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C4 vegetation evolution over the last 7.0 Myr in the Chinese Loess Plateau: evidence from pedogenic carbonate δ13C
Palaeogeography, Palaeoclimatology, Palaeoecology 160 (2000) 291–299 www.elsevier.nl/locate/palaeo
C /C vegetation evolution over the last 7.0 Myr in 3 4 the Chinese Loess Plateau: evidence from pedogenic carbonate d13C Z.L. Ding *, S.L. Yang Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, People’s Republic of China Received 16 September 1999; received in revised form 3 February 2000; accepted for publication 10 February 2000
Abstract The stable carbon and oxygen isotopic compositions of soil carbonate were measured on an eolian loess and red clay sequence at Lingtai, the Chinese Loess Plateau. This sequence is composed of 130 m of Tertiary red clay deposits with a basal age of 7.05 Ma overlain by 175 m of Pleistocene loess. In the field we identified ca. 110 carbonate nodule horizons in the red clay and 27 nodule horizons in the loess. These carbonate nodule horizons are formed by leaching and re-precipitation of carbonate from the eolian material. The d13C record of soil carbonate indicates a major expansion of C plants at ca. 4.0 Myr in the Loess Plateau. This event is comparable in timing with the expansion of 4 C plants in northern North America (Cerling et al., 1997. Nature 389, 153–158) but is ca. 3 million years later than 4 the C biomass expansion in Pakistan (Quade et al., 1989. Nature 342, 163–166). The pedogenic characteristics of 4 the soils and the d18O record in the red clay suggest that the C plant expansion in the Loess Plateau was not driven 4 by local climatic changes, which may support Cerling et al.’s (1997) assertion that the decline of atmospheric CO 2 levels in the Neogene is responsible for this global vegetation change. Our record also implies that the Tibetan Plateau could have been uplifted to a critical height in the late Miocene, thus resulting in the formation of the atmospheric Great East-Asia Trough. © 2000 Elsevier Science B.V. All rights reserved. Keywords: atmospheric CO levels; carbon isotope record; eolian deposits; paleovegetation 2
1. Introduction The relative proportion of C and C plants in 3 4 local biomass can be inferred from the d13C compositions of soil organic matter and pedogenic carbonate (Cerling, 1984; Cerling et al., 1989). Pure C and C plants have d13C values ranging 3 4 from ca. −22 to −30‰ and −10 to −14‰, * Corresponding author. Tel.: +86-10-6200-8111; fax: +86-10-6491-9140. E-mail address: [email protected] (Z.L. Ding)
respectively (Bender, 1971; Winter et al., 1976; Brown, 1977; Vogel et al., 1978; Farquhar et al., 1989). Soil organic matter preserves this isotopic distinction with little or no isotopic fractionation, but pedogenic carbonate and carbonate in the dental enamel of fossil herbivores are enriched in 13C by ca. 14 (25°C ) to 17‰ (0°C ) with respect to source carbon (Lee-Thorp and van der Merwe, 1987; Cerling et al., 1989; Wang et al., 1994). Therefore, under conditions of moisture stress, d13C values of −8 and +4‰ in pedogenic carbonate indicate nearly pure C and C ecosystems, 3 4
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respectively ( Ehleringer et al., 1986; Ehleringer and Cooper, 1988; Farquhar et al., 1989). Recently, Cerling et al. (1997) have reported changes in the carbon isotope ratios of fossil tooth enamel from Pakistan, Africa, North America and South America since the late Miocene. The authors determined that there was a global expansion of C biomass between 8 and 4 million years (Myr) 4 ago, with the C plant expansion occurring first in 4 low latitudes and later in higher latitudes. In western Europe, there is no evidence suggesting a significant C component in the diet of large 4 mammals at any time. To explain this variable spatial–temporal pattern, Cerling et al. (1997) developed a model in which both atmospheric CO level and temperature are thought to have 2 controlled the global spatial and temporal evolution of the C /C plants. As the continental weath3 4 ering increased, due to the quick uplifting of the Tibetan Plateau in the Neogene (Raymo and Ruddiman, 1992), atmospheric CO concentration 2 has been lowered to a critical level that favors C 4 over C plants in the low latitudes with an elevated 3 temperature. With the further decrease in atmospheric CO levels, the crossover point favoring 2 C over C plants will be reached later in higher 4 3 latitudes with relatively low temperatures (Cerling et al., 1997). This means that d13C records of geological sediments are ideal in reconstructing the history of mutual interactions between the lithosphere, biosphere and atmosphere of the Earth system. However, this atmospheric CO -climate2 ecosystem linkage is challenged by two new pieces of evidence from marine sediments. Based on boron isotope composition measurements of planktonic foraminifera from tropical Pacific Ocean sediments, Pearson and Palmer (1999) established a pH profile for ancient ocean surface water, which suggests that atmospheric CO partial 2 pressure during the middle Eocene was probably similar to modern concentrations or only slightly higher. This means that the extreme warmth of middle Eocene environment might not be caused mainly by high atmospheric CO level. A pCO 2 2 record derived from alkenone d13C measurements (Pagani et al., 1999) shows that atmospheric pCO was surprisingly low (180–300 ppmv) over 2
the interval of 15 to 9 Myr and stabilized at preindustrial values by 9 million years ago. Absence of evidence for major changes in pCO 2 during the late Miocene led Pagani et al. (1999) to suggest that the sudden expansion of C vegeta4 tion at ca. 7 Myr could be initiated by the development of low-latitude seasonal aridity and changes in growing conditions on a global scale, rather than a decrease in pCO . 2 As plant-type distributions are determined by several factors, more records of paleovegetation changes are needed to understand the mechanisms for the late Miocene C plant expansion. The 4 Chinese Loess Plateau is located to the northeast of the Tibetan Plateau, where the climate is controlled essentially by the Asian monsoon system. Many studies have shown that the wind-blown loess–soil sequence in the Loess Plateau is one of the most complete records of late Cenozoic climatic changes in the world (Liu, 1985; Kukla and An, 1989; Rutter et al., 1991; Ding et al., 1999). The Asian monsoon system has experienced a longterm evolution with the uplift of the Tibetan Plateau during the late Cenozoic, which may in turn result in fundamental changes in the ecosystem of the Loess Plateau. In this context, paleovegetation change records from the eolian successions and comparison with other records in the world would help to clarify the factors causing the global expansion of C ecosystems 4.0– 4 7.0 Myr ago. In this paper we report a 7.0 Myr d13C record from soil carbonate nodules formed in an eolian loess and red clay sequence at Lingtai, the Chinese Loess Plateau.
2. Setting and stratigraphy Lingtai (Fig. 1) is situated in the middle of the Loess Plateau at an elevation of ca. 1340 m above sea level ( latitude 35°00∞33◊N, longitude 107°30∞33◊E ). At present, the mean annual temperature and precipitation at Lingtai are ca. 9.1°C and 600 mm, respectively. The averaged temperature of July, the warmest month, is ca. 22.3°C. The modern climate of this region is essentially controlled by the Asia monsoon system. In the
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Fig. 1. Schematic map showing the locations where the d13C records used in the paper are derived. $, Lingtai section of the Chinese Loess Plateau; +, the Siwalik sediments in Pakistan. The dashed line in the North American continent indicates 37°N latitude.
summer season, the East-Asia summer monsoon brings warm, moisture-laden air masses from tropical oceans to the Loess Plateau, resulting in heavy rainfall in the region. Over 50% of the annual precipitation occurs from July to September. During the winter season, the winter monsoon winds flowing from the Siberian region prevail in the Loess Plateau, leading to a dry and cold climate. Recently, a loess and red clay sequence, ca. 305 m thick, was discovered and studied at Lingtai (Ding et al., 1998a, 1999). The loess deposit is ca. 175 m thick, and contains ca. 40 paleosols. Field observation and magnetic susceptibility and median grain size (Md) records (Fig. 2) show that the Lingtai loess and soil sequence can be well correlated with the standard sections of the Loess Plateau (Liu, 1985; Kukla and An, 1989; Rutter et al., 1991). Paleomagnetic studies indicate a basal age of ca. 2.6 Myr for the Lingtai loess ( Fig. 2). The paleosols within the loess are generally characterized by a brownish or reddish color, and substantial clay skins. Horizontal carbonate nodule horizons are generally not observed immediately below the pedogenic B horizons for the paleosols in the upper part of the loess. However, scattered nodules are common at the base of the soils. Most of the paleosols in the lower part of the loess sequence have ca. 20–80 cm thick nodule horizons (Ding et al., 1999). The red clay underlying the loess is ca. 130 m
thick. Paleomagnetic measurements (Ding et al., 1998a) suggest that it accumulated from 2.6 to 7.05 Myr (Fig. 2). The magnetic polarity stratigraphy shown in Fig. 2 was established by analyses of 625 orientated samples at an interval of 15– 25 cm from the red clay sequence. The magnetic remanence was measured at the Paleomagnetism Laboratory in the Institute of Geology and Geophysics, Chinese Academy of Sciences, with a 2G three-axis cryogenic magnetometer. Remanence data are obtained after alternative demagnetization, usually at 20 or 25 mT. In most cases, an alternative demagnetization at 15–30 mT leads to an unambiguous polarity assignment. The Matuyama/Gauss (M/G) reversal occurred in the upper part of the oldest loess of L33 ( Fig. 2). Within the Gauss Chron, there are two short reversed zones occurring respectively at depths of 190 and 196 m. They are interpreted as subchrons 2An.1r and 2An.2r (i.e. the Kaena and Mammoth Events). The 2An/2Ar boundary (Gauss/Gilbert boundary) falls near a depth of 206 m, and the bottom of 2Ar can be defined at the depth of ca. 219.5 m. Below ca. 220 m, the polarity zonation is readily obtained from the remanence data, which can be easily correlated to the standard polarity time scale (Cande and Kent, 1995). Sedimentological and geochemical studies have demonstrated that this Tertiary red clay is windblown in origin as is the overlying Pleistocene loess (Ding et al., 1998b). One of the most striking
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Fig. 2. Magnetic susceptibility and median grain size (Md ) records of the Lingtai red clay–loess sequence with the magnetic reversal polarity [modified from Ding et al. (1998a)]. Most of the soil units (Si) and some thick loess beds (Li) are indicated.
features of the red clay sequence is the existence of many horizontal carbonate nodule horizons. The thickness of these horizons ranges from 10 cm to >100 cm, with most of the nodules <10 cm in diameter. The groundmass within most of the nodule horizons is reddish, weathered soil material. In addition to the nodule horizons, individual nodules are commonly scattered throughout the red clay. The red clay between the nodule horizons
shows a redder color than that of the paleosols in the overlying Pleistocene loess. The horizons have been subjected to relatively strong pedogenesis, as indicated by the relative abundance of clay and Fe–Mn skins, and are identified as pedological B horizons. A remarkable difference between the soil horizons of the red clay and loess records is that pedogenic A horizons are generally lacking in the red clay sequence, whereas a complete A–B–C
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sequence is readily recognized in the loess. A total of ca. 120 soil B horizons and ca. 110 intervening carbonate nodule horizons are identified in the Lingtai red clay sequence (Ding et al., 1999).
3. Sampling and analysis In the Pleistocene loess sequence, we took one carbonate nodule from each of the carbonate nodule horizons if nodules existed at the base of the paleosols; 27 nodule samples were collected in the loess–soil record. In the red clay sequence, one sample was taken from each thin nodule layer, while in every thick nodule horizon, two samples from the upper and lower parts were collected. A total of 178 samples were taken in the red clay record. Nodules were ground to a powder, and converted to CO with 100% phosphoric acid. Carbon 2 and oxygen isotopic ratios were both measured in a DeltaS gas-source mass spectrometer. Isotopic results are presented in the usual d notation as the permil (‰) deviation of the sample CO from the 2 PDB standard
A
R
B
sample −1 ×1000 R standard where R=13C/12C or 18O/16O. Replicate analyses (n=13) of a single, homogenized sample show that this procedure yields a standard deviation of ±0.05 and ±0.06‰, respectively, for d13C and d18O measurements.
d=
4. Results The carbon and oxygen isotopic records of soil carbonate nodules and magnetic susceptibility are shown in Fig. 3 for the Lingtai loess and red clay sequence. The time scale was established by linear interpolation between paleomagnetic reversal boundaries (Ding et al., 1998a). The carbon isotope record can be grouped roughly into three phases. From 7.05 to ca. 4.0 Myr, the d13C values of soil carbonate fall between −9.6 and −7.8‰, indicating a C -dominated ecosystem in the Lingtai 3 area. At ca. 4.0 Myr, the stable carbon isotopic
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compositions of soil carbonate begin to shift toward more positive values. Most of the d13C values lie between −7.8 and −5.3‰ from 4.0 to 2.0 Myr (Fig. 3). A single sample at the 3.4 Myr level yielded the d13C value of −4.7‰, representing the highest d13C value in the entire red clay sequence. Over this interval, a mixed ecosystem of C and C plants can be interpreted from the 3 4 stable carbon isotopic compositions of soil carbonate. Therefore, the expansion of C plants on the 4 Loess Plateau occurred at ca. 4.0 Myr, but at no time did C plants dominate the vegetation on the 4 plateau. From 2.0 Myr to the Holocene, most of the d13C values in soil carbonate range from −6.5 to −9.2‰ with an averaged value of ca. −8.0‰ ( Fig. 3). Compared to the interval of 4.0 to 2.0 Myr, C plants in the Loess Plateau ecosystem 3 increased significantly in the Pleistocene. However, in the youngest two paleosols (S1 and S2), soil carbonate is enriched in 13C (Fig. 3), suggesting another expansion of C plants on the plateau. 4 In the Pakistan stable isotopic records (Quade et al., 1989; Quade and Cerling, 1995), the shift of d13C compositions toward more positive values at ca. 7.0 Myr was accompanied by a clear increase in the d18O record. This feature is not found in our d18O record. Changes in soil carbonate d18O compositions range scatteringly from −11.8 to −8.0‰ (Fig. 3). There is an overall trend of d18O compositions toward more positive values from bottom to top of the loess and red clay sequence. Changes in d18O compositions of soil carbonate may be controlled by complicated factors, including temperature, regional moisture sources, rainfall amount, and seasonality of precipitation (Quade and Cerling, 1995). The causes for the trend of d18O changes of soil carbonate in the Loess Plateau are unclear. We preliminarily think that this may be partly related with the increase in seasonality during the past 7.0 Ma due to the long-term intensification of both the summer and winter monsoon systems over East Asia (Liu and Ding, 1993; Ding et al., 1999).
5. Discussion and conclusions Our carbon isotopic record indicates a significant expansion of C plants at ca. 4.0 Myr on the 4
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Fig. 3. Magnetic susceptibility and carbon and oxygen isotope records plotted against time. The timescale is constructed by linear interpolation between magnetic reversal boundaries.
Chinese Loess Plateau. A major question concerning this event is whether or not it was driven by local climatic changes. The pedogenic characteristics of the soils within the red clay sequence have recorded the history of the East-Asia summer monsoon evolution during the late Miocene and Pliocene in the Loess Plateau. Recent observations (Ding et al., 1999) show that the Lingtai red clay sequence can be subdivided into five units according to soil features. The first ( lowest) unit accumulated from 7.05 to 6.2 Myr. Soils in this unit are characterized by a reddish brown color and a weak subangular blocky structure with few clay skins, suggesting a relatively weak summer monsoon. The second unit, formed during 6.2 to 5.5 Myr, contains soils showing a moderate subangular blocky structure and some or common clay skins, implying a strengthened summer monsoon relative
to the earlier time. The soils of the third unit (5.5 to 3.85 Myr) display many thick clay skins and many dark Fe–Mn films, interpreted as having developed under strong summer monsoon conditions. The summer monsoon was significantly weakened over the interval of 3.85 to 3.15 Ma (the fourth unit), as suggested by greatly reduced clay skins on the pedogenic structure surfaces. In the fifth unit (3.15 to 2.6 Myr), translocated clay skins in the soils again significantly increase, implying an intensified summer monsoon climate. This interpreted monsoon evolution suggests that there is no significant local climatic event at ca. 4.0 Myr, and thus it seems unlikely that the ecological change registered in the d13C record was be forced solely by changes in local climatic conditions. This idea may be supported also by the soil carbonate d18O record ( Fig. 3) which shows an almost linear
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enrichment in 18O from bottom to top of the sequence. No peculiar change ca. 4.0 Myr is found in the record. Therefore, if the d18O trend indeed represents the increase in seasonality of precipitation, this factor is not the most important in driving the middle Pliocene C plant expansion on 4 the Loess Plateau. The modern spatial distribution of C and C 4 3 grasses shows that C grasses dominate in the 4 tropical and subtropical regions, that the transition to C grasses takes place between ca. 30 and 45° 3 latitudes, and that C grasses dominate at high 3 latitudes. According to the model developed by Cerling et al. (1997), the photosynthetic efficiency of C grasses relative to C grasses varies with 3 4
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both atmospheric CO levels and temperature. 2 This model predicts that with the continuous decrease in atmospheric CO levels during the late 2 Cenozoic, as deduced from geological records, expansion of C plants will occur first in the 4 tropical and subtropical regions, and later in the transitional latitudes because of lower temperatures. Fig. 4 shows the comparison of d13C records between Pakistan, the Chinese Loess Plateau and northern North America (>37°N ). It is evident that the Loess Plateau and northern North America experienced a synchronous expansion of C plants 4.0 Myr ago, ca. 3 million years 4 later than the C expansion in Pakistan. This may 4 support the model of Cerling et al. (1997),
Fig. 4. Comparison of the C -plant expansion event between the Loess Plateau (pedogenic carbonate d13C ), Pakistan (pedogenic 4 carbonate d13C ) and northern North America ( Equid d13C ). The arrows indicate the positions of the event. The Pakistan data are from Quade et al. (1989) and the northern North America data from Cerling et al. (1997).
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although direct evidence for high concentration of atmospheric CO in the Neogene is lacking (Pagani 2 et al., 1999; Pearson and Palmer, 1999). The Pakistan carbon isotopic records are derived from between 32 and 33°N, only 2–3° of latitude lower than the Lingtai section (35°N ). The difference in the timing of the C -plant expan4 sion between the two regions is, however, as great as ca. 3 million years. On the other hand, northern North America has substantially higher latitudes than Lingtai, but the two regions witnessed an approximately synchronous C -plant expansion. 4 This evidence suggests an anomalously low temperature of the Loess Plateau ca. 7.0 Myr ago, according to the Cerling et al. (1997) model. At present, cold air masses propagated onto the Plateau are mainly delivered by the northerly winter monsoon and the westerlies. Changes in grain size values of four red clay–loess sections demonstrate that the southward gradient of decreasing grain size in the Tertiary red clay on the Plateau is negligible (Ding et al., 1998b), whereas there is a strong negative gradient in particle size from north to south in the Loess Plateau for sediments younger than 2.6 Ma. This decrease in loess particle size is consistent with the direction of the winter monsoonal winds flowing from Siberia. It is suggested that the winter monsoon may not have formed until ca. 2.6 Myr ago (Ding et al., 1998b). In this case, the lowered temperature background over the Loess Plateau during the late Miocene may be caused by increased propagation of cold air masses by the westerlies. Meteorological observations ( Wallace, 1983) show that at 500 mb, there is a broad trough, called the Great East-Asia Trough (GEAT ), centered about the longitudes of Japan and forming a broad ridge to the west, which produces northwesterly winds in the upper troposphere over China. The GEAT is the strongest of the quasistationary waves on the globe, which is partly responsible for the relatively low temperature over eastern and central China. GCM modeling experiments (Manabe and Terpstra, 1974; Kutzbach et al., 1989) show that this stationary wave over Asia is caused by the uplifted Tibetan Plateau. The greatly-delayed shift of C plants in the Loess 4 Plateau suggests that the Tibetan Plateau may have reached a critical height in the late Miocene,
thus leading to the formation of the GEAT. This is consistent with the observations of several geological studies (Mercier et al., 1987; Copeland and Harrison, 1990; Harrison et al., 1993). Our d13C record shows a subsequent diminishment of the C -plant component in the period of 4 2 Myr to the late Pleistocene ( Fig. 3). This phenomenon is also observed in the dental enamel carbon isotope record of fossil herbivores from northern North America (Cerling et al., 1997). To date, the cause for this is unknown. As no evidence is obtained for a subsequent increase in atmospheric CO levels during the Pleistocene, we pro2 pose that decrease in temperature could be the main factor for this event. Dramatic expansion of polar ice sheets in the Northern Hemisphere during late Pliocene and early Pleistocene has been indicated by many geological records (Ruddiman et al., 1989; Raymo et al., 1989; Shackleton et al., 1990). Large-scale ice sheet expansion in the polar area will lead to a decrease in temperature background and an increase in north–south temperature gradient, thereby causing a southward shift of C taxa in the Northern Hemisphere. 3 Acknowledgements This research is supported by the NNSF of China (49525203) and CAS ( KZ951-A1-402). The authors are greatly indebted to Dr Eve Arnold for critical comments on the manuscript.
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C /C vegetation evolution over the last 7.0 Myr in 3 4 the Chinese Loess Plateau: evidence from pedogenic carbonate d13C Z.L. Ding *, S.L. Yang Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, People’s Republic of China Received 16 September 1999; received in revised form 3 February 2000; accepted for publication 10 February 2000
Abstract The stable carbon and oxygen isotopic compositions of soil carbonate were measured on an eolian loess and red clay sequence at Lingtai, the Chinese Loess Plateau. This sequence is composed of 130 m of Tertiary red clay deposits with a basal age of 7.05 Ma overlain by 175 m of Pleistocene loess. In the field we identified ca. 110 carbonate nodule horizons in the red clay and 27 nodule horizons in the loess. These carbonate nodule horizons are formed by leaching and re-precipitation of carbonate from the eolian material. The d13C record of soil carbonate indicates a major expansion of C plants at ca. 4.0 Myr in the Loess Plateau. This event is comparable in timing with the expansion of 4 C plants in northern North America (Cerling et al., 1997. Nature 389, 153–158) but is ca. 3 million years later than 4 the C biomass expansion in Pakistan (Quade et al., 1989. Nature 342, 163–166). The pedogenic characteristics of 4 the soils and the d18O record in the red clay suggest that the C plant expansion in the Loess Plateau was not driven 4 by local climatic changes, which may support Cerling et al.’s (1997) assertion that the decline of atmospheric CO 2 levels in the Neogene is responsible for this global vegetation change. Our record also implies that the Tibetan Plateau could have been uplifted to a critical height in the late Miocene, thus resulting in the formation of the atmospheric Great East-Asia Trough. © 2000 Elsevier Science B.V. All rights reserved. Keywords: atmospheric CO levels; carbon isotope record; eolian deposits; paleovegetation 2
1. Introduction The relative proportion of C and C plants in 3 4 local biomass can be inferred from the d13C compositions of soil organic matter and pedogenic carbonate (Cerling, 1984; Cerling et al., 1989). Pure C and C plants have d13C values ranging 3 4 from ca. −22 to −30‰ and −10 to −14‰, * Corresponding author. Tel.: +86-10-6200-8111; fax: +86-10-6491-9140. E-mail address: [email protected] (Z.L. Ding)
respectively (Bender, 1971; Winter et al., 1976; Brown, 1977; Vogel et al., 1978; Farquhar et al., 1989). Soil organic matter preserves this isotopic distinction with little or no isotopic fractionation, but pedogenic carbonate and carbonate in the dental enamel of fossil herbivores are enriched in 13C by ca. 14 (25°C ) to 17‰ (0°C ) with respect to source carbon (Lee-Thorp and van der Merwe, 1987; Cerling et al., 1989; Wang et al., 1994). Therefore, under conditions of moisture stress, d13C values of −8 and +4‰ in pedogenic carbonate indicate nearly pure C and C ecosystems, 3 4
0031-0182/00/$ - see front matter © 2000 Elsevier Science B.V. All rights reserved. PII: S0 0 3 1 -0 1 8 2 ( 0 0 ) 0 0 07 6 - 6
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respectively ( Ehleringer et al., 1986; Ehleringer and Cooper, 1988; Farquhar et al., 1989). Recently, Cerling et al. (1997) have reported changes in the carbon isotope ratios of fossil tooth enamel from Pakistan, Africa, North America and South America since the late Miocene. The authors determined that there was a global expansion of C biomass between 8 and 4 million years (Myr) 4 ago, with the C plant expansion occurring first in 4 low latitudes and later in higher latitudes. In western Europe, there is no evidence suggesting a significant C component in the diet of large 4 mammals at any time. To explain this variable spatial–temporal pattern, Cerling et al. (1997) developed a model in which both atmospheric CO level and temperature are thought to have 2 controlled the global spatial and temporal evolution of the C /C plants. As the continental weath3 4 ering increased, due to the quick uplifting of the Tibetan Plateau in the Neogene (Raymo and Ruddiman, 1992), atmospheric CO concentration 2 has been lowered to a critical level that favors C 4 over C plants in the low latitudes with an elevated 3 temperature. With the further decrease in atmospheric CO levels, the crossover point favoring 2 C over C plants will be reached later in higher 4 3 latitudes with relatively low temperatures (Cerling et al., 1997). This means that d13C records of geological sediments are ideal in reconstructing the history of mutual interactions between the lithosphere, biosphere and atmosphere of the Earth system. However, this atmospheric CO -climate2 ecosystem linkage is challenged by two new pieces of evidence from marine sediments. Based on boron isotope composition measurements of planktonic foraminifera from tropical Pacific Ocean sediments, Pearson and Palmer (1999) established a pH profile for ancient ocean surface water, which suggests that atmospheric CO partial 2 pressure during the middle Eocene was probably similar to modern concentrations or only slightly higher. This means that the extreme warmth of middle Eocene environment might not be caused mainly by high atmospheric CO level. A pCO 2 2 record derived from alkenone d13C measurements (Pagani et al., 1999) shows that atmospheric pCO was surprisingly low (180–300 ppmv) over 2
the interval of 15 to 9 Myr and stabilized at preindustrial values by 9 million years ago. Absence of evidence for major changes in pCO 2 during the late Miocene led Pagani et al. (1999) to suggest that the sudden expansion of C vegeta4 tion at ca. 7 Myr could be initiated by the development of low-latitude seasonal aridity and changes in growing conditions on a global scale, rather than a decrease in pCO . 2 As plant-type distributions are determined by several factors, more records of paleovegetation changes are needed to understand the mechanisms for the late Miocene C plant expansion. The 4 Chinese Loess Plateau is located to the northeast of the Tibetan Plateau, where the climate is controlled essentially by the Asian monsoon system. Many studies have shown that the wind-blown loess–soil sequence in the Loess Plateau is one of the most complete records of late Cenozoic climatic changes in the world (Liu, 1985; Kukla and An, 1989; Rutter et al., 1991; Ding et al., 1999). The Asian monsoon system has experienced a longterm evolution with the uplift of the Tibetan Plateau during the late Cenozoic, which may in turn result in fundamental changes in the ecosystem of the Loess Plateau. In this context, paleovegetation change records from the eolian successions and comparison with other records in the world would help to clarify the factors causing the global expansion of C ecosystems 4.0– 4 7.0 Myr ago. In this paper we report a 7.0 Myr d13C record from soil carbonate nodules formed in an eolian loess and red clay sequence at Lingtai, the Chinese Loess Plateau.
2. Setting and stratigraphy Lingtai (Fig. 1) is situated in the middle of the Loess Plateau at an elevation of ca. 1340 m above sea level ( latitude 35°00∞33◊N, longitude 107°30∞33◊E ). At present, the mean annual temperature and precipitation at Lingtai are ca. 9.1°C and 600 mm, respectively. The averaged temperature of July, the warmest month, is ca. 22.3°C. The modern climate of this region is essentially controlled by the Asia monsoon system. In the
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Fig. 1. Schematic map showing the locations where the d13C records used in the paper are derived. $, Lingtai section of the Chinese Loess Plateau; +, the Siwalik sediments in Pakistan. The dashed line in the North American continent indicates 37°N latitude.
summer season, the East-Asia summer monsoon brings warm, moisture-laden air masses from tropical oceans to the Loess Plateau, resulting in heavy rainfall in the region. Over 50% of the annual precipitation occurs from July to September. During the winter season, the winter monsoon winds flowing from the Siberian region prevail in the Loess Plateau, leading to a dry and cold climate. Recently, a loess and red clay sequence, ca. 305 m thick, was discovered and studied at Lingtai (Ding et al., 1998a, 1999). The loess deposit is ca. 175 m thick, and contains ca. 40 paleosols. Field observation and magnetic susceptibility and median grain size (Md) records (Fig. 2) show that the Lingtai loess and soil sequence can be well correlated with the standard sections of the Loess Plateau (Liu, 1985; Kukla and An, 1989; Rutter et al., 1991). Paleomagnetic studies indicate a basal age of ca. 2.6 Myr for the Lingtai loess ( Fig. 2). The paleosols within the loess are generally characterized by a brownish or reddish color, and substantial clay skins. Horizontal carbonate nodule horizons are generally not observed immediately below the pedogenic B horizons for the paleosols in the upper part of the loess. However, scattered nodules are common at the base of the soils. Most of the paleosols in the lower part of the loess sequence have ca. 20–80 cm thick nodule horizons (Ding et al., 1999). The red clay underlying the loess is ca. 130 m
thick. Paleomagnetic measurements (Ding et al., 1998a) suggest that it accumulated from 2.6 to 7.05 Myr (Fig. 2). The magnetic polarity stratigraphy shown in Fig. 2 was established by analyses of 625 orientated samples at an interval of 15– 25 cm from the red clay sequence. The magnetic remanence was measured at the Paleomagnetism Laboratory in the Institute of Geology and Geophysics, Chinese Academy of Sciences, with a 2G three-axis cryogenic magnetometer. Remanence data are obtained after alternative demagnetization, usually at 20 or 25 mT. In most cases, an alternative demagnetization at 15–30 mT leads to an unambiguous polarity assignment. The Matuyama/Gauss (M/G) reversal occurred in the upper part of the oldest loess of L33 ( Fig. 2). Within the Gauss Chron, there are two short reversed zones occurring respectively at depths of 190 and 196 m. They are interpreted as subchrons 2An.1r and 2An.2r (i.e. the Kaena and Mammoth Events). The 2An/2Ar boundary (Gauss/Gilbert boundary) falls near a depth of 206 m, and the bottom of 2Ar can be defined at the depth of ca. 219.5 m. Below ca. 220 m, the polarity zonation is readily obtained from the remanence data, which can be easily correlated to the standard polarity time scale (Cande and Kent, 1995). Sedimentological and geochemical studies have demonstrated that this Tertiary red clay is windblown in origin as is the overlying Pleistocene loess (Ding et al., 1998b). One of the most striking
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Fig. 2. Magnetic susceptibility and median grain size (Md ) records of the Lingtai red clay–loess sequence with the magnetic reversal polarity [modified from Ding et al. (1998a)]. Most of the soil units (Si) and some thick loess beds (Li) are indicated.
features of the red clay sequence is the existence of many horizontal carbonate nodule horizons. The thickness of these horizons ranges from 10 cm to >100 cm, with most of the nodules <10 cm in diameter. The groundmass within most of the nodule horizons is reddish, weathered soil material. In addition to the nodule horizons, individual nodules are commonly scattered throughout the red clay. The red clay between the nodule horizons
shows a redder color than that of the paleosols in the overlying Pleistocene loess. The horizons have been subjected to relatively strong pedogenesis, as indicated by the relative abundance of clay and Fe–Mn skins, and are identified as pedological B horizons. A remarkable difference between the soil horizons of the red clay and loess records is that pedogenic A horizons are generally lacking in the red clay sequence, whereas a complete A–B–C
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sequence is readily recognized in the loess. A total of ca. 120 soil B horizons and ca. 110 intervening carbonate nodule horizons are identified in the Lingtai red clay sequence (Ding et al., 1999).
3. Sampling and analysis In the Pleistocene loess sequence, we took one carbonate nodule from each of the carbonate nodule horizons if nodules existed at the base of the paleosols; 27 nodule samples were collected in the loess–soil record. In the red clay sequence, one sample was taken from each thin nodule layer, while in every thick nodule horizon, two samples from the upper and lower parts were collected. A total of 178 samples were taken in the red clay record. Nodules were ground to a powder, and converted to CO with 100% phosphoric acid. Carbon 2 and oxygen isotopic ratios were both measured in a DeltaS gas-source mass spectrometer. Isotopic results are presented in the usual d notation as the permil (‰) deviation of the sample CO from the 2 PDB standard
A
R
B
sample −1 ×1000 R standard where R=13C/12C or 18O/16O. Replicate analyses (n=13) of a single, homogenized sample show that this procedure yields a standard deviation of ±0.05 and ±0.06‰, respectively, for d13C and d18O measurements.
d=
4. Results The carbon and oxygen isotopic records of soil carbonate nodules and magnetic susceptibility are shown in Fig. 3 for the Lingtai loess and red clay sequence. The time scale was established by linear interpolation between paleomagnetic reversal boundaries (Ding et al., 1998a). The carbon isotope record can be grouped roughly into three phases. From 7.05 to ca. 4.0 Myr, the d13C values of soil carbonate fall between −9.6 and −7.8‰, indicating a C -dominated ecosystem in the Lingtai 3 area. At ca. 4.0 Myr, the stable carbon isotopic
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compositions of soil carbonate begin to shift toward more positive values. Most of the d13C values lie between −7.8 and −5.3‰ from 4.0 to 2.0 Myr (Fig. 3). A single sample at the 3.4 Myr level yielded the d13C value of −4.7‰, representing the highest d13C value in the entire red clay sequence. Over this interval, a mixed ecosystem of C and C plants can be interpreted from the 3 4 stable carbon isotopic compositions of soil carbonate. Therefore, the expansion of C plants on the 4 Loess Plateau occurred at ca. 4.0 Myr, but at no time did C plants dominate the vegetation on the 4 plateau. From 2.0 Myr to the Holocene, most of the d13C values in soil carbonate range from −6.5 to −9.2‰ with an averaged value of ca. −8.0‰ ( Fig. 3). Compared to the interval of 4.0 to 2.0 Myr, C plants in the Loess Plateau ecosystem 3 increased significantly in the Pleistocene. However, in the youngest two paleosols (S1 and S2), soil carbonate is enriched in 13C (Fig. 3), suggesting another expansion of C plants on the plateau. 4 In the Pakistan stable isotopic records (Quade et al., 1989; Quade and Cerling, 1995), the shift of d13C compositions toward more positive values at ca. 7.0 Myr was accompanied by a clear increase in the d18O record. This feature is not found in our d18O record. Changes in soil carbonate d18O compositions range scatteringly from −11.8 to −8.0‰ (Fig. 3). There is an overall trend of d18O compositions toward more positive values from bottom to top of the loess and red clay sequence. Changes in d18O compositions of soil carbonate may be controlled by complicated factors, including temperature, regional moisture sources, rainfall amount, and seasonality of precipitation (Quade and Cerling, 1995). The causes for the trend of d18O changes of soil carbonate in the Loess Plateau are unclear. We preliminarily think that this may be partly related with the increase in seasonality during the past 7.0 Ma due to the long-term intensification of both the summer and winter monsoon systems over East Asia (Liu and Ding, 1993; Ding et al., 1999).
5. Discussion and conclusions Our carbon isotopic record indicates a significant expansion of C plants at ca. 4.0 Myr on the 4
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Fig. 3. Magnetic susceptibility and carbon and oxygen isotope records plotted against time. The timescale is constructed by linear interpolation between magnetic reversal boundaries.
Chinese Loess Plateau. A major question concerning this event is whether or not it was driven by local climatic changes. The pedogenic characteristics of the soils within the red clay sequence have recorded the history of the East-Asia summer monsoon evolution during the late Miocene and Pliocene in the Loess Plateau. Recent observations (Ding et al., 1999) show that the Lingtai red clay sequence can be subdivided into five units according to soil features. The first ( lowest) unit accumulated from 7.05 to 6.2 Myr. Soils in this unit are characterized by a reddish brown color and a weak subangular blocky structure with few clay skins, suggesting a relatively weak summer monsoon. The second unit, formed during 6.2 to 5.5 Myr, contains soils showing a moderate subangular blocky structure and some or common clay skins, implying a strengthened summer monsoon relative
to the earlier time. The soils of the third unit (5.5 to 3.85 Myr) display many thick clay skins and many dark Fe–Mn films, interpreted as having developed under strong summer monsoon conditions. The summer monsoon was significantly weakened over the interval of 3.85 to 3.15 Ma (the fourth unit), as suggested by greatly reduced clay skins on the pedogenic structure surfaces. In the fifth unit (3.15 to 2.6 Myr), translocated clay skins in the soils again significantly increase, implying an intensified summer monsoon climate. This interpreted monsoon evolution suggests that there is no significant local climatic event at ca. 4.0 Myr, and thus it seems unlikely that the ecological change registered in the d13C record was be forced solely by changes in local climatic conditions. This idea may be supported also by the soil carbonate d18O record ( Fig. 3) which shows an almost linear
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enrichment in 18O from bottom to top of the sequence. No peculiar change ca. 4.0 Myr is found in the record. Therefore, if the d18O trend indeed represents the increase in seasonality of precipitation, this factor is not the most important in driving the middle Pliocene C plant expansion on 4 the Loess Plateau. The modern spatial distribution of C and C 4 3 grasses shows that C grasses dominate in the 4 tropical and subtropical regions, that the transition to C grasses takes place between ca. 30 and 45° 3 latitudes, and that C grasses dominate at high 3 latitudes. According to the model developed by Cerling et al. (1997), the photosynthetic efficiency of C grasses relative to C grasses varies with 3 4
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both atmospheric CO levels and temperature. 2 This model predicts that with the continuous decrease in atmospheric CO levels during the late 2 Cenozoic, as deduced from geological records, expansion of C plants will occur first in the 4 tropical and subtropical regions, and later in the transitional latitudes because of lower temperatures. Fig. 4 shows the comparison of d13C records between Pakistan, the Chinese Loess Plateau and northern North America (>37°N ). It is evident that the Loess Plateau and northern North America experienced a synchronous expansion of C plants 4.0 Myr ago, ca. 3 million years 4 later than the C expansion in Pakistan. This may 4 support the model of Cerling et al. (1997),
Fig. 4. Comparison of the C -plant expansion event between the Loess Plateau (pedogenic carbonate d13C ), Pakistan (pedogenic 4 carbonate d13C ) and northern North America ( Equid d13C ). The arrows indicate the positions of the event. The Pakistan data are from Quade et al. (1989) and the northern North America data from Cerling et al. (1997).
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although direct evidence for high concentration of atmospheric CO in the Neogene is lacking (Pagani 2 et al., 1999; Pearson and Palmer, 1999). The Pakistan carbon isotopic records are derived from between 32 and 33°N, only 2–3° of latitude lower than the Lingtai section (35°N ). The difference in the timing of the C -plant expan4 sion between the two regions is, however, as great as ca. 3 million years. On the other hand, northern North America has substantially higher latitudes than Lingtai, but the two regions witnessed an approximately synchronous C -plant expansion. 4 This evidence suggests an anomalously low temperature of the Loess Plateau ca. 7.0 Myr ago, according to the Cerling et al. (1997) model. At present, cold air masses propagated onto the Plateau are mainly delivered by the northerly winter monsoon and the westerlies. Changes in grain size values of four red clay–loess sections demonstrate that the southward gradient of decreasing grain size in the Tertiary red clay on the Plateau is negligible (Ding et al., 1998b), whereas there is a strong negative gradient in particle size from north to south in the Loess Plateau for sediments younger than 2.6 Ma. This decrease in loess particle size is consistent with the direction of the winter monsoonal winds flowing from Siberia. It is suggested that the winter monsoon may not have formed until ca. 2.6 Myr ago (Ding et al., 1998b). In this case, the lowered temperature background over the Loess Plateau during the late Miocene may be caused by increased propagation of cold air masses by the westerlies. Meteorological observations ( Wallace, 1983) show that at 500 mb, there is a broad trough, called the Great East-Asia Trough (GEAT ), centered about the longitudes of Japan and forming a broad ridge to the west, which produces northwesterly winds in the upper troposphere over China. The GEAT is the strongest of the quasistationary waves on the globe, which is partly responsible for the relatively low temperature over eastern and central China. GCM modeling experiments (Manabe and Terpstra, 1974; Kutzbach et al., 1989) show that this stationary wave over Asia is caused by the uplifted Tibetan Plateau. The greatly-delayed shift of C plants in the Loess 4 Plateau suggests that the Tibetan Plateau may have reached a critical height in the late Miocene,
thus leading to the formation of the GEAT. This is consistent with the observations of several geological studies (Mercier et al., 1987; Copeland and Harrison, 1990; Harrison et al., 1993). Our d13C record shows a subsequent diminishment of the C -plant component in the period of 4 2 Myr to the late Pleistocene ( Fig. 3). This phenomenon is also observed in the dental enamel carbon isotope record of fossil herbivores from northern North America (Cerling et al., 1997). To date, the cause for this is unknown. As no evidence is obtained for a subsequent increase in atmospheric CO levels during the Pleistocene, we pro2 pose that decrease in temperature could be the main factor for this event. Dramatic expansion of polar ice sheets in the Northern Hemisphere during late Pliocene and early Pleistocene has been indicated by many geological records (Ruddiman et al., 1989; Raymo et al., 1989; Shackleton et al., 1990). Large-scale ice sheet expansion in the polar area will lead to a decrease in temperature background and an increase in north–south temperature gradient, thereby causing a southward shift of C taxa in the Northern Hemisphere. 3 Acknowledgements This research is supported by the NNSF of China (49525203) and CAS ( KZ951-A1-402). The authors are greatly indebted to Dr Eve Arnold for critical comments on the manuscript.
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