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Keywords:

  • East Asian Monsoon;
  • northern China;
  • climate history;
  • lake sediment;
  • dolomite;
  • calcite

Abstract

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Material and Methods
  5. 3. Results and Discussion
  6. 4. Conclusions
  7. Acknowledgments
  8. References
  9. Supporting Information

[1] To determine the timing and spatial distribution of mid-Holocene drying over northern China, the mineralogical and oxygen isotopic composition of authigenic carbonate from a closed lake at Bayanchagan, southern Inner Mongolia, were measured. Further analysis and synthesis of the spatial geological data were performed. Results from Lake Bayanchagan show a significant drying at 6000 calendar years (cal years) B.P., indicated by dolomite precipitation and a striking rise in δ18O values. The synthesis of spatial data reveals a zonal distribution for timing of drying over northern China in the mid-Holocene, which began at 9000–7000 cal years B.P. in deserts of north-central China. At 7000–5500 cal years B.P., drying extended into the desert-steppe transitional zone and at ∼4500 cal years B.P. into northeastern and south-central China. This pattern indicates that the east Asian summer monsoon has significantly retreated southeastward since the mid-Holocene, which may be related to orbitally induced Northern Hemisphere insolation changes. A retreat of ∼400–550 km is inferred for the front of the summer monsoon from 6500 to 4500 cal years B.P.

1. Introduction

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Material and Methods
  5. 3. Results and Discussion
  6. 4. Conclusions
  7. Acknowledgments
  8. References
  9. Supporting Information

[2] The climate of China is mainly controlled by the East Asian Monsoon, which comprises two seasonally alternating atmospheric circulations. In winter, a dry-cold air mass from Siberia leads to a cold and dry climate, while in summer, a southeast monsoon transports heat and moisture inland from the low-latitude oceans, with a gradient of decreasing rainfall going from the southeast to the northwest (Figure 1). The steep inland precipitation gradient is observed in the northern transitional zone between the desert and steppe landscapes, which are particularly sensitive to monsoon precipitation changes.

image

Figure 1. Map showing the location of Lake Bayanchagan, deserts and distribution of mean annual precipitation (mm) in China. Precipitation data are from the National Climate Centre of China Meteorological Administration. Deserts are marked as follows: A, Taklimakan; B, Gurbantunggut (Junggar); C, Kumtag; D, Qaidam; E, Badain Jaran; F, Tengger; G, Mu Us; H, Hobq; I, Ulan Buh; J, Otindag; K, Horqin; and L, Hulun Buir. The arrows indicate the advance of the east Asian summer monsoonal rainfall belt.

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[3] Previous studies have shown that a dry climate followed early Holocene humid conditions in northern China [e.g., An et al., 2000; Zhou et al., 2001, 2002; C.-T. A.Chen et al., 2003; Jiang et al., 2006]. However, the initiation of mid-Holocene drying varies between sites in northern China, starting between 9000 and 4000 years ago [e.g., An et al., 1993; Chen et al., 1999; Zhou et al., 2001, 2002; Liu et al., 2002; Shi et al., 2002; C.-T. A.Chen et al., 2003; F.-H.Chen et al., 2003; An et al., 2003; Li et al., 2003; He et al., 2004; Jiang et al., 2006]. Clearly, the spatial distribution of the mid-Holocene drying event needs to be identified, not only because this information is crucial to an understanding of climate mechanisms and as a means of defining model boundary conditions, but also because dry events in northern China played a significant role in the collapse and substitution of Chinese Neolithic cultures [Wu and Liu, 2004]. Solutions to this problem largely depend on more precisely dated, high-resolution geological records in climatically sensitive regions of northern China, and further analysis and synthesis of the published data.

[4] In this study, we first present mineralogical and δ18O records of authigenic carbonate from a basin-closed lake at Bayanchagan (BY), southern Inner Mongolia. This site is situated at the current northern edge of the summer monsoon [Gao, 1962], in one of the key regions for reconstructing East Asian Monsoon history. Precise age-constrained, high-resolution paleoclimatic records throughout northern China have been collected and compiled to identify the spatial pattern of mid-Holocene drying and its relationship with the East Asian Monsoon.

2. Material and Methods

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Material and Methods
  5. 3. Results and Discussion
  6. 4. Conclusions
  7. Acknowledgments
  8. References
  9. Supporting Information

[5] Lake Bayanchagan (115.21°E, 41.65°N; 1355 m asl) was a ∼15 km2 closed-basin lake in 1959. Today, it is almost completely dry because of anthropogenic water use, with only small patches of shallow water maintained by summer rain. A 1.8 m trench was cut (core BY, Figure 2a) into the center of Lake Bayanchagan sediment. The chronology of the core is constrained by seven radiocarbon dates of bulk organic carbon (Figure 2b) [Jiang et al., 2006]. Assuming a relatively constant proportion of aquatic to nonaquatic organic carbon throughout the core, the reservoir effect is ∼570 years for Lake Bayanchagan [Jiang et al., 2006].

image

Figure 2. (a) Lithology, (b) chronology, (c) carbonate contents, (d) δ18O values, (e) PFT scores of deciduous trees, (f) pollen concentrations, and (g) estimated mean annual precipitation at core BY. Figures 2a, 2b and 2e–2g are from Jiang et al. [2006].

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[6] Carbonate content and δ18O values of carbonate were determined at 2 cm intervals. Prior to CO2 extraction, all samples were sieved and the <40 μm fraction collected for further analysis, as the carbonate in the <40 μm fraction of lake sediment was considered to be authigenic origin [Fontes et al., 1996]. δ18O values were determined using a Finnigan MAT252 mass spectrometer and are reported in per mil units (‰) relative to the Vienna Peedee belemnite (VPDB) standard. Replicate analyses (n = 5) show that this procedure yields a precision of better than ±0.2‰. X-ray diffraction (XRD, RINT2000 Wide-angle goniometer), and scanning electron microscopy (SEM, LEO 1450VP) were performed to identify the composition and morphology of the authigenic carbonate.

3. Results and Discussion

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Material and Methods
  5. 3. Results and Discussion
  6. 4. Conclusions
  7. Acknowledgments
  8. References
  9. Supporting Information

3.1. Mineralogical and Oxygen Isotopic Composition of the Authigenic Carbonate From BY Core

[7] Carbonate content is abundant in the BY core (Figure 2c), increasing from 8–40% before 10,800 calendar years (cal years) B.P. to 40–55% between 10,800 and 7200 cal years B.P., and subsequently decreasing gradually from 55% to 34% from 7200 cal years B.P. to the present. XRD results show that calcite predominates in the carbonate fraction between 12,400 and 6000 cal years B.P. (depth 180–48 cm), and dolomite occurs after 6000 cal years B.P. (depth 47 cm) (Figure 3). Using SEM, the calcite fraction consists of small (1–5 μm long), flaky and lenticular idiomorphic crystals. Dolomite occurs as a knobbly coating (<1 μm) on the surface of feldspar (Figure 4). The morphological features of calcite and dolomite indicate rapid carbonate precipitation [Fontes et al., 1996].

image

Figure 3. X-ray diffractogram (28° to 32° 2θ) for the <40 μm fraction from the core BY. Below depth of 48 cm, the major peak occurs at 3.01Å (29.62° 2θ), reflecting dominance of calcite. The flat peak at depth of 137 cm represents a relatively high content of magnesium in calcite. At depth of 47 cm, the small increase in the intensity at 2.89Å (30.96° 2θ) indicates the presence of dolomite. Above a depth of 43 cm, the major peak occurs at 2.89Å (30.96° 2θ), indicating an increase of dolomite [Vasconcelos et al., 1995].

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image

Figure 4. SEM photographs of <40 μm fraction from BY core showing (a) an overview of calcite aggregates on the surface of feldspar, (b) an enlarged picture of the rectangle in Figure 4a, (c) two calcite aggregates, (d) calcite crystals precipitated on the surface of quartz as marked by circles, (e) knobbly dolomite coating on the surface of feldspar, and (f) an enlarged picture of the rectangle in Figure 4e. The dolomite marked by circles has morphological characteristics similar to those of dolomite precipitated in experiments [Vasconcelos et al., 1995].

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[8] From 12,400 to 10,800 cal years B.P., all δ18O values of authigenic carbonate are between −3 and −1‰ VPDB (Figure 2d). From 10,800 to 7800 cal years B.P., the oxygen isotope value decreases with lowest δ18O values (averaging −7.0‰) observed between 7800 and 7200 cal years B.P. From 7200 to 6000 cal years B.P., the oxygen isotope value gradually increases to between −6.2 and −5.3‰. This trend is followed by a rapid increase of 4.1‰ in δ18O values from 6000 to 4800 cal years B.P. After 4800 cal years B.P., the δ18O values show small variations between −1.5 and −0.2‰.

3.2. Mid-Holocene Drying Inferred From Dolomite Precipitation and δ18O Record in BY Core

[9] In lake environments, carbonate phases are controlled by salinities and Mg/Ca ratios of lake water [Müller et al., 1972]. As salinity increases, low-Mg calcite precipitates first, followed by high-Mg calcite, aragonite, and finally dolomite. In core BY, the presence of dolomite around 6000 cal years B.P. provides depositional evidence for salinization, indicative of increasing aridity since the mid-Holocene in southern Inner Mongolia.

[10] Increased aridity beginning in the mid-Holocene is supported by the δ18O record of the authigenic carbonate. δ18O values of authigenic carbonate are controlled by temperature and the oxygen isotope value of lake water [e.g., Fontes et al., 1996], and carbonate phases [Land, 1980; McKenzie, 1981]. Previous studies have shown that dolomite is enriched in 18O relative to cogenetic calcite by ∼3.0‰ [Land, 1980; McKenzie, 1981]. Assuming a maximum percentage of 100% for dolomite after 4800 cal years B.P., then a minimum variation of ∼3.0‰ in δ18O values is expected for cogenetic calcite during the time interval 7200–4800 cal years B.P. (Figure 2d). If temperature were the major factor controlling carbonate δ18O values, then the minimum variation of ∼3.0‰ would have resulted from at least a 14°C temperature change [Kim and O'Neil, 1997]. Because such a large change in temperature is unlikely, we suggest that the δ18O values of authigenic carbonate are mainly controlled by the oxygen isotope value of lake water, i.e., the balance between precipitation and evaporation [Leng and Marshall, 2004]. Increased monsoonal precipitation produces lower water δ18O values while decreased precipitation generates higher δ18O values, as demonstrated in previous studies [e.g., Wei and Gasse, 1999]. Evaporation of lake water results in higher δ18O values in authigenic carbonate [Leng and Marshall, 2004]. Our δ18O values suggest that precipitation gradually decreased from 7200 to 6000 cal years B.P., and then decreased more rapidly from 6000 to 4800 cal years B.P., indicating progressive drying between 7200 and 4800 cal years B.P. Here we consider the presence of dolomite at 6000 cal years B.P. as marking the beginning of the mid-Holocene drying event at core BY. The mineralogical and oxygen isotopic changes are consistent with pollen data and mean annual precipitation estimated from the same core (Figures 2e–2g) [Jiang et al., 2006].

3.3. Timing and Spatial Distribution of Mid-Holocene Drying Over Northern China and Its Implications for the East Asian Monsoon Changes

[11] Paleoclimatic records that range from Tengger Desert to Bohai Bay (Figure 1) over northern China (Table 1, selection criteria are the same as An et al. [2006]) have been collected and compiled from lacustrine deposits, peat, loess and stalagmites. Most sites are located in the transitional area between desert and steppe, a climatically sensitive zone, approximately paralleling the 400 mm isohyet of mean annual precipitation.

Table 1. Paleoclimatic Records Collected in This Studya
Site NumberSiteRecordProxiesReference
  • a

    CE, chemical elements; TOC, total organic carbon; LOI, loss on ignition; MS, magnetic susceptibility.

1Hongshui Riverfluvial-lacustrinepollen, TOC, δ18O, δ13C, CaCO3, CEZhang et al. [2000]
2Yema Lakelakegrain size, carbonateShi et al. [2002]
3Alashan PlateaulakespollenF.-H.Chen et al. [2003]
4Yanhaizi Lakelakegrain size, MS, TOC, C/NC.-T. A.Chen et al. [2003]
5Midiwanloess-soil-sandpollen, δ13CZhou et al. [2001], Li et al. [2003]
6Loess Plateauloess-soilpollen, δ13C, TOCZhou et al. [2001]
7Dadiwanloess-wetlandpollen, mollusk, TOC, grain sizeAn et al. [2003]
8Sujiawanloess-wetlandpollen, mollusk, TOC, grain sizeAn et al. [2003]
9Baxieloess-soilMS, grain size, TOC, δ13CAn et al. [1993]
10ZoigepeatTOC, gray scaleZhou et al. [2002]
11Shanbao Cavestalagmiteδ18OShao et al. [2006]
12MianyanglakepollenYang et al. [1998]
13QidongYangtze deltapollen, TOC, CaCO3Liu et al. [1992]
14Huanghe deltaHuanghe deltapollenYi et al. [2003]
15Taishizhuangpeatpollen, δ18OJin and Liu [2002], Tarasov et al. [2006]
16Bayanchagan Lakelakepollen, Pediastrum, carbonate mineral, δ18OJiang et al. [2006], this study
17Haolukuancient lakepollen, grain size, LOI, CELiu et al. [2002]
18Xiaoniuchangancient lakepollen, grain size, LOI, CELiu et al. [2002]
19GushantunpeatpollenLiu [1989], Sun et al. [1991]
20JinchuanpeatpollenSun et al. [1991], W. Y. Jiang and T. S. Liu (unpublished data, 2006)
21SandaolaoyefupeatpollenYuan and Sun [1990], Sun et al. [1991]
22Hulun Lakelakepollen, diatom, OstracodeYang et al. [1995]

[12] We identified mid-Holocene dry events through vegetation changes, the termination of soil development and the decline in lake levels [e.g., An et al., 1993; Zhou et al., 2001; Liu et al., 2002; Shi et al., 2002; F.-H.Chen et al., 2003], as these changes are largely related to precipitation [An et al., 2000]. For vegetation records from arid and semiarid regions, the date of increase in desert vegetation (e.g., Ephedra and Chenopodiaceae) was taken as the time of drying [e.g., Zhang et al., 2000], while the date of decrease in broadleaved deciduous/evergreen trees and/or increase in steppe vegetation was selected for those from semihumid and humid regions [e.g., Liu et al., 1992; Yang et al., 1998; Liu et al., 2002; An et al., 2003; Tarasov et al., 2006].

[13] Humidity reconstruction in deserts was obtained mainly from pollen and geochemical records. A fluvial-lacustrine record from the terrace of the Hongshui River in the Tengger Desert (site 1, Figure 5) [Zhang et al., 2000] suggests a significant increase in desert vegetation (Ephedra and Chenopodiaceae) at 7500 cal years B.P., reflecting enhanced aridity. Lake records indicate a severe drought starting between 7200 and 7000 cal years B.P. in the Tengger Desert (sites 2 and 3, Figure 5) [Shi et al., 2002; F.-H.Chen et al., 2003], and 8800 cal years B.P. in the Hobq Desert (site 4, Figure 5) [C.-T. A.Chen et al., 2003]. A drought that started at 8300 cal years B.P. is also indicated by a pollen record near the margin of Mu Us Desert (site 5, Figure 5) [Li et al., 2003].

image

Figure 5. Map showing the timing (solid lines, cal years B.P.) and spatial distribution of mid-Holocene drying over northern China. Sites are numbered according to Table 1. Dashed lines are isohyets of mean annual precipitation (mm). Deserts are marked as follows: A, Taklimakan; B, Gurbantunggut (Junggar); C, Kumtag; D, Qaidam; E, Badain Jaran; F, Tengger; G, Mu Us; H, Hobq; I, Ulan Buh; J, Otindag; K, Horqin; and L, Hulun Buir. The arrows indicate the retreat of the summer monsoonal rainfall belt.

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[14] In the desert-steppe transitional zone, paleoclimatic records mainly come from loess and lake sediments. Loess deposits have long been regarded as one of the most important archives of climatic evolution in China during the past 2.6 Ma [Liu and Ding, 1998]. Loess beds were deposited when climate was arid and cold, whereas soils developed under relatively warm and humid conditions [Liu, 1985; Kukla, 1987; Liu and Ding, 1998]. Previous studies have shown that development of Holocene soil started at 11,500–10,000 cal years B.P., and terminated at 6800–5700 cal years B.P. (sites 6 and 9, Figure 5) [An et al., 1993; Zhou et al., 2001]. Pollen records from lake sediment demonstrate that vegetation type and cover deteriorated at the same time as soil termination, characterized by decreases in pollen concentrations (sites 16 and 22, Figure 5) [Yang et al., 1995; Jiang et al., 2006] and arboreal pollen percentages (sites 16–18, Figure 5) [Liu et al., 2002; Jiang et al., 2006]. The good agreement in timing between soil termination and vegetation changes suggests a mid-Holocene drying at 7000–5500 cal years B.P. in the desert-steppe transitional zone.

[15] There are few paleoclimatic records, with the exception of pollen and stalagmites, available for northeastern and south-central China. Broadleaved deciduous forest was replaced by a mixed forest of deciduous and coniferous trees at 4400–4200 cal years B.P. in northeastern China (sites 19–21, Figure 5) [Liu, 1989; Yuan and Sun, 1990; Sun et al., 1991; W. Y. Jiang et al., unpublished data, 2006]. An increase in steppe vegetation occurred at 4800 cal years B.P. at the Taishizhuang peat bog (site 15, Figure 5) [Jin and Liu, 2002; Tarasov et al., 2006] and 4500 cal years B.P. at the delta of the Yellow River (site 14, Figure 5) [Yi et al., 2003]. Pollen records from south-central China have shown that a decrease in subtropical broadleaved evergreen trees, and an increase in trees with greater moisture tolerances, such as the deciduous genera Quercus, Corylus and the coniferous genus Pinus, occurred at 4300–4200 cal years B.P. (sites 12 and 13, Figure 5) [Liu et al., 1992; Yang et al., 1998]. A high-resolution oxygen-isotope record of a precisely dated stalagmite from Shanbao Cave (site 11, Figure 5) reveals an abrupt decrease in monsoonal precipitation at 4400 cal years B.P. [Shao et al., 2006]. These records suggest a dry, cold climate after 4800–4200 cal years B.P.

[16] On the basis of the above geological records, together with the BY record, we compiled a contour map showing the spatial changes in timing of mid-Holocene drying over northern China using Surfer software. It shows a clear zonal distribution from northwest to southeast over northern China (Figure 5). Drying began at 9000–7000 cal years B.P. in the deserts of north-central China, extending into the transitional zone between desert and steppe at 7000–5500 cal years B.P. and at ∼4500 cal years B.P. into northeastern and south-central China. This zonal distribution of drying time in the mid-Holocene roughly parallels the isohyets of mean annual precipitation in northern China (Figure 5), indicating a close relationship with the east Asian summer monsoon.

[17] The summer monsoonal precipitation is produced by the interaction between warm moist southerly air masses and cold northerly airflows from middle and high latitudes. In general, the more northerly the penetration of the rainfall belt into interior continental northern China, the greater is the intensity of the summer monsoon [Tao and Chen, 1987]. The Pleistocene witnessed numerous retreat-advance cycles of the east Asian summer monsoon in response to glacial-interglacial cycles [Liu and Ding, 1998; Ding et al., 2001, 2002]. For any specific site in northern China, it would usually become drier when the summer monsoon retreated southward. The progressive drying from 7200 to 4800 cal years B.P. recorded in Lake BY is consistent with the progressive weakening of the east Asian summer monsoon. In this context, the zonal distribution of the desiccation suggests that the East Asian Monsoon has significantly retreated southeastward since the mid-Holocene. Furthermore, a ∼400–550 km retreat is inferred for the summer monsoon from 6500 to 4500 cal years B.P.

[18] Changes in the intensity of the east Asian summer monsoon during the Holocene may be related to orbitally induced Northern Hemisphere insolation variability [An et al., 2000; Kutzbach et al., 2001; Wang et al., 2005]. According to Berger and Loutre [1991], Northern Hemisphere summer solar insolation has been decreasing since the early Holocene. Numerical modeling experiments have shown that a decrease in insolation would have resulted in a weakened monsoon circulation [Ganopolski et al., 1998; Kutzbach et al., 2001]. The forcing mechanism lies in that the cooling of the continental interior gives rise to a decrease in thermal contrast between the east Asian continent and the adjacent ocean, thus leading to a weakened monsoon circulation, i.e., a southeastward retreat of the east Asian summer monsoon.

4. Conclusions

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Material and Methods
  5. 3. Results and Discussion
  6. 4. Conclusions
  7. Acknowledgments
  8. References
  9. Supporting Information

[19] In Lake Bayanchagan, precipitation of dolomite started at around 6000 cal years B.P. δ18O values of authigenic carbonate increased gradually from 7200 to 6000 cal years B.P., and then rapidly from 6000 to 4800 cal years B.P. All these data indicate the initiation of significant drying at 6000 cal years B.P. in southern Inner Mongolia.

[20] Synthesis of the spatial data suggests a zonal pattern for the timing of mid-Holocene drying over northern China. The drying began at 9000–7000 cal years B.P. in the deserts of north-central China, then extended into the desert-steppe transitional zone at 7000–5500 cal years B.P., and at ∼4500 cal years B.P. into northeastern and south-central China. This pattern suggests that the east Asian summer monsoon has retreated significantly southeastward since the mid-Holocene, which may be related to orbitally forced Northern Hemisphere insolation variation. A retreat of ∼400–550 km is inferred for the front of the summer monsoon from 6500 to 4500 cal years B.P. This information will provide valuable records for validating GCM models for ancient climate in this region.

Acknowledgments

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Material and Methods
  5. 3. Results and Discussion
  6. 4. Conclusions
  7. Acknowledgments
  8. References
  9. Supporting Information

[21] This work was supported by National Basic Research Program of China (grant 2004CB720203), the National Natural Science Foundation of China (grant 40273037), Chinese Academy of Sciences (grant KZCX2-YW-117), and the cooperative project between CAS and CNRS. We are greatly indebted to V. A. Hall, J. P. Smol and W. P. Patterson for critical comments on an earlier version of this manuscript. We also thank A. Alexandre, G. Q. Chu, S. Gachet, J. Guiot, Z. T. Guo, H. B. Wu and B. Y. Yuan for field work; Z. Y. Zhang for lab work; and Z. T. Guo, Z. Y. Gu, J. Guiot, C. Hatté, J. M. Sun and S. L. Yang for valuable suggestions.

References

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Material and Methods
  5. 3. Results and Discussion
  6. 4. Conclusions
  7. Acknowledgments
  8. References
  9. Supporting Information
  • An, C. B., Z.-D. Feng, and L. Y. Tang (2003), Evidence of humid mid-Holocene in the western part of the Chinese Loess Plateau, Chin. Sci. Bull., 48, 24722479.
  • An, C. B., Z.-D. Feng, and L. Barton (2006), Dry or humid? Mid-Holocene humidity changes in arid and semi-arid China, Quat. Sci. Rev., 25, 351361.
  • An, Z. S., S. C. Porter, W. J. Zhou, Y. C. Lu, D. J. Donahue, M. J. Head, X. H. Wu, J. Z. Ren, and H. B. Zheng (1993), Episode of strengthened summer monsoon climate of Younger Dryas age on the Loess Plateau of central China, Quat. Res., 39, 4554.
  • An, Z. S., S. C. Porter, J. E. Kutzbach, X. H. Wu, S. M. Wang, X. D. Liu, X. Q. Li, and W. J. Zhou (2000), Asynchronous Holocene optimum of the East Asian Monsoon, Quat. Sci. Rev., 19, 743762.
  • Berger, A., and M. F. Loutre (1991), Insolation values for the last 10 million years, Quat. Sci. Rev., 10, 297317.
  • Chen, C.-T. A., H.-C. Lan, J.-Y. Lou, and Y.-C. Chen (2003), The dry Holocene Megathermal in Inner Mongolia, Palaeogeogr. Palaeoclimatol. Palaeoecol., 193, 181200.
  • Chen, F.-H., Q. Shi, and J.-M. Wang (1999), Environmental changes documented by sedimentation of Lake Yiema in arid China since the Late Glaciation, J. Paleolimnol., 22, 159169.
  • Chen, F.-H., W. Wu, J. A. Holmes, D. B. Madsen, Y. Zhu, M. Jin, and C. G. Oviatt (2003), A mid-Holocene drought interval as evidenced by lake desiccation in the Alashan Plateau, Inner Mongolia, China, Chin. Sci. Bull., 48, 14011410.
  • Ding, Z. L., Z. W. Yu, S. L. Yang, J. M. Sun, S. F. Xiong, and T. S. Liu (2001), Coeval changes in grain size and sedimentation rate of eolian loess, the Chinese Loess Plateau, Geophys. Res. Lett., 28, 20972100.
  • Ding, Z. L., E. Derbyshire, S. L. Yang, Z. W. Yu, S. F. Xiong, and T. S. Liu (2002), Stacked 2.6-Ma grain size record from the Chinese loess based on five sections and correlation with the deep-sea δ18O record, Paleoceanography, 17(3), 1033, doi:10.1029/2001PA000725.
  • Fontes, J. C., F. Gasse, and E. Gibert (1996), Holocene environmental changes in Lake Bangong basin (western Tibet) part 1: Chronology and stable isotopes of carbonates of a Holocene lacustrine core, Palaeogeogr. Palaeoclimatol. Palaeoecol., 120, 2547.
  • Ganopolski, A., C. Kubatzki, M. Claussen, V. Brovkin, and V. Petoukhov (1998), The influence of vegetation-atmosphere-ocean interaction on climate during the mid-Holocene, Science, 280, 19161919.
  • Gao, Y. X. (1962), On some problems of Asian monsoon (in Chinese), in Some Questions About the East Asian Monsoon, edited by Y. X. Gao, pp. 149, China Sci. Press, Beijing.
  • He, Y., W. H. Theakstone, Z. L. Zhang, D. Zhang, T. D. Yao, T. Chen, Y. P. Shen, and H. X. Pang (2004), Asynchronous Holocene climatic change across China, Quat. Res., 61, 5263.
  • Jiang, W. Y., Z. T. Guo, X. J. Sun, H. B. Wu, G. Q. Chu, B. Y. Yuan, C. Hatté, and J. Guiot (2006), Reconstruction of climate and vegetation changes of Lake Bayanchagan (Inner Mongolia): Holocene variability of the East Asian Monsoon, Quat. Res., 65, 411420.
  • Jin, G. Y., and T. Liu (2002), Mid-Holocene climate change in north China, and the effect on cultural development, Chin. Sci. Bull., 47, 408413.
  • Kim, S. T., and J. R. O'Neil (1997), Equilibrium and nonequilibrium oxygen isotope effects in synthetic carbonates, Geochim. Cosmochim. Acta, 61, 34613475.
  • Kukla, G. (1987), Loess stratigraphy in central China, Quat. Sci. Rev., 6, 191219.
  • Kutzbach, J. E., S. P. Harrison, and M. T. Coe (2001), Land-ocean-atmosphere interactions and monsoon climate change: A paleo-perspective, in Global Biogeochemical Cycles in the Climate System, edited by E.-O. Schulze et al., pp. 7386, Academic, San Diego, Calif.
  • Land, L. S. (1980), The isotopic and trace element geochemistry of dolomite: The state of art, in Concepts and Models of Dolomitization, edited by D. H. Zenger et al., Spec. Publ. Soc. Econ. Paleontol. Mineral., 28, 87110.
  • Leng, M. J., and J. D. Marshall (2004), Palaeoclimate interpretation of stable isotope data from lake sediment archives, Quat. Sci. Rev., 23, 811831.
  • Li, X. Q., W. J. Zhou, Z. S. An, and J. Dodson (2003), The vegetation and monsoon variations at the desert-loess transition belt at Midiwan in northern China for the last 13 ka, Holocene, 13(5), 779784.
  • Liu, H. Y., L. H. Xu, and H. T. Cui (2002), Holocene history of desertification along the woodland-steppe border in northern China, Quat. Res., 57, 259270.
  • Liu, J. L. (1989), Vegetational and climatic changes at Gushantun Bog in Jilin, NE China since 13000 yr (in Chinese with English abstract), Acta Palaeontol. Sin., 28, 496511.
  • Liu, K.-B., S. C. Sun, and X. H. Jiang (1992), Environmental change in the Yangtze River Delta since 12000 years B.P. Quat. Res., 38, 3245.
  • Liu, T. S. (1985), Loess and the Environment, China Ocean Press, Beijing.
  • Liu, T. S., and Z. L. Ding (1998), Chinese loess and the paleomonsoon, Annu. Rev. Earth Planet. Sci., 26, 111145.
  • McKenzie, J. A. (1981), Holocene dolomitization of calcium carbonate sediments from the coastal sabkhas of Abu Dhabi, U.A.E.: A stable isotope study, J. Geol., 89, 185198.
  • Müller, G., G. Irion, and U. Forstner (1972), Formation and diagenesis of inorganic Ca-Mg carbonates in the lacustrine environment, Naturwissenschaften, 59, 158164.
  • Shao, X. H., Y. J. Wang, H. Cheng, X. G. Kong, J. Y. Wu, and E. R. Lawrence (2006), Long-term trend and abrupt events of the Holocene Asian monsoon inferred from a stalagmite δ18O record from Shennongjia in central China, Chin. Sci. Bull., 51, 221228.
  • Shi, Q., F.-H. Chen, Y. Zhu, and D. Madsen (2002), Lake evolution of the terminal area of shiyang River drainage in arid China since the last glaciation, Quat. Int., 93–94, 3143.
  • Sun, X. J., et al. (1991), The vegetation history of mixed Korean pine and deciduous forests in Changbai Mt. area, Jilin Province, northeast China during the last 13000 years, Chin. J. Bot., 3, 4761.
  • Tao, S. Y., and L. X. Chen (1987), A review of recent research on the east Asian summer monsoon, in Monsoon Meteorology, edited by C. P. Chang, and T. N. Krishnamurti, pp. 6092, Oxford Univ. Press, Oxford, U. K.
  • Tarasov, P., G. Y. Jin, and M. Wanger (2006), Mid-Holocene environmental and human dynamics in northeastern China reconstructed from pollen and archaeological data, Palaeogeogr. Palaeoclimatol. Palaeoecol., 241, 284300.
  • Vasconcelos, C., J. A. McKenzie, S. Bernasconi, D. Grujic, and A. J. Tiens (1995), Microbial mediation as a possible mechanism for natural dolomite formation at low temperatures, Nature, 377, 220222.
  • Wang, Y. J., et al. (2005), The Holocene Asian monsoon: Links to solar changes and North Atlantic climate, Science, 308, 854857.
  • Wei, K., and F. Gasse (1999), Oxygen isotopes in lacustrine carbonates of west China revisited: Implications for post glacial changes in summer monsoon circulation, Quat. Sci. Rev., 18, 13151334.
  • Wu, W. X., and T. S. Liu (2004), Possible role of the “Holocene Event 3” on the collapse of Neolithic cultures around the central Plain of China, Quat. Int., 117, 153166.
  • Yang, X. D., S. M. Wang, and B. Xue (1995), Vegetational development and environmental changes in Hulun Lake since late Pleistocene (in Chinese with English abstract), Acta Palaeontol. Sin., 34, 647656.
  • Yang, X. D., Y. X. Zhu, X. Z. Jiang, Y. H. Wu, and S. M. Wang (1998), Environmental changes from spore-pollen record of Mianyang region over the past 10000 years (in Chinese with English abstract), J. Lake Sci., 10, 2329.
  • Yi, S., Y. Saito, H. Oshima, Y. Q. Zhou, and H. L. Wei (2003), Holocene environmental history inferred from pollen assemblages in the Huanghe (Yellow River) delta, China: Climatic change and human impact, Quat. Sci. Rev., 22, 609628.
  • Yuan, S. M., and X. J. Sun (1990), The vegetational and environmental history at the west foot of Changbai Mountain, northeast China during the last 10000 years (in Chinese with English abstract), Acta Bot. Sin., 32(7), 558567.
  • Zhang, H. C., Y. Z. Ma, B. Wünnemann, and H.-J. Pachur (2000), A Holocene climate record from arid northwestern China, Palaeogeogr. Palaeoclimatol. Palaeoecol., 162, 389401.
  • Zhou, W. J., M. J. Head, and L. Deng (2001), Climate changes in northern China since the late Pleistocene and its response to global change, Quat. Int., 83–85, 285292.
  • Zhou, W. J., X. F. Lu, Z. K. Wu, L. Deng, A. J. T. Jull, D. Donahue, and W. Beck (2002), Peat record reflecting Holocene climatic change in the Zoige Plateau and AMS radiocarbon dating, Chin. Sci. Bull., 47, 6670.

Supporting Information

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Material and Methods
  5. 3. Results and Discussion
  6. 4. Conclusions
  7. Acknowledgments
  8. References
  9. Supporting Information
FilenameFormatSizeDescription
jgrd14189-sup-0001-t01.txtplain text document2KTab-delimited Table 1.

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