Cenozoic organic carbon isotope and pollen records from the Xining Basin, NE Tibetan Plateau, and their palaeoenvironmental significance

Palaeogeography, Palaeoclimatology, Palaeoecology 386 (2013) 436–444

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Cenozoic organic carbon isotope and pollen records from the Xining Basin, NE Tibetan Plateau, and their palaeoenvironmental significance Yunping Chi a,c, Xiaomin Fang a,b,c,⁎, Chunhui Song a,c, Yunfa Miao d, Xiaohua Teng c, Wenxia Han e, Fuli Wu b, Jiwei Yang a a

School of Earth Sciences & Key Laboratory of Western China's Mineral Resources of Gansu Province, Lanzhou University, Lanzhou 730000, China Key Laboratory of Continental Collision and Plateau Uplift, Institute of Tibetan Plateau Research, Chinese Academy of Sciences, Beijing 100101, China MOE Key Laboratory of Western China's Environmental Systems & Research School of Arid Environment & Climate Change, Lanzhou University, Lanzhou 730000, China d Key Laboratory of Desert and Desertification, Cold and Arid Regions Environmental and Engineering Institute, Chinese Academy of Sciences, Lanzhou 730000, China e Key Laboratory of Salt Lake Resources and Chemistry, Qinghai Institute of Salt Lakes, Chinese Academy of Sciences, Xining 810008, China b c

a r t i c l e

i n f o

Article history: Received 3 February 2013 Received in revised form 13 June 2013 Accepted 14 June 2013 Available online 21 June 2013 Keywords: Total organic carbon isotope Pollen Xining Basin Cenozoic Global temperature Monsoon

a b s t r a c t Marine oxygen isotope records show that the Cenozoic global climate has undergone remarkable fundamental changes. However, how terrestrial ecosystems respond to these changes remains unclear. The Xining Basin on the northeastern Tibetan Plateau holds continuous fine fluvial-lacustrine sediments from the early Eocene to the early Miocene (between ~52 Ma and 17 Ma). This paper first presents total organic carbon isotope (δ13CTOC) and pollen records from the Xiejia section in the Xining Basin, as indicators of the long-term terrestrial paleoecosystem and paleoclimate changes in the Asian interior of NW China. The δ13CTOC record shows a long term persistently decreasing trend (from ~−20‰ to −23.5‰) superimposed with obvious cycles of five distinctive phases of heavier and lighter δ13CTOC values, along with short pulses of especially low values occurring at ~32.5 Ma, 23 Ma, 21.5 Ma, and 18 Ma. These variations generally correlate with changes of gymnosperm content or the ratio of gymnosperm to angiosperm (G/A); i.e., heavier δ13CTOC values correspond with higher gymnosperm content or higher G/A, as well as with global temperatures (for both long term trends and shorter term fluctuations). We propose that global temperatures might provide the major factor controlling the evolution of the gymnosperm content (and the G/A ratio) that determines the δ13CTOC in the sediments. The East Asian monsoon may not have reached this region during the deposition of the sediments, or alternatively, its signals could not clearly be identified in these low resolution paleoecologic records. © 2013 Published by Elsevier B.V.

1. Introduction The marine carbon and oxygen isotope records have demonstrated that the global climate undergone a series of tremendous changes in the Cenozoic, characterized by three periods of remarkable stepwise cooling and the initiation and development of polar ice-sheets. Such progressive changes (from a “greenhouse” to an “icehouse” earth) have also led to great transitions of the global carbon reservoir and related ecosystems (e.g., Miller et al., 1987; Zachos et al., 2001, 2008). However, how terrestrial ecosystems respond to these changes remains unclear. Organic carbon isotopes (δ13CTOC) have been regarded as a basic and sensitive indicator of ecosystem evolution. Previous studies have examined the carbon isotopic compositions within various types of plants and their relationships with ecological environments, especially, with temperature and precipitation (e.g., Bender, 1971; ⁎ Corresponding author at: Institute of Tibetan Plateau Research, Chinese Academy of Sciences (CAS), Beijing, 100101, China. Tel.: +86 10 8409 7090; fax: +86 10 8409 7079. E-mail address: [email protected] (X. Fang). 0031-0182/$ – see front matter © 2013 Published by Elsevier B.V. http://dx.doi.org/10.1016/j.palaeo.2013.06.013

Smith and Epstein, 1971; Deines, 1980; O'Leary, 1981; Farquhar et al., 1982; O'Leary, 1988; Farquhar et al., 1989). The knowledge regarding these relationships has been widely applied to lacustrine sediments (e.g., Street-Perrott et al., 1997; Brincat et al., 2000; Huang et al., 2001), eolian loess–paleosol sequences (e.g., Gu et al., 2003; Zhang et al., 2003; Liu et al., 2005; Chen et al., 2006), and marine sediments (e.g., Yamada and Ishiwatari, 1999; Jia et al., 2003), for the reconstruction of paleovegetation, paleoenvironments, and paleoclimates. However, studies of terrestrial ecosystems have mainly focused on the Quaternary, especially the last glacial cycle. These studies have indicated patterns of how ecosystems have evolved with the Quaternary climates, e.g., the relative abundance of C4 plants (which form a 4-carbon molecule instead of two 3-carbon molecules) decreased during glaciation and increased during interglacial intervals, although opinions vary regarding the mechanisms of changes of the δ13C in sediments (e.g., Lin et al., 1991; Lin and Liu, 1992; He et al., 2002; Liu et al., 2002; Gu et al., 2003; Zhang et al., 2003; Rao et al., 2005, 2012). The Cenozoic long-term δ13C and paleoenvironmental evolution on continents are poorly understood, due to the scarcity of precisely time controlled

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paleoclimatic records. Until now, only two rough δ13CTOC records from the SE Australian Gippsland Basin and European Rhine Valley Basin have been reported to reveal the evolution of the ecological environment since the Eocene; and these show vegetation types and changes of atmospheric CO2 concentration closely related to δ13C variations of the sediments (Bechtel et al., 2008; Holdgate et al., 2009). In addition, studies have shown that changes of the ecological environment in Asia have been controlled not only by the global cooling, but have also been affected by other significant events related to their geographical locations, such as the uplift of the Tibetan Plateau, the development of the East Asian monsoon (e.g., Ruddiman and Kutzbach, 1989; Manabe and Broccoli, 1990; Raymo and Ruddiman, 1992; Li and Fang, 1999; An et al., 2001; Lu et al., 2010; Miao et al., 2011, 2012), and the retreat of the Paratethys Sea (e.g., Ramstein et al., 1997; Zhang et al., 2007). Therefore, determining how the continental ecology evolves in response to these factors is an important scientific question. The Xining Basin, located on the northeastern Tibetan Plateau, and filled with thick Cenozoic continuous fine grain sediments, provides an ideal place to investigate these scientific questions. Dupont-Nivet et al. (2007, 2008) and Hoorn et al. (2012) discuss the Tibetan Plateau uplift and the climate changes during the middle Cenozoic based on the first appearance of coniferous trees (beginning ~38 Ma) using very low resolution pollen samples (n b 20), but only for the stratigraphic interval ~33–41.5 Ma. Long et al. (2011) show early Eocene to early Oligocene (52–28 Ma) n-alkane distributions, compared with the pollen for the

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same intervals, and point out that vegetation changes resulted directly from the global cooling and Tibetan Plateau uplift. However, the paleoenvironmental research for the later periods (e.g., the Oligocene– Miocene) has not been carried out for the Xining Basin. Here we present long-term (38 Ma to 17 Ma) δ13CTOC and pollen records from a continuous fluvial-lacustrine sediment sequence of the well-known Xiejia section in the Xining Basin to investigate the trends and variations of the paleoecologic environment and their possible causes. 2. Geologic setting The Xining Basin lies on the northeastern Tibetan Plateau, bordered by the Daban Shan (Mts.) to the north, the Laji Shan to the south, the Riyue Shan to the west, and a small rise to the east (Fig. 1). The Basin lies at an average elevation of about 2000–2100 m, while the surrounding mountains reach about 3000–4000 m. The semi-arid climate presently controls the Basin's weather. Recent mean annual temperatures in the Basin range between and 5–6 °C, with annual precipitation between 350 and 500 mm. Precipitation occurs mainly in the summer, derived from the East Asian summer monsoon. The natural vegetation shows dramatic vertical variations because of the large elevation differences from the lowest basins and valleys to the high surrounding mountains. Stipa bungeana, Stipa glareosa, and Artemisia gmelinii grow mainly in areas below 2600 m. Picea crassifolia and Picea wilsonii are distributed mainly on the shady slopes between 2600 m and 3000 m.

Fig. 1. Tectonic geomorphology and stratigraphic distribution of the Xining Basin.

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Shrubs, including Potentilla fruticosa, Salix oritrepha, and Rhododendron, grow above 3000 m. Alpine meadows (with species such as Kobresia bipartite) lie widely distributed in areas above 3200 m (Zhou et al., 1987). The Cenozoic strata in the basin lie widely distributed and completely exposed along deeply incised river valleys (Fig. 1). On the basis of lithology and fossil mammals, these strata can be divided into two groups and seven formations above the Paleogene Xining Group: (1) the Paleocene to Eocene Qijiachuan Formation (Fm.) (55–215 m thick in measured control sections), (2) the Eocene Honggou Fm. (81–367 m thick), (3) the Oligocene Mahalagou Fm. (245–522 m thick); and the Neogene Guide Group: (4) the early Miocene Xiejia Fm. (78–278 m thick), (5) the middle Miocene Chetougou Fm. (80–245 m thick), (6) the middle to late Miocene Xianshuihe Fm. (22–104 m thick), and (7) the Pliocene Linxia Fm. (10–30 m thick) (Dai et al., 2006; Fang et al., 2007). The lithology of the strata changes from gypsum and gypsiferous mudstones from playa environments since the Eocene, to red mudstone and silty mudstone from distal alluvial fan deposits during the Oligocene to Miocene. The strata generally thicken southward, with thick bedded sulfate sediments (gypsum and mirabilite) deposited in the center of the basin. 3. Sampling and measurements This study collected samples from the Xiejia section (36° 31′ N, 101° 52′ E) in the Xining Basin (Fig. 1). The exposed Xiejia section is 819 m thick and consists of, from the bottom upwards, the Qijiachuan Fm. (0–69 m, from the base), the Honggou Fm. (69–224 m), the Mahalagou Fm. (224–504 m), the Xiejia Fm.

(504–667 m), the Chetougou Fm. (667–778 m), and the Xianshuihe Fm. (778–819 m). The Paleogene Qijiachuan, Honggou, and lower Mahalagou Fms. consist mainly of red mudstones and siltstones, interbedded with many marlstone, gypsum, and gypsiferous mudstone layers. The upper member of the Mahalagou Fm. and the Xiejia Fm. are brownish red mudstones and siltstones, occasionally interbedded with thin gypsum layers. The Chetougou and Xianshuihe Fms. are yellowish brown siltstones and mudstones intercalated with some lenses of fine sandstone. A large number of fossil mammals, named the Xiejian Fauna (Li and Qiu, 1980; Qiu et al., 1981), were found at the thickness of 632 m, and high-resolution paleomagnetic dating has determined the section age at approx. 52.5–17 Ma, with the following approximate formation ages: the Qijiachuan — 52.5–50 Ma; the Honggou — 50–41.5 Ma; the Mahalagou — 41.5–30 Ma; the Xiejia — 30–23 Ma; the Chetougou — 23–18 Ma; and the Xianshuihe — 18–b17 Ma (Dai et al., 2006; Fang et al., 2007) (Fig. 2a and f). Recent paleomagnetic dating of the Tashan section (about 3 km away to the northwest of the Xiejia section) supports the chronology of the Xiejia section (Xiao et al., 2012). In order to avoid the complexity of organic matter sources from the Qijiachuan Fm. to the lower Mahalagou Fm. (caused by frequent large sedimentary facies changes from distal alluvial, fluvial, and salt lakes), this study carried out detailed studies of the carbon isotopes and pollen compositions of the strata from the upper Mahalagou Fm. to the Xianshuihe Fm., which are mainly fine-grained floodplain sediments. This study collected more than 250 samples, at intervals of about 2 m along the upper 518 m of the Xiejia section (from 300 m to 818 m) (Fig. 2b). For δ13CTOC analysis, the samples were dried at

Fig. 2. The Xiejia section: (a) lithologic stratigraphy and depth indicator; (b) the total organic carbon isotope values; (c) percentages of gymnosperms; (d) percentages of angiosperms (note the reversed scale); (e) the ratio of gymnosperms to angiosperms; and (f) observed paleomagnetic zones and their correlations to the geomagnetic polarity timescale (GPTS) of Cande and Kent (1995) for tracking the chronology. Modified from Dai et al. (2006) and Fang et al. (2007).

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40 °C, ground in an agate mortar, then sieved through a 100 mesh screen and homogenized. The samples were then treated with 10% HCl for 24 h to remove carbonates. Samples were then washed to a pH N 6 by distilled water and dried at 40 °C, and finally weighed and sealed in tin vessels for δ13CTOC analysis. The measurements were carried out on a Thermo Delta V mass spectrometer interfaced with a Flash EA 1112 elemental analyzer at the Institute of Tibetan Plateau Research (ITP), CAS. All samples were measured in duplicate and are reported in per mil units (‰) relative to the Vienna Peedee belemnite (VPDB) standard. Statistical analysis of the δ13CTOC of the duplicate samples and standard indicate the total error to be less than ± 0.2‰. Samples for sporopollen analysis were collected at 2–20 m intervals. Approximate 150–350 g samples were treated with 36% HCl and 40% HF to remove carbonates and silica. Separation of the palynomorphs from the residue was carried out using a 10-μm nylon sieve. Finally, the palynomorphs were mounted in glycerin jelly (for detailed pollen results see Miao, 2010). Lanzhou University and the ITP, performed the CAS Pollen analysis and identification of the samples. 4. Variations of carbon isotopes and pollen records 4.1. Total organic carbon isotopes Figs. 2b and 3b show the δ13CTOC record from the upper Xiejia section and its time series (calculated by linear intropolation between correlated magnetic polarity chrons in Fig. 2f). These demonstrate an overall pattern (first order feature) of a long-term decreasing δ13CTOC trend from about −20‰ in the middle and late Eocene, to −23.5‰ in the early Miocene. The pattern is characterized by five distinctive phases, I to V, of alternating heavier and lighter values with boundaries at ~32.5 Ma, 25 Ma, 22 Ma, and 18 Ma (Figs. 2b and 3b). Phase I spans from ~38 to 32.5 Ma and has the heaviest average δ13CTOC (−21.7‰) in the whole analyzed section, and shows a clear

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long-term decrease trend. Phase II covers between ~32.5 Ma and 25 Ma and shows generally lighter values, with the average δ13CTOC of −22.3‰ and minor fluctuations. Phase III straddles 25–22 Ma and presents generally heavier values with an average δ13CTOC of −21.9‰, with large fluctuations. Phase IV covers ages ~22–18 Ma and exhibits a gradually decreasing trend, with an average δ13CTOC of −22.6‰. Phase V covers the time ~18–17 Ma, during which values return to heavier ones, around −21.8‰ (Figs. 2b and 3b). Short fast decreases of δ13CTOC occur superimposed on the longer trends of the δ13CTOC record, at ~32.5 Ma, ~23 Ma, ~21.5 Ma, and ~18 Ma. These short fast decreases span only about one to several hundred thousands of years, and encompass almost the lightest values (−23‰ or lower) over the entire section (Figs. 2b and 3b). 4.2. Pollen records This study analyzed a total of 59 samples for pollen assemblages. Figs. 2c–e and 3a show the percentages of gymnosperms and angiosperms. Of the gymnosperms, Ephedripites, Abiespollenites, Piceaepollenites, Cedripites, Abietineaepollenites, Pinuspollenites, Podocarpidites, Tsugaepollenites, and others were identified. Identified angiosperms include Fraxinoipollenites, Rutaceoipollenites, Rhoipites, Meliaceoidites, Euphorbiacites, Proteacidites, Lonicerapollis, Jianghanpollis, Potamogetonacidites, Sparganiaceaepollenites, Magnolipollis, Betulaceoipollenites, Alnipollenites, Chenopodipollis, Artemisiaepollenites, Graminidites, Nitrariadites, Cupuliferoipollenites, Quercoidites, Ulmipollenites, Celtispollenites, Salixipollenites, Juglanspollenites, and others. The variations in percentages of the gymnosperms generally show long-term decreasing trends, while the angiosperms show long-term increasing trends, fluctuating oppositely. Obvious short term peaks of the gymnosperms or valleys of the angiosperms occurred at approx. 32 Ma, 25 Ma, 23 Ma, 21 Ma, and 18 Ma (Figs. 2c and d and 3a). Fig. 2e shows the ratio of

Fig. 3. Comparisons in time domain of (a) the percentages of the gymnosperms; (b) the total organic carbon isotope record in the upper Xiejia section; (c) the oxygen isotope record of deep sea sediments at Site 1148 in the South China Sea (Wang et al., 2003b); (d, e) global marine oxygen and carbon isotope records (Zachos et al., 2001); (f) organic carbon isotope records from the Gippsland Basin, Australia (Holdgate et al., 2009); (g) organic carbon isotope records from Central Europe (Bechtel et al., 2008); and (h) the synthesized atmospheric CO2 records from Beerling and Royer (2011).

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gymnosperms/angiosperms with a similar pattern, but with greater amplitude.

5. Discussion 5.1. Sources of organic matter and δ13CTOC composition Total of organic carbon isotope in sediment relates closely to the origin and composition of the organic matter. Based on the photosynthetic pathways used, terrestrial higher plants can be divided into C3, C4, and CAM categories, each with characteristic δ13C values: −20‰ to −35‰ with an average of −27‰ for C3 plants; −9‰ to −17‰ with an average of −13‰ for C4 plants; and intermediate values for CAM plants (succulents such as cacti and some yuccas in extremely arid environment) (see e.g., Deines, 1980; O'Leary, 1981; Farquhar et al., 1982, 1989; Cerling et al., 1997; Tipple and Pagani, 2007). However, lacustrine plants produce large amounts of endogenic organic matter, which differs from the exogenous organic matter of higher terrestrial plants. Great differences in δ13C exist among submerged and emerged plants, phytoplankton, and algae. The δ13C values of submerged plants lie between −12‰ and −20‰, for emerged plants between −24‰ and −30‰, for phytoplankton near −27‰, and algae have overall lower δ13C values (Dana and Deevey, 1960; Smith and Epstein, 1971; Aravena et al., 1992). Therefore, the δ13CTOC compositions can be greatly affected by the contributions of the exogenous and endogenous organic matter in the lake. Hence, we chose to examine the upper Xiejia section, where only one floodplain facies occurs and all the organic matter is exogenous (land sourced). Studies of the n-alkanes components from the Xiejia section have shown that the distribution of n-alkanes ranges from C16 to C35, maximizing at C29 or C31 (Long et al., 2011), confirming that the organic matter came mostly from higher terrestrial plants. The Carbon Preference Index, CPI (the ratio of the sum of high odd numbered carbon peaks C25– C33 to the sum of high even carbon peaks C24–C34: Bray and Evans, 1961), ranges from 1.45 to 6.23, further indicating a higher terrestrial plant origin (Long et al., 2011). The δ13CTOC values of the measured Xiejia section range from − 23.5‰ to −20.0‰, with an average of ~−22.0‰, much higher than that for C4 plants at ~−13‰ (Deines, 1980; Farquhar et al., 1982, 1989; Cerling et al., 1997). Furthermore, C4 plants had not yet appeared in the late Eocene to the early Miocene (Quade et al., 1989; Cerling et al., 1993, 1997), and pollen records of the Xiejia section show little influence from the CAM plants (e.g., Deines, 1980). We regard the δ13CTOC values of the Xiejia section to mainly reflect changes in terrigenous C3 plants.

5.2. Relationships between carbon isotopic composition and pollen records The δ13CTOC and pollen records show a generally good correlation, showing high contents of gymnosperms, or ratios of gymnosperms/ angiosperms, consistent with heavy δ13CTOC values (Fig. 2b–e). Statistical analysis demonstrates positive correlations between the gymnosperm contents and δ13CTOC values in the upper Xiejia section, with a correlation coefficient of 0.69 (Fig. 4a); while the angiosperms contents show clear negative correlations with the δ13CTOC values, with a correlation coefficient of 0.68 (Fig. 4b). Vegetation can be divided into gymnosperms, angiosperms, and ferns. However, the Xiejia section contains a very low fern content (with an average of b 5%). Hoorn et al. (2012) similarly report very low fern content in the Shuiwan section just to the north of the Xiejia section, supporting the dominant role of gymnosperms and angiosperms in the area's vegetation. Their contents not only are reversely correlated, but also generally control the variation of the δ13CTOC in the sediments. The Cenozoic brown coal seams in Australia (Holdgate et al., 2009) and Central Europe (Bechtel et al., 2008) also demonstrate a roughly positive relationship between the gymnosperm contents and δ13CTOC values (Bechtel et al., 2008; Holdgate et al., 2009) (Fig. 3f and g), even though the seams are discontinuous, hard to date, and the pollen-spore records irregular.

5.3. Development of the ecological environment and associated mechanisms Comparisons of δ13CTOC records and the content of the gymnosperms with global temperature records demonstrate great similarities between them (Fig. 3a–d). The late Eocene (Phase I) persistent decrease of the δ13CTOC agrees with the late Eocene global long term cooling. The Oligocene low and relatively stable δ13CTOC values in Phase II (32.5–25 Ma) can readily be correlated with the initiation of the Antarctic ice sheet (the first time in the Cenozoic for a warming earth to turn into a cooling earth). The late Oligocene–earliest Miocene high values of the δ13CTOC in Phase III are probably analogs of the late Oligocene–earliest Miocene global warming. The relative low and high values of the δ13CTOC in Phases IV and V may be correlated with the early Miocene relative cooling interval and the middle Miocene climatic optimum, respectively (Fig. 3). Furthermore, the remarkably rapid decreases of the δ13CTOC at ~ 32.5 Ma, ~ 23 Ma, ~ 21.5 Ma, and ~ 18 Ma in the upper Xiejia section can also be correlated in time with the well known rapid cooling events Oi-1, Mi-1.1, Mi-1a, and Mi-1b, of the early Oligocene and Miocene (Zachos et al., 2001; Wang et al., 2003b; Zachos et al., 2008) (Fig. 3b–d). Detailed paleomagnetic and lithofacies studies have revealed the remarkable

Fig. 4. Scatter plots of the percentage contents for the upper Xiejia section of (a) gymnosperms and (b) angiosperms, with their corresponding δ13CTOC values, showing high correlations between them.

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sedimentary response to the Oi-1 event in the Xining Basin (Dupont-Nivet et al., 2007; Xiao et al., 2010). Therefore, we argue that global temperatures may provide the key factor in controlling the evolution and ecologic environment of gymnosperms and angiosperms. Thus, global temperatures also control the contents of the gymnosperms and angiosperms, the ratio of the gymnosperms to angiosperms, and finally the δ13CTOC values of the organic matter in the sediments in the upper Xiejia section. Indeed, the most striking feature of the evolution of the vegetation on earth is the retreat of gymnosperms and appearance and expansion of angiosperms from the late Mesozoic to the Cenozoic, accompanied by a generally long term global cooling. On the regional scale, the investigation of carbon isotopic compositions of five spruce trees (gymnosperms) from southern Germany reveals that their average δ13C values get heavier as the temperature increases (Schleser et al., 1999) (Fig. 5a). Similar relationships have been found for the gymnosperm fir Abies alba in the famous Black Forest of Germany (Edwards et al., 2000) and for gymnosperm pine trees in other parts of the world (Stuiver and Braziunas, 1987). Studies of the δ13CTOC of the surface soils in the North American Great Plains have shown that the δ13CTOC becomes gradually more negative with increasing latitudes and decreasing temperatures between 30°N and 52°N (Tieszen et al., 1997). Much wider reviews of δ13C values of C3 plants and temperatures corroborate the positive relationships between them (Heaton, 1999; Schleser et al., 1999; McCarroll and Loader, 2004). These general positive relationships between δ13C values of C3 plants and temperature are more robustly supported by the latest investigation of 118 C3 plant species (gymnosperms and angiosperms) along an isoline of 400 mm mean annual precipitation across North China (Wang et al., 2013) (Fig. 5b). Pollen records from the Xiejia section also show a gradually decrease in the thermophilic taxa percentages, indicating a long-term climate cooling trend (Miao, 2010). The coexistence approach (Mosbrugger and Utescher, 1997) of the pollen at the Shuiwan section, Xining Basin, indicates that air temperature changed from about 20 °C to 16 °C, showing a long-term cooling process (Hoorn et al., 2012); although some of the possibly related tectonics may have occurred at about 38 Ma (Dupont-Nivet et al., 2008). The sediment colors (redness) within the Xiejia section become lighter (from red via brownish red to yellowish brown), roughly coupled with the long-term cooling trend in the Xining Basin.

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Based on underlying theory, Farquhar et al. (1982) proposed a carbon isotope fractionation model of C3 plants as follows:     δ13 C ¼ δ13 Ca –a–ðb–aÞ PiðtÞ =Pa or Δ≈δ13 Ca –δ13 C ¼ a þ ðb–aÞ PiðtÞ =Pa ;

where ‘δ13Ca’ is the isotopic composition of CO2 in the ambient atmosphere; ‘a’ represents the isotope fractionation factor of CO2 into plants from the atmosphere (about 4.4‰); ‘b’ is the isotope fractionation factor of carboxylation (27‰ to 30‰); ‘Pi (t)’ is the partial pressure of CO2 in the intercellular spaces, dependent on temperature t; ‘Pa’ is the partial pressure of CO2 in the external atmosphere; and ‘Δ’ is the carbon isotopic discrimination. This model shows that when the temperature declines, plant photosynthesis efficiency decreases, resulting in higher values of Pi (t) and Pi (t)/Pa, followed by increasing plant carbon isotope fractionation, and finally causing decreases of the δ13C within the plant. This model also indicates that decreasing of the partial pressure of atmospheric CO2 (Pa) can lead to higher Pi (t)/Pa values as well, thus also causing the δ13C in the plants to decrease. Marine carbon isotope records indicate that the Cenozoic global carbon reservoir remains relatively stable near the value of 1‰, except when carbon isotope excursions occur associated with global climatic events and serious corresponding biological events (Vincent and Berger, 1985; Zachos et al., 2001) (Fig. 3e). This pattern of varying marine carbon isotope records differs from those of the continental carbon isotope records from flood plain sediments in the Xining Basin (Fig. 3b) and seams also to those in Australia and Europe (Bechtel et al., 2008; Holdgate et al., 2009) (Fig. 3f and g), showing that the balance of the global carbon reservoir is not the major factor influencing the variation of these continental carbon isotopes. Synthesized records of geological partial pressure of atmospheric CO2 (Pa) generally show a long term decline during the Cenozoic (Beerling and Royer, 2011) (Fig. 3h), suggesting that Pa is one of the factors controlling the long term evolution of continental carbon isotope compositions. These geological records agree well with Farquhar et al.'s (1982) theoretical model prediction (above). Since the Pa is also a major contributor to global temperature, we regard simply that the Cenozoic continental carbon isotope variation is mostly controlled by the global temperature change. The onset and evolution of the Asian monsoons and aridification provided major climatic changes for the earth. Increasing lines of

Fig. 5. (a) Relationship between mean July temperatures and δ13C values in gymnosperm wood cellulose of tree rings (average of five spruce trees from southern Germany) during the period 1957–1992. Correlation coefficients (r), and number of analyses (n), based on regression and standard deviation analyses by Schleser et al. (1999); (b) δ13C variations in 118 C3 plant species (gymnosperms and angiosperms), with the annual mean temperature along a temperature gradient in north China. The hollow squares represent site-averaged δ13C values after precipitation-correction and the vertical bars indicate the δ13C variability within the samples of each site (±1 standard deviation) (Wang et al., 2013).

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evidence reveal that the East Asian monsoon may have originated in the late Oligocene to early Miocene (e.g., Liu et al., 1998; Shi et al., 1999; Guo et al., 2002; Sun and Wang, 2005; Qiang et al., 2011), causing increased rainfall in East Asia. Investigations of carbon isotopes of modern C3 plants in the North China loess area show that the average δ13C values decrease with the increasing of annual rainfall (Wang et al., 2003a) (Fig. 6). Such relationships have been widely reported from various parts of the world (e.g., Stewart et al., 1995; Schulze et al., 2006; Diefendorf et al., 2010; Kohn, 2010). However, in our study, we found no clear signals of significant decreases of the δ13CTOC from the late Oligocene and early Miocene, where instead, δ13CTOC values appear to mimic global temperatures (Fig. 3b and d). This implies that the East Asian monsoon may have not reached this region during that time period, or alternatively, that our low resolution sedimentary carbon isotope and pollen records could not detect this event clearly. Many other factors, however, such as pedogenesis and the local environment, can modify the carbon isotopic composition of TOC in sediments (e.g., Melillo et al., 1989; Connin et al., 2001; Boström et al., 2007; Liu and Huang, 2008). But we believe all these factors only exert secondary influences on the sedimentary TOC carbon isotope composition because: 1) many of these factors inter-relate and are subject to the same climatic environmental changes that the plant composition reflects, and 2) pedogenesis mostly increases δ13C values of organic matter (Melillo et al., 1989; Connin et al., 2001; Boström et al., 2007), but does not change its pattern identically for C3 and C4 plants (Cerling et al., 1993, 1997). This means that the evolution of climate over geologic time controls the trends and major phases that determine the isotopic composition of plants, and thus the major patterns of the isotopic composition of the sedimentary organic carbon. Other factors only modify these major carbon isotope patterns in sediments to a lesser extent, superimposing minor fluctuations. Therefore, we hope to continue this effort in the future, expanding the records to the late Miocene and Pliocene, to clearly reveal the signals and evolution of these events in the Xining Basin. 6. Conclusions This study obtained a late Eocene to early Miocene continental sedimentary δ13CTOC record from the Xining Basin, NW China. The δ13CTOC values from the flood plain sediments in the basin range between − 23.5‰ and − 20.0‰ and show a long-term decreasing trend over time. Five distinct alternating phases of heavier and lighter values superimpose on this long term trend, with short periods with very low δ13CTOC values occurring at approx. 32.5 Ma, 23 Ma,

Fig. 6. The variation of δ13C in C3 plants (defined in the Discussion section) with mean annual precipitation in the North China loess region (each bar represents the mean δ13C value and standard error of C3 plants in different sampled regions) (Wang et al., 2003a).

21.5 Ma, and 18 Ma. These characteristic variations generally correlate with the changes of the gymnosperm concentrations, and the ratio of gymnosperm/angiosperm (G/A), i.e., higher δ13CTOC values correlate with higher gymnosperm concentrations (and higher G/A), as well as with global temperatures, both in long term trends and during short term fluctuations or events. For example, short periods of low δ13CTOC values in the earliest Oligocene and early Miocene at ~32.5 Ma, 23 Ma, 21.5 Ma, and 18 Ma correlate with the rapid cooling events of Oi-1, Mi-1.1, Mi-1a, and Mi-1b; while intervals of high δ13CTOC values occur at approx. 25–23 Ma, 23–21.5 Ma, and 18–b17 Ma, corresponding to the late Oligocene and early Miocene periods of Climatic Warming and the Middle Miocene Climatic Optimum, respectively. Global temperatures might be the major factor controlling the evolution of gymnosperm concentrations (and the G/A ratios) that determines the δ13CTOC in the sediments. Either the East Asian monsoon did not reach this region during the deposition of these sediments, or its signals cannot clearly be identified from our paleoecologic records due to their low resolution.

Acknowledgment This work is co-supported by the Strategic Priority Research Program of the Chinese Academy of Sciences (Grant No. XDB03020402), the National Basic Research Program of China (Grant Nos. 2013CB956400, 2010CB833401), and the National Natural Science Foundation (NSFC Grant Nos. 41021001, 41272183, 41272128, 41172153). We thank Dai Shuang, Zhang Zhigao, Yan Xiaoli, and Xu Li for field sampling assistance. Thanks are also due to Bai Yan, Yang Yibo, Zhu Zhiyong, Qu Dongmei, and Li Xiangyu for their assistance in the experiments and their valuable discussions.

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