumulation rates implied by these ages are characterized by a drastic increase in the lower part of the section (Fig. 4A,B). Between 22.4F0.7 and 18.2F0.3 Ma (at 10 and 58.8 m above the base of the section, respectively), the mean sediment accumulation rate (not corrected for compaction) was 12F2 m/Myr. In the younger part of the section, sediment accumulation occurred at substantially higher rates; for example, between the 18.2F0.3 Ma tuff at 58.8 m above the base of the section and the z14.2F0.8 Ma top of the section, the mean rate of sediment accumulation was z114F25 m/Myr. Our data only bracket the main increase in deposition rates between 18.2F0.3 Ma (at 58.8 m in the section) and 16.7F0.6 Ma (at 181.2 m in the section); it is tempting, however, to speculate that the main increase in deposition rates may be marked by the change to more frequent and thicker sandstone beds at ~110 m in the section.

5. Stable isotope data 5.1. Methods The samples used for isotopic analysis are carbonate nodules from paleosol horizons with evidence for extensive pedogenic alteration (bioturbation, calcic and argillic zones, root traces, burrows). The studied nodules are from 5 mm to 3 cm, and typically ~1 cm in diameter. They were cleaned ultrasonically in two steps, first in 5% acetic acid, then in deionized water. To obtain the powders used for analysis, nodules b1 cm in diameter were crushed completely, whereas dental burrs were used to drill powders from representative transects of larger nodules. Sample analysis was carried out in the stable isotope laboratories at Dartmouth College, the University of Texas at Austin, and Stanford University. At Dartmouth and Austin, the sample powders were processed by phosphoric acid digestion at 25 8C [43] and analyzed on a Finnigan Delta E mass spectrometer (Dartmouth), or by reaction at 90 8C in a Multiprep system interfaced with a VG Prism II mass spectrometer (Austin). At Stanford, the samples were

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analyzed on a fully automated Finnigan Gas-Bench interfaced with a Finnigan Delta-Plus XL using phosphoric acid digestion at 72 8C. During the measurements, we routinely analyzed NBS (National Bureau of Standards) and internal laboratory standards, using the same gas extraction procedure as for the samples. For all measurements, we estimate the 2r uncertainty to be V0.1x for C and V0.2x for O (Table 2). The data are reported with respect to VPDB for carbon and VSMOW for oxygen (expressed in dnotation, with d=1000[R sample/R standard 1], where R is 13C/12C or 18O/16O). 5.2. Results The main trend in the y13C values (Fig. 4C) is a ~3x increase at ~16.5 Ma (between ~200 and ~215 m from the base of the measured section), following a general decrease in the older part of the section. Between 14.4 and 93.0 m from the base of the section, site averages range from 8.4x to 11.9x, with a mean of 9.8F1.0x (1r, n=14), whereas between 106.6 and 200.8 m, they range from 9.1x to 13.6x, with a mean of 11.8F1.4x (1r, n=19). Above and including the paleosol at 215.5 m from the base of the section, the site-averaged y13C values range from 3.7x to 12.8x, with a mean of 8.9F1.5x (1r; n=69). The increase at 200–215 m is very pronounced; a T-test comparing mean y13C values at stratigraphic levels from Table 3 yields a very significant difference ( p=3.810 5) between samples from 150–201 m versus 201–250 m. Similarly a T-test comparing y13C data from all stratigraphic levels below and above 201 m yields a p of 8.110 9. Compared to the y13C data, the y18O data (Fig. 4D) are characterized by a more gradual trend towards lower values throughout the studied section. From the lowermost 100 m to the uppermost 100 m in the section, the mean of the site-averaged y18O values decreases by ~2x from 19.2F1.3x (n=14) to 17.1F1.9x (n=34), respectively. This decrease, however, appears most pronounced between ~180 m and ~200 m from the base of the section. In the lower part of the section (from 14.4 m and 179.3 m above the base), the mean of the site-averaged y18O values is 18.9F1.2x (1r, n=32), ~1.5x lower than the corresponding value of 17.4F2.0x (1r, n=70) for

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Table 2 Results of stable isotope analyses of all individual carbonate nodules used for this study (n=268) Sample ID

Stratigraphic level (m)

y18O

LPC 14.4 144a 144b LPC 15.4 154a 154b LPC 19.3 193a 193b LPC 21.9 219a 219b LPC 22.5 225a 225b LPC 53.3 533a 533b LPC 60.7-1 LPC 60.7-2 LPC 69.2 692 LPC 78.7 787a 787b LPC 79.0 790a 790b LPC 81.6_av 816a 816b LPC 83.9 839a 839b LPC 86.7_av 867a 867b LPC 93.0_1_av LPC 93.0-2 LPC 93.0-3 LPC 93.0-4 LPC 93.0-5 LPC 106.6 1066a 1066b LPC111.2 1112a 1112b LPC 144.1 1441a 1441b

14.4 14.4 14.4 15.4 15.4 15.4 19.3 19.3 19.3 21.9 21.9 21.9 22.5 22.5 22.5 53.3 53.3 53.3 60.7 60.7 69.2 69.2 78.7 78.7 78.7 79.0 79.0 79.0 81.6 81.6 81.6 83.9 83.9 83.9 86.7 86.7 86.7 93.0 93.0 93.0 93.0 93.0 106.6 106.6 106.6 111.2 111.2 111.2 144.1 144.1 144.1

20.43 20.14 20.08 19.25 20.18 20.15 16.88 16.90 17.55 21.34 21.22 20.73 19.02 20.96 19.79 18.57 18.16 17.69 16.48 18.80 18.53 20.18 16.75 17.59 16.85 19.16 19.21 19.55 19.75 20.03 20.02 20.87 20.79 18.80 20.36 20.36 20.53 18.14 19.05 17.81 18.12 18.24 16.82 17.69 19.17 16.22 16.38 16.50 17.77 18.73 18.94

y13C 11.09 9.38 10.47 10.48 9.81 9.69 9.75 10.20 10.21 9.21 10.05 9.23 8.63 8.63 10.32 9.36 10.16 10.71 12.45 11.35 8.95 8.60 10.05 10.03 10.20 8.77 9.11 9.32 9.00 8.26 7.92 8.18 8.44 10.73 8.14 9.15 8.50 11.47 10.44 12.24 11.75 11.76 12.76 13.45 10.13 12.12 12.39 12.60 9.75 12.04 11.89

Table 2 (continued) Sample ID

Stratigraphic level (m)

y18O

LPC 144.6 1446a 1446b LPC 145.1 1451 1451b LPC 150.3 1503a 1503b LPC 150.6 1506a 1506b LPC 152.9 1529a 1529b LPC 154.5 1545a 1545b LPC 155.6 1556a 1556b LPC 159.1 1591a 1591b LPC 160.2 1602a 1602b LPC 160.9 1609a 1609b LPC 162.8 1628a 1628b LPC 164.7 1647a 1647b LPC 168.6-1 LPC 168.6-2 1686 1748 1748b LPC 179.3 1793 LPC 200.8-1 LPC 200.8-2 2008a 2008b LPC 215.5 2155a 2155b LPC 220.9 2209a 2209b

144.6 144.6 144.6 145.1 145.1 145.1 150.3 150.3 150.3 150.6 150.6 150.6 152.9 152.9 152.9 154.5 154.5 154.5 155.6 155.6 155.6 159.1 159.1 159.1 160.2 160.2 160.2 160.9 160.9 160.9 162.8 162.8 162.8 164.7 164.7 164.7 168.6 168.6 168.6 174.8 174.8 179.3 179.3 200.8 200.8 200.8 200.8 215.5 215.5 215.5 220.9 220.9 220.9

19.88 19.12 20.48 18.93 19.02 19.74 18.02 20.71 19.68 19.63 18.67 18.76 19.46 19.80 19.21 18.67 17.96 17.88 17.89 17.11 18.20 19.30 19.74 19.81 19.53 20.43 20.62 20.85 20.57 20.45 18.71 19.33 18.92 18.85 18.36 18.48 18.01 18.07 18.51 19.29 19.28 17.66 14.73 13.47 13.41 13.73 13.63 17.91 16.60 16.72 19.52 17.52 17.72

y13C 8.87 9.67 8.94 9.02 9.17 9.38 12.45 12.83 13.40 12.54 13.70 13.13 12.63 13.05 13.15 10.85 11.17 12.36 12.26 11.09 11.73 13.82 12.65 12.91 11.43 10.40 10.36 10.70 11.12 10.91 11.66 10.13 11.94 9.42 9.26 9.54 13.09 13.52 12.65 12.41 12.88 12.56 13.70 13.66 13.70 13.54 13.50 9.46 10.32 9.84 10.20 11.38 11.02

P.M. Blisniuk et al. / Earth and Planetary Science Letters 230 (2005) 125–142 Table 2 (continued)

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Table 2 (continued) 18

Sample ID

Stratigraphic level (m)

y O

LPC 222.7 2227a 2227b LPC 223.4 2234a 2234b LPC 224.7 2247a 2247b LPC 225.8 2258a 2258b LPC 226.8 2268a 2268b LPC 250.6 2506 LPC 252.1 2521a 2521b LPC 254.7 2547a 2547b LPC 261.2 LPC 282.0 2820a 2820b LPC 282.3-1 LPC 282.3-2 LPC 306.8-1 LPC 306.8-2 LPC 306.8-3 LPC 306.8-4 LPC 306.8-5 LPC 322.4 3224a 3224b LPC 322.6 3226a 3226b LPC 326.8 3268a 3268b LPC 332.0 3320 LPC 335.0 3350a 3350b LPC 335.5 3355a 3355b LPC 343.2_av 3432

222.7 222.7 222.7 223.4 223.4 223.4 224.7 224.7 224.7 225.8 225.8 225.8 226.8 226.8 226.8 250.6 250.6 252.1 252.1 252.1 254.7 254.7 254.7 261.2 282.0 282.0 282.0 282.3 282.3 306.8 306.8 306.8 306.8 306.8 322.4 322.4 322.4 322.6 322.6 322.6 326.8 326.8 326.8 332.0 332.0 335.0 335.0 335.0 335.5 335.5 335.5 343.2 343.2

18.67 17.06 16.57 16.45 16.20 16.62 16.21 17.73 16.41 17.41 16.02 17.84 18.25 17.47 17.51 19.59 14.65 15.68 15.76 14.83 18.04 15.49 15.78 18.97 21.22 20.23 20.19 20.20 19.13 19.55 18.91 19.90 20.11 19.71 16.66 17.66 17.69 17.05 15.33 16.43 17.79 19.07 18.88 17.95 18.50 20.75 18.82 20.39 21.76 21.65 21.57 15.61 15.16

13

y C 7.94 8.47 8.64 9.90 10.01 9.97 8.13 9.09 8.93 7.48 7.87 7.92 8.32 9.08 9.52 11.19 11.07 6.17 6.48 6.21 11.67 10.89 11.21 10.08 7.55 9.00 9.08 5.19 2.10 8.40 9.38 8.96 8.70 9.02 9.13 8.26 9.01 8.13 8.50 9.09 7.13 7.91 7.86 13.13 12.49 9.14 12.14 9.52 7.87 8.53 9.45 9.75 10.31

Sample ID

Stratigraphic level (m)

y18O

LPC 344.5 3445a 3445b LPC 350.2-1 LPC 350.2-2 LPC 358.2 3582a 3582b LPC 358.6 LPC 361.8 3618 LPC 364.5 3645a 3645b LPC 369.2 3692a 3692b LPC 370.0 3700a 3700b LPC 371.8 LPC 379.4 LPC 382.5-1 LPC 382.5-2 3825 LPC 383.9 3839a 3839b LPC 389.1 LPC 398.5 LPC 400.2-1 LPC 400.2-2 LPC 400.2-3 LPC 400.2-4 LPC 400.2-5 4145a 4145b LPC 417.5-1 LPC 417.5-2 4175 LPC 433.8 4338 LPC 435.1 4351 LPC 443.1 LPC 444.1 4441a 4441b LPC 444.6 4446 4458a 4458b

344.5 344.5 344.5 350.2 350.2 358.2 358.2 358.2 358.6 361.8 361.8 364.5 364.5 364.5 369.2 369.2 369.2 370.0 370.0 370.0 371.8 379.4 382.5 382.5 382.5 383.9 383.9 383.9 389.1 398.5 400.2 400.2 400.2 400.2 400.2 414.5 414.5 417.5 417.5 417.5 433.8 433.8 435.1 435.1 443.1 444.1 444.1 444.1 444.6 444.6 445.8 445.8

18.05 17.69 20.97 14.89 15.27 16.34 16.99 18.37 18.70 17.00 16.92 18.75 19.78 17.98 13.93 13.70 14.24 13.45 15.31 13.88 18.68 20.12 19.85 20.01 20.45 15.36 15.09 17.22 20.91 14.45 19.07 18.94 16.55 18.96 16.37 17.64 17.50 18.16 19.19 18.71 20.49 20.83 18.82 19.25 17.17 16.31 18.60 17.53 16.88 17.21 18.55 16.84

y13C 8.76 9.20 8.95 12.05 11.94 11.52 12.24 10.57 8.23 8.24 8.65 7.66 8.17 9.63 6.07 7.24 6.75 6.28 6.35 7.26 7.63 9.45 8.96 9.09 7.89 6.64 7.79 7.87 9.94 10.01 7.69 7.98 9.18 9.79 9.60 9.87 8.38 7.63 7.82 7.51 8.04 9.23 7.73 8.65 8.69 8.49 9.44 9.24 8.91 9.96 8.59 8.93

(continued on next page)

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Table 2 (continued)

Table 2 (continued) 18

Sample ID

Stratigraphic level (m)

y O

LPC 447.9 4479a 4479b LPC 451.7 4517a 4517b LPC 455.1 4551a 4551b LPC 458.1 4581 LPC 468.4 LPC 470.9 4709a 4709b LPC 471.7 4717a 4717b LPC 476.9 LPC 477.1 4771a 4771b LPC 478.4 LPC 480.0 4800a 4800b LPC 480.4 4804a 4804b LPC 481.3 4813a 4813b LPC 484.0 4840a 4840b LPC 486.4 4864a 4864b LPC 487.6 LPC 488.2 4882a 4882b LPC 490.2 4902a 4902b LPC 492.4 LPC 493.2 4932 LPC 495.0 4950a 4950b LPC 495.4 4954a

447.9 447.9 447.9 451.7 451.7 451.7 455.1 455.1 455.1 458.1 458.1 468.4 470.9 470.9 470.9 471.7 471.7 471.7 476.9 477.1 477.1 477.1 478.4 480.0 480.0 480.0 480.4 480.4 480.4 481.3 481.3 481.3 484.0 484.0 484.0 486.4 486.4 486.4 487.6 488.2 488.2 488.2 490.2 490.2 490.2 492.4 493.2 493.2 495.0 495.0 495.0 495.4 495.4

14.97 15.58 17.27 18.12 18.19 18.26 17.81 17.21 16.52 17.65 20.16 21.54 16.31 16.98 14.69 20.29 20.71 20.14 19.06 16.73 17.07 17.24 13.24 14.20 15.59 14.41 15.63 15.54 14.96 14.71 14.08 13.76 14.30 16.97 14.67 13.07 16.78 16.52 17.56 18.37 18.92 15.13 16.42 16.25 16.98 18.38 14.72 15.10 19.27 18.06 19.44 14.89 17.11

13

y C 8.60 8.84 8.81 6.80 7.41 7.10 10.14 10.08 10.00 9.34 8.66 10.47 9.61 10.28 10.57 7.18 7.79 8.37 9.90 9.49 10.63 9.98 9.99 9.45 10.79 11.16 8.62 9.19 9.13 8.96 9.45 9.12 8.65 8.88 8.54 9.24 9.39 8.72 10.07 8.66 10.92 9.21 9.85 8.32 9.56 8.11 7.68 7.65 7.96 7.07 7.41 6.52 6.88

Sample ID

Stratigraphic level (m)

y18O

4954b LPC 495.8 LPC 497.0 LPC 497.4-1 LPC 497.4-2 LPC 497.4-3

495.4 495.8 497.0 497.4 497.4 497.4

16.86 17.22 15.92 14.79 13.42 13.62

y13C 6.53 7.97 9.06 6.42 6.73 6.83

For all measurements, we estimate the 2r uncertainty to be V0.1x for C and V0.2x for O.

the younger part of the section. We also note a marked increase in the scatter of the oxygen values from the lower to the upper part of the section, which is largely directed towards relatively low y18O values. In the upper part of the section, the range of the oxygen values has an upper limit similar to that in the lower part of the section, whereas the lower limit is decreased by 3–4x.

6. Interpretation of stable isotope data The carbon isotope data indicate a significant ecosystem change at ~16.5 Ma. The y13C values of pedogenic carbonate nodules formed prior to ~16.5 Ma (from 14.4 m to 200.8 m above the base of the measured section) are consistent with a pure C3 ecosystem, and imply the presence of balanced woodlands and grasslands. The trend towards lower values throughout this stratigraphic interval (with mean values of 9.8x from 0 to 100 m, and 11.8x from 100 to 201 m) suggests increasing humidity in the eastern foreland during this time interval. In the younger part of the section, in contrast, most nodules have y13C values higher than 10x. Some of this y13C increase could be due to effects of low water availability on the carbon isotope composition of C3 plants [26], implying increased aridity since ~16.5 Ma. Moreover, many of these nodules have y13C values higher than 8x, which indicates a significant proportion of C4 grasses and is also consistent with increased aridity. This appearance or expansion of C4 plants is significantly earlier than the expansion of C4 grasslands observed at 7–8 Ma in many other regions around the globe, e.g., northwest Argentina [44,45],

P.M. Blisniuk et al. / Earth and Planetary Science Letters 230 (2005) 125–142 Table 3 Stable isotope values for individual paleosol horizons (n=102) given as averages (with F indicating one standard deviation) of the values determined for each individual nodule Stratigraphic level (m)

#

y18O

F

14.4 15.4 19.3 21.9 22.5 53.3 60.7 69.2 78.7 79.0 81.6 83.9 86.7 93.0 106.6 111.2 144.1 144.6 145.1 150.3 150.6 152.9 154.5 155.6 159.1 160.2 160.9 162.8 164.7 168.6 174.8 179.3 200.8 215.5 220.9 222.7 223.4 224.7 225.8 226.8 250.6 252.1 254.7 261.2 282.0 282.3 306.8 322.4 322.6 326.8 332.0

3 3 3 3 3 3 2 2 3 3 3 3 3 5 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 3 2 2 4 3 3 3 3 3 3 3 2 3 3 1 3 2 5 3 3 3 2

20.22 19.86 17.11 21.09 19.92 18.14 17.64 19.35 17.06 19.31 19.93 20.15 20.42 18.27 17.89 16.37 18.48 19.83 19.23 19.47 19.02 19.49 18.17 17.73 19.61 20.19 20.62 18.99 18.56 18.20 19.29 16.19 13.56 17.08 18.26 17.43 16.42 16.78 17.09 17.74 17.12 15.42 16.44 18.97 20.55 19.66 19.63 17.34 16.27 18.58 18.22

0.18 0.53 0.38 0.32 0.97 0.44 1.64 1.17 0.46 0.21 0.16 1.18 0.10 0.47 1.19 0.14 0.62 0.68 0.45 1.36 0.53 0.30 0.43 0.56 0.28 0.58 0.20 0.32 0.26 0.27 0.00 2.07 0.15 0.72 1.10 1.10 0.21 0.83 0.95 0.44 3.49 0.52 1.40 n.a. 0.58 0.75 0.46 0.59 0.87 0.69 0.39

y13C 10.31 9.99 10.05 9.50 9.19 10.08 11.90 8.78 10.10 9.07 8.39 9.12 8.60 11.53 12.11 12.37 11.23 9.16 9.19 12.89 13.12 12.94 11.46 11.69 13.12 10.73 10.91 11.24 9.41 13.08 12.65 13.13 13.60 9.87 10.87 8.35 9.96 8.71 7.76 8.97 11.13 6.29 11.26 10.08 8.54 3.65 8.89 8.80 8.58 7.63 12.81

F 0.87 0.43 0.26 0.48 0.98 0.68 0.78 0.25 0.09 0.28 0.55 1.40 0.51 0.67 1.75 0.24 1.28 0.44 0.18 0.48 0.58 0.28 0.80 0.58 0.61 0.61 0.21 0.97 0.14 0.44 0.33 0.81 0.10 0.43 0.61 0.37 0.05 0.52 0.24 0.61 0.09 0.17 0.39 n.a. 0.86 2.19 0.37 0.47 0.49 0.44 0.45

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Table 3 (continued) Stratigraphic level (m)

#

y18O

F

335.0 335.5 343.2 344.5 350.2 358.2 358.6 361.8 364.5 369.2 370.0 371.8 379.4 382.5 383.9 389.1 398.5 400.2 414.5 417.5 433.8 435.1 443.1 444.1 444.6 445.8 447.9 451.7 455.1 458.1 468.4 470.9 471.7 476.9 477.1 478.4 480.0 480.4 481.3 484.0 486.4 487.6 488.2 490.2 492.4 493.2 495.0 495.4 495.8 497.0 497.4

3 3 2 3 2 3 1 2 3 3 3 1 1 3 3 1 1 5 2 3 2 2 1 3 2 2 3 3 3 2 1 3 3 1 3 1 3 3 3 3 3 1 3 3 1 2 3 3 1 1 3

19.99 21.66 15.38 18.90 15.08 17.23 18.70 16.96 18.84 13.96 14.21 18.68 20.12 20.10 15.89 20.91 14.45 17.98 17.57 18.69 20.66 19.03 17.17 17.48 17.04 17.69 15.94 18.19 17.18 18.90 21.54 16.00 20.38 19.06 17.01 13.24 14.73 15.38 14.19 15.31 15.46 17.56 17.47 16.55 18.38 14.91 18.92 16.29 17.22 15.92 13.95

1.03 0.09 0.32 1.80 0.27 1.04 n.a. 0.06 0.90 0.27 0.97 n.a. n.a. 0.31 1.16 n.a. n.a. 1.39 0.10 0.51 0.24 0.30 n.a. 1.14 0.24 1.21 1.19 0.07 0.64 1.78 n.a. 1.18 0.30 n.a. 0.26 n.a. 0.75 0.37 0.48 1.44 2.07 n.a. 2.05 0.38 n.a. 0.27 0.75 1.21 n.a. n.a. 0.74

y13C 10.27 8.62 10.03 8.97 11.99 11.44 8.23 8.44 8.48 6.69 6.63 7.63 9.45 8.65 7.43 9.94 10.01 8.85 9.13 7.65 8.64 8.19 8.69 9.06 9.44 8.76 8.75 7.11 10.07 9.00 10.47 10.15 7.78 9.90 10.03 9.99 10.47 8.98 9.18 8.69 9.12 10.07 9.60 9.24 8.11 7.66 7.48 6.65 7.97 9.06 6.66

F 1.63 0.79 0.39 0.22 0.08 0.84 n.a. 0.29 1.02 0.59 0.54 n.a. n.a. 0.66 0.69 n.a. n.a. 0.96 1.05 0.15 0.84 0.65 n.a. 0.50 0.75 0.24 0.13 0.31 0.07 0.48 n.a. 0.49 0.59 n.a. 0.57 n.a. 0.90 0.31 0.25 0.17 0.35 n.a. 1.18 0.82 n.a. 0.02 0.45 0.20 n.a. n.a. 0.21

Symbol # denotes number of nodules analyzed from each individual paleosol horizon. For completeness, data from paleosols where only a single nodule was analyzed are also listed.

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North America [46], East Africa [47], and South Asia [48,49]. It is, however, consistent with similar evidence for C4 plants since 15 Ma or earlier in Bolivia [50], East Africa [51,52], and East Asia [53]. This supports the view that, although the evolution of C4 ecosystems has largely been controlled by a decrease of the atmospheric CO2 content at 7–8 Ma [46], C4 ecosystems have evolved gradually since at least 15 Ma, in response to both global and local environmental factors (e.g., [51]). The most plausible explanations for the ecologic change implied by our y13C data are either that aridity increased as the result of global climate change, or that the orographic rain shadow of the southern Patagonian Andes was established or significantly enhanced at that time. Reconstruction of global paleoclimate implies relatively stable paleoclimate conditions during deposition of the Santa Cruz Formation, with only minor warming 20–15 Ma (see [23] for a review). It therefore seems most likely that significant surface uplift of this Andean segment had established a pronounced orographic rain shadow by ~16.5 Ma. In this scenario, the trend towards lower carbon isotope values in the older part of the section may reflect that the initial phase of this uplift caused an increased supply of moisture to the eastern foreland prior to rain shadow formation. To estimate the effect of topography on the oxygen isotope composition of soil water and pedogenic carbonate, it is important to consider the influence of temperature and evaporation. Other influences such as changing moisture source, latitude and continentality are unimportant in this location, as discussed above. Also, there is no evidence for a significant change in the oxygen isotope value of the source water (Pacific Ocean) during the deposition of the Santa Cruz Formation. Benthic and planktonic foraminifera y18O values as a proxy for global deepsea and southwestern Pacific deep-sea and surface water y18O values imply that the oxygen isotope composition of the source water has varied by b0.4x between 24 and 14 Ma; even the y18O increase of ocean water attributed to major growth of the Antarctic ice-sheet between 14.5 and ~10 Ma was only ~1.2x [22,23]. Changes in ocean circulation patterns and sea surface temperatures are potentially more important. The Neogene opening of the Drake Passage likely caused southern hemi-

sphere high-latitude sea surface cooling of ~3 8C, based on an ocean circulation model coupled to a one-dimensional atmospheric model (Toggweiler and Bjornsson [54]). Shevenell et al. [55] document a 6– 7 8C decrease of sea surface temperatures in the southwest Pacific between 14.2 and 13.8 Ma, based on oxygen isotope and Mg/Ca compositions of surface-dwelling foraminifera. In the absence of a temperature change on land, decreasing sea surface temperatures would lead to increasing y18O values of precipitation due to the influence of temperature on the starting moisture content of an airmass; this causes a ~0.55x increase in the y18O of precipitation per 1 8C decrease in sea surface temperature (Jouzel et al. [56]). Thus both ice sheet growth and cooling of the sea surface would cause increases in the average y18O value of precipitation, opposite from our observed decrease in the y18O values of pedogenic calcite. Changes of temperatures on land, in contrast, will lead to a ~0.37x decrease in the y18O values of pedogenic carbonate per 1 8C decrease in temperature, based on the temperature dependence of the calcitewater fractionation factor [57] and the correlation between the oxygen isotope composition of precipitation and temperature [58]. Although the net effect of changes in sea surface temperatures and temperatures on land cannot be estimated with much precision, its influence on the isotopic composition of precipitation during deposition of the Santa Cruz Formation thus is likely to be small. Most importantly, while global cooling since ~14.5 Ma [22,23] may have affected samples in the uppermost part of the section, global climate change cannot explain the major decrease in the y18O values of pedogenic carbonate at ~16.5 Ma. Correspondingly, the main trend in the oxygen isotope data is best explained by rain shadow formation. Increased aridity after rain shadow formation also provides a plausible explanation for the relatively large scatter of the y18O values in the top 200–300 m of the section. The effect of evaporation is hard to quantify exactly, but can clearly be significant: along a transect of the Andes at 47.58S, many present-day surface waters sampled in the rain shadow east of the mountains had deuterium excess values implying substantial evaporation, and these samples had y18O values up to N5x higher than those of apparently non-evaporated samples (Fig. 3) [39]. Since surface

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water composition is an important factor controlling soil water composition, evaporation enrichment of 18 O is likely to have affected many of our pedogenic carbonate samples, and caused an increase in the scatter of the data after the rain shadow formed. Correspondingly, our data recording a ~1.5x decrease of y18O values in the mean soil water composition at ~16.5 Ma probably underestimate the shift in the y18O values of precipitation. We therefore suggest that the 3–4x decrease of the lowest y18O values from the older to the younger part of the section is a better measure for the effect of surface uplift in the southern Patagonian Andes on the y18O values of precipitation in the eastern foreland. Assuming a lapse rate of 0.28x/100 m [34], this would imply that 1.3F0.2 km of surface uplift occurred in the southern Patagonian Andes at ~16.5 Ma.

7. Discussion The stable isotope data presented above imply that surface uplift of the southern Patagonian Andes led to a considerable aridification in the leeward eastern foreland at ~16.5 Ma. Since there is no evidence for an increase in the convergence rate between the Nazca and South American plates at that time [7], the most likely reason for this surface uplift would have been an increase of compression and tectonic shortening due to (1) the subduction of progressively younger and more buoyant oceanic lithosphere as the Chile ridge approached the trench, and (2) an increase in the strength of coupling between the Nazca and South America plates because of the absence of a significant sediment fill in the trench. Furthermore, the buoyancy-driven surface uplift caused by subduction of a spreading center relative to older oceanic lithosphere can be significant [59]. Thus an inferred surface uplift of z1 km is within the expected range of tectonic responses to spreading ridge subduction. We note, however, that the resulting aridification in the eastern foreland did not lead to a decrease in deposition rates, at least not one discernible at the resolution of the available age data. It thus seems that any trend towards decreasing erosion rates resulting from the more arid climate was roughly balanced by a trend towards higher erosion rates resulting from increased

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relief, consistent with the observed upward increase in the content of coarse clastics in Santa Cruz Formation deposits. A much more important change of mass flux rates in the eastern foreland occurred at ca. 14–15 Ma, when deposition of the Santa Cruz Formation terminated. Since then, sedimentation in that region has been almost exclusively limited to short-lived episodes of conglomerate deposition during and immediately after glacial periods [14,15], implying a drastic increase in aridity at that time. As compression and surface uplift in the southern Patagonian Andes presumably increased until the ridge–trench collision started (b15 Ma), it seems possible that at 15–14 Ma a threshold elevation was reached at which the orographic rain shadow effect became much stronger. However, the apparent temporal coincidence of this change with significant global cooling after 14.5 Ma suggests that the transition to a cooler climate at least contributed to the increased aridity east of the mountain belt. Although the relative importance of these processes is uncertain, both presumably resulted in an increased sediment flux from the western side of the mountains into the trench. Increasing surface elevation would have blocked more moisture from reaching the leeward eastern side of the mountains, while leading to an increase of precipitation and erosion rates on their windward western side. Global cooling may have caused an overall drier climate and increased aridity on both sides of the mountains, but also increased storm frequency [60], and thus is also likely to have increased net erosion on the humid windward side of the mountains. Since a thicker trench fill promotes weaker coupling along the plate interface (e.g., [61]), this increased sediment flux into the trench may have contributed to the apparent cessation of subduction erosion and eastward migration of uplift and denudation in the southern Patagonian Andes that had occurred between 30–23 and ~12 Ma [8]. Thus changes in earth surface processes resulting from mountain uplift may feed back on tectonically controlled surface uplift in subduction orogens. In settings like that of the Central Andes, where high surface elevation causes increased aridity on the oceanward side of the mountains and sediment starvation in the trench, such a feedback would be positive [61]. A negative feedback, in contrast, would

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occur in settings like that of the southern Patagonian Andes, where surface uplift leads to increased precipitation and erosion rates on the oceanward side of the mountains.

Acknowledgements We thank George Hilley, Travis Horton, Victor Ramos, Manfred Strecker, and Jacob Waldbauer for discussions, and Susana and Pedro Fortuny for help with local logistics. We are also grateful to Fre´de´ric Fluteau and Bruce MacFadden for constructive reviews that improved the original manuscript. This project has been funded by the Deutsche Forschungsgemeinschaft (SFB 267).

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