Introduction

With average elevations of ~4 km, low internal relief, closed and partially coalesced sedimentary basins, the Andean Plateau (AP, Altiplano-Puna) constitutes the world’s second largest orogenic plateau1 (Fig. 1). Topographic growth of the AP and its high eastern margin causes the intercept of moisture-bearing easterly winds, resulting in an efficient orographic barrier with steep E-W rainfall and vegetation gradients2,3,4 (Fig. 1). Today, the undeformed foreland and the eastern flanks of the AP between 13° and 30°S receive summer precipitation of up to 3 m/yr mainly via the southward transport of Amazonian moisture by the South American Low-Level Jet (LLJ) of the South American Monsoon (SAM)5. In contrast, leeward basins to the west of this orographic barrier are arid and receive rainfall amounts of less than 0.2 m/yr2. These conditions are mirrored by dense C3 vegetation on the eastern flanks of the AP, whereas the orogen interior hosts a mixed C3/C4 vegetation, reflecting aridity6.

Figure 1
figure 1

Rainfall and topographic patterns of the Central Andes and structural context of the study area.

Figure 1 was created using Adobe Illustrator CS6 and the base map was created with ArcGIS 10.1. (a) TRMM 3B42 annual rainfall map and morphotectonic provinces (EC: Eastern Cordillera, SBS: Santa Barbara System)2. Arrows depict moisture transport direction (SALLJ - South American Low-Level Jet). (b) Topography and structural context of the south-central Andes and location of the study area. (c) Schematic geological cross section along profile line shown in B33,34. Cloud symbol indicates region of present-day enhanced orographic rainout, i.e. the position of the effective orographic barrier.

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Stable isotope data suggest that the establishment of present-day climatic conditions, topography, and vegetation patterns in the Bolivian Andes (20–22°S) was coeval with the onset of the LLJ at ca. 9 Ma7,8. This has been linked to surface uplift of the central AP and redirection of air-mass trajectories in South America4,8,9,10. However, discontinuous uplift11,12, long-term surface evaporation13, global climate change14, and changes in the water stable isotope compositions occur simultaneously with surface uplift in the Andes15. Recent paleotopographic reconstructions challenge these interpretations and postulate the existence of a high-elevation AP in NW Argentina already during the Oligocene16,17. To evaluate these inconsistencies additional information on the topographic growth and climate south of 22°S is needed, especially in NW Argentina. Previous studies suggest that global climate change did not have a profound affect on South American climate during the time period in question, but rather that the growth of Andean topography reinforced the preexisting climatic regime and promoted further aridification within the AP and along the Pacific cost18,19. This particularly applies to surface uplift during the establishment of the LLJ and associated reconfiguration in the provision of nutrients, biodiversity, fluvial networks, and other eco-hydrological changes that affected sediment flux toward the foreland and ocean basins20.

One commonly encountered impediment in stable isotope paleoaltimetry is the potential transience of stable isotope lapse rates through time, e.g. changes in moisture sources and trajectories due to topographic uplift or climate change10,21. In this study such transients are considered to be of minor relevance for the intermontane Angastaco Basin (AB) of NW Argentina, as it is located within the Eastern Cordillera (Fig. 1) immediately east of the AP, with former sediment transport directions pointing to unrestricted eastward transport22 and no indication for modifications of airflow trajectories due to topographic uplift or changes in the moisture source2,8,10. Prior to deformation and surface uplift the basin was an integral part of the low-elevation (<500 m) foreland22. In our study we use a multi-isotope approach combining inorganic and organic proxies to decode moisture supply changes to the AB that reflect evolving eco-hydrological conditions along the AP flanks. Today the basin is located at ca. 2 km elevation and hosts up to 6 km of sediments, deposited during tectonic deformation along the eastern flank of the AP beginning at ca. 9 Ma. Geochronology of abundant intercalated volcanic ashes in the AB furnishes a unique, well-dated record of interlayered organic-rich beds and carbonate-bearing paleosols (Fig. 2; Fig. S1). The former contain lipid biomarkers, specifically n-C27 to n-C33 alkanes derived from leaf waxes of higher terrestrial plants that originated in the basin catchment along the eastern AP margin (Fig. 1; Figs S2–7 and 10; Table S1)23. We focus on paleoenvironmental proxy data from the stable isotopic composition of pedogenic carbonates (δ18Osc), volcanic glass (δDvg), and n-C29 and n-C31 alkanes (δDwax) (see SI)24. The δDwax signal typically forms during the early season as leaves flush over variable timescales. This corresponds to weeks in deciduous plants25,26 to months in evergreen plants27 and coincides with the monsoonal rainy season in the study area. In the southern central Andes δ18Osc have also been reported to form during the rainy season, or at its end28,29. Therefore, both proxies record similar seasonal timeframes (i.e. the austral summer rainy season). We use δDwax as a proxy for the leaf water hydrogen isotopic composition (a function of plant transpiration), soil-carbonate δ18Osc as a proxy for soil water δ18O values (a function of soil evaporation), and volcanic glass δDvg as a proxy for precipitation δD values, and their respective isotopic differences as a measure of plant transpiration (Δ δDleaf − δ18Osoil) and soil evaporation (Δ δ18Osoil − δDprecipitation) (Figs S8 and 9; see methods)25,30,31. Our stable isotope data span the time between 8.5 to 3.5 Ma, corresponding to the depositional time frame of the exceptionally well exposed Miocene Palo Pintado and Pliocene San Felipe formations of the AB.

Figure 2
figure 2

Stratigraphy, water stable isotopes, and tectono-sedimentary evolution of the Angastaco Basin.

(a) Measured stratigraphic section showing fossil-bearing horizons (Caiman latirostris42 and fossil plant data from this study); Geochronological data from aCarrapa et al.35, bPingel et al.41, and cBywater-Reyes et al.43; and position of leaf wax and soil carbonate samples. (b) Stable isotope compilation of leaf wax δDwax (this study), soil carbonate δ18Osc (this study and data from *Bywater-Reyes et al.43 and volcanic glass shards δDvg Pingel et al.41. All isotope data has a 2-sigma age error of 0.1–0.3 Ma and δDwax and δ18Osc represent pooled precisions of 5% based on external isotope standards and analytical errors, respectively. δD and δ18O data are presented with respect to Vienna Standard Mean Ocean Water (VSMOW). (c) Deformation history based on own observations and low-temperature thermochronology33,34,35.

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Our isotope records distinguish between two stages in eco-hydrological and topographic conditions along the eastern AP: (1) a humid phase, which corresponds to a contiguous open foreland setting; and (2) an arid phase, representing an intermontane basin setting following the establishment of an orographic barrier (Sierra León Muerto) east of the present-day AB (Figs 1 and 2). Below we discuss the sequence of environmental changes and the corresponding regional tectonic forcing.

Paleoenvironmental and sedimentologic data prior to deposition of the Palo Pintado formation indicate arid conditions within the AB32,33. With the onset of deposition of the Palo Pintado formation between ca. 8.5 to 7.6 Ma δDwax and δ18Osc values decreased from ca. −110 to −150%, and 23 to 18%, respectively, indicating a trend towards wetter conditions in this inherently dry region (Fig. 2). For the same time interval the δ13Cwax and δ13Csc records, pollen and plant macrofossil data suggest a stable C3 forest ecosystem in the AB (Fig. 3)32 likely similar to present-day environmental conditions in the foreland ca. 150 km to the East of the AB22. The isotopic composition of intercalated volcanic ash deposits (glass shards) within the AB section (δDvg) supports this notion as it indicates that meteoric water in the contiguous AB had a similar stable isotope composition at that time as modern meteoric waters from the undeformed foreland (Fig. 2)21. The wetter conditions indicated by the decrease in δDwax and δ18Osc values until 7.6 Ma may have resulted from enhanced orographic rainfall as a result of regional topographic growth along the Eastern Cordillera as inferred from low-temperature thermochronology (Sierra Luracatao, Cerro Durazno and Sierra de Quilmes)8,32,34,35. The evapotranspiration data (Fig. 3) suggest that the relative amounts of evaporation and transpiration were reduced in the AB in concert with enhanced orographic precipitation, whereas plant transpiration appears to have been unaffected (Fig. 3).

Figure 3
figure 3

Carbon stable isotopes, vegetation and reconstructed evapotranspiration.

(a) Stable isotope records of leaf wax δ13Cwax (this study) and soil carbonate δ13Csc (this study and data from *Bywater-Reyes et al.43. Carbonate isotope records are reported according to the Pee Dee Belemnite (PDB). All isotope data has a 2-sigma age error of 0.1–0.3 Ma and δDwax and δ18Osc represent pooled precisions of 0.5% based on external isotope standards and analytical errors, respectively. (b) Reconstruction of evaporation, transpiration and evapotranspiration based on proxy source-water reconstruction. Reconstructed source-waters for lipid-biomarkers, pedogenic carbonates and volcanic glass shards have different source-waters for the individual proxies: volcanic glass - precipitation; pedogenic carbonates – soil water; lipid biomarkers – leaf water. For comparison δ18Osc source waters are converted to δD values.

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Between 7.6 and 6.5 Ma the sedimentation rate in the AB doubled from 0.5 to 1.0 km/Myr, and sedimentary environments changed from alluvial fans to meandering rivers with overbank deposits and paleosol formation (Fig. 2). Over the same time period both δDwax and δ18Osc records consistently display a large variability in isotope values (Fig. 2) of more than 50%, and 8%, respectively. These large variations in stable isotope values imply considerable changes in rainfall variability on timescales larger than each proxy record’s integration time. δDwax and δ18Osc integrate their isotopic signal over similar timespans, whereas δDwax reflect soil formation timescales on the order of ca. 1 to 5 ka as a function of soil thickness36 and δ18Osc represents integration time on the order of 0.1 to 10 ka37,38. Collectively, the isotope data point to the existence of a paleo LLJ south of 22°S as early as 7.6 Ma that facilitated moisture transport into the inherently dry region. Most likely the intensity was modulated by solar insolation linked to orbital variations reflected by the large variability in δDwax and δ18Osc values. The inference of a regionally extensive humid climate at that time is supported by the character of sedimentary facies, fossil content and isotope geochemistry of paleosols from other basins along the eastern flanks of the AP. These areas include basins close to the Argentina-Bolivia border at 22°53′S and 64°36′W39 and regions in southern Bolivia at 21°S8. Interestingly, the inferred onset of the LLJ in NW Argentina is about one million years later than its postulated impact in southern Bolivia8, suggesting a latitudinal lag time along the eastern Andes. The AB carbon isotope records of δ13Cwax and δ13Csc during this period document little variation of 3.2% and 3.9%, respectively, reflecting a sustained C3 forest ecosystem (Fig. 3)30,40. This is corroborated by abundant tree trunks, leaf macrofossils, and pollen assemblages from the Palo Pintado Formation that resemble the present-day Paranaense flora of southwest Brazil32 (Fig. 2).

After approximately 6.5 Ma the variability in δDwax and δ18Osc values decreased to less than 38% and 1%, respectively. Prior to the decrease in variability δDwax and δ18Osc values increase over a short period of only 0.2 Myrs from −160 to −102% and from 18 to 31%, respectively. We interpret this signal to reflect the rapid establishment of arid conditions in the AB in the absence of regional circulation and moisture source changes at that time8,10. Instead, tectonic shortening and uplift at the eastern basin margin (the present-day Sierra de León Muerto) at ca. 7 Ma35 caused the formation of an effective orographic barrier and the transition from a contiguous foreland with E directed drainage into an intermontane basin setting with axial, N-S oriented drainage conditions (Figs 1 and 2)32. Today, the threshold elevation for effective orographic barriers blocking approximately 90% of incoming moisture in the south-central Andes is ca. 1.5 to 2 km2,22. Therefore, the increase in δDwax and δ18Osc values at 6.5 Ma can be interpreted to reflect rapid aridification in response to the Sierra de León Muerto reaching a threshold elevation of 2 km. This is compatible with a previous study that suggest that the coeval decrease in δDvg values (Fig. 2) resulted from stronger Rayleigh rainout due to uplift41. Surface uplift along the eastern basin margin coincides with a rapid increase in plant transpiration in response to reduced moisture supply into the AB (Fig. 3). In contrast to the basin margin, the basin itself did not gain significant elevation during this initial stage of deformation and uplift, as indicated by a ca. 6 Ma caiman fossil (Caiman latirostris), whose extant relatives live below 800 m in the modern Andean foreland (Fig. 2)32,42. Subsequently, however, surface uplift of the intermontane basin involved strata within the AB and resulted in the deposition of coarser sediments of the San Felipe Formation (Figs 1 and 2)35. As a result of continued surface uplift (5.5 and 3.5 Ma) the basin further aridified as documented by overall increasing δDwax values (ca. 130% to −98%), thick calcrete-bearing paleosols, and deposition of halite and gypsum in restricted fluvio-lacustrine environments (Fig. 2)43.

During this episode the ecosystem fundamentally changed as revealed by δ13Cwax and δ13Csc values increasing from −32 to −28% and −11 to −8%, respectively (Fig. 3). These new environmental conditions could either correspond to a mixed C3/C4 high-altitude grassland ecosystem or represent greater water use efficiency of plants in an environment with low water availability, or both6,44. An environmental shift towards a C4 successor vegetation is also supported by rodent tooth enamel isotope data from the San Felipe Formation, which points to increased availability of C4 plants43. Reconstructed post-6 Ma clumped isotope temperatures indicate lower temperatures between 18° to 25 °C, similar to present-day summer temperatures21 and much cooler than previously inferred foreland temperatures of up to 38 °C at 7 Ma17. Uplift of the orographic barrier and transition to cooler, more arid conditions are reflected by sustained high levels of plant transpiration, evaporation, and evapotranspiration in the AB (Fig. 3). The large increase in leaf water enrichment of >60% inferred from terrestrial organic material during orographic barrier uplift is similar to the present-day difference between the Amazonian rainforest and the semi-arid basins along the eastern Andes (Figs 3 and S9)25.

In summary, our new multi-proxy, multi-isotope record from the AB unravels the intimate coupling of tectonic and climatic events in the southern central Andes and major Mio-Pliocene eco-hydrological changes that were part of a continent-wide redistribution of moisture, as predicted by General Circulation Models e.g.4,10. Our observations validate these models of enhanced convective rainfall and moisture transport via the LLJ in concert with progressive surface uplift of the AP10. The resulting seasonally hot and humid climate of the southern central Andes was characterized by variable amounts of evapotranspiration related to the strength of regional moisture transport via the LLJ (Fig. 3) and the establishment of a C3 forest ecosystem independent of global climate change. Our findings agree with climate models that point to a strong control of Andean topography on regional climate and a rather weak link to late Miocene - Pliocene global climate change e.g.4,14,45,46. Although decreasing atmospheric carbon dioxide levels, air and sea surface temperatures throughout the Late Miocene47,48,49 also impact deuterium and oxygen stable isotope values in South America14 the close proximity of our record to the uplifting eastern flanks of the AP dominates the climatic and eco-hydrologic signal in the AB deposits. This is remarkably well illustrated by the correlation of tectonic events and related environmental changes within the AB (Figs 2 and 3). However, humid eco-hydrologic conditions related to the orographic barrier of the AP flanks rapidly ended after 6.5 Ma as local orographic barrier uplift farther east forced cooler and drier conditions in the AB. As a result, we observe that stable isotope proxy records of plant wax and pedogenic carbonates primarily respond to thresholds of moisture supply in the lee of orographic barriers rather than surface uplift or regional to global climate change. The large eco-hydrologic gradients observed over time support the notion that mountain building forces landscape and climate change, and thus ecological gradients and physical habitats that set the stage for speciation and biodiversity20.

Methods

The multi-proxy-isotope approach presented here represents one of the first steps to reconstruct the different amounts of evaporation, transpiration and evapotranspiration from a single record on a million of year timescale. Assuming that individual δDvg and δ18Osc source waters are representative of precipitation and soil water, respectively, the isotopic difference between both values must then represent evaporation30,31 (Fig. S8). In addition, the difference between δ18Osc and δDwax source water values reflects leaf-water isotope enrichment above soil water, and thus transpiration25,30,31 (Fig. 3; Figs S8 and S9). Ultimately, the difference between reconstructed source-water values from δDvg and δDwax therefore represents evapotranspiration (Fig. 3; Figs S8 and S9). The relationships between the isotopically distinct water sources recorded by the proxies (leaf wax – leaf-water; soil carbonates – soil water; volcanic glass – meteoric water) and evaporation, transpiration and evapotranspiration are schematically sketched in Fig. S8. In combination the three isotope proxies can be used to reconstruct the total relative evapotranspiration trends in a region (Figs S8 and S9). To compare the data time series on a common timescale with the same temporal resolution we calculated equidistant time series based on a 100 ka time window. For calculating the equidistant time series we used MATLAB© (version R2010b) and used the method “interpl” with the specification “linear”. From the equidistant time series for leaf wax, soil carbonates and volcanic glass source waters we calculated the differences between each proxy representing evaporation, transpiration and evapotranspiration. The underlying individual source water reconstructions are discussed in further detail below.

δDwax values are determined by the isotopic composition of water that is available for biosynthesis of leaf wax n-alkanes representing leaf-water and the net biosynthetic fractionation called εbio that accounts for various biochemical hydrogen isotope fractionations during biosynthesis of n-alkanes (Fig. S8). Assuming εbio is constant over the timescales of proxy integration in sediments, it is possible to reconstruct the isotopic composition of leaf-water from leaf-wax n-alkanes. The isotopic difference between the plants source water, i.e. soil-water, and leaf-water should thus arise due to (D)-enrichment in response to plant transpiration (Fig. S8)25. While εbio of leaf waxes from sedimentary deposits, integrating a vast number of plant individuals and species, can be regarded constant, recent studies reported a significant spread in biosynthetic fractionation among different plant species between ca. −130 to −190%25. For leaf-water reconstruction we use a combined εbio value of −160 ± 30% representing the range of reported εbio value for higher terrestrial plants from several available studies27,50,51. Reconstructed leaf-water is presented in Table S1 and Fig. S8 and source water differences in Fig. S9.

Pedogenic carbonates usually form at soil depths of ~20 to 30 cm as the result of evaporating surface water. As calcite forms it records a signal of δ18O and δ13C stable isotopes. The δ18O signal reflects the isotopic composition of the surface water and the δ13C signal has been interpreted to reflect bioproductivity, CO2 plant respiration or C4/C3 vegetation type. The calcite formation and the δ18O isotope values are strongly controlled by temperature-dependent isotope fractionation between water and carbonate52. Therefore, for the reconstruction of paleosurface water the carbonate formation temperature needs to be known. Here, we use reconstructed clumped-isotope temperatures from Carrapa et al.17 for deciphering the temperature-dependent isotope fractionation using the fractionation factor of Kim and O’Neil17,52. We linearly interpolate in between clumped-isotope temperatures points, where our sampling of soil-carbonates was denser. The error reported for reconstructed source water of δ18Osc incorporates the analytical error for each analysis and was 0.1 to 0.3% (Table S2). In addition to analytical uncertainties other processes such as the location of pedogenic carbonate formation (e.g. open versus shaded for example) can have an effect on clumped-isotope temperatures of up to 10 °C that are used during reconstruction of source water53. Calculated values of the δ18Ocarb source waters are reported in Table S2 and Figs 3 and S9. For comparison with δDwax and δDvg source waters the pedogenic carbonate δ18Ocarb source water was converted to deuterium values using the present-day local meteoric water line of δD = 8.44*δ18O + 15.91 for 26°S from Rohrmann et al.21 (Table S2 and Figs 3 and S9)21.

After deposition rhyolitic glass incorporates large amounts of meteoric water (3 to 8 wt%) over a time frame of 5 to 10 ky31. The final δDvg signal represents an integrated meteoric water signal during hydration over geological time scales. To compare δDvg with other isotope proxy materials, e.g. lipid-biomarker and soil-carbonates, we converted the δDvg signal to meteoric water δD values (δDmw), using the equation δDmw = 1.0343 [1000 + δDvg ] − 100031. Converted meteoric-water and δDvg values are presented in Table S3 and Figs 3 and S9.

Additional Information

How to cite this article: Rohrmann, A. et al. Miocene orographic uplift forces rapid hydrological change in the southern central Andes. Sci. Rep. 6, 35678; doi: 10.1038/srep35678 (2016).