Teleconnection of climatic events between East Asia and polar, high latitude areas during the last deglaciation

ELSEVIER

Palaeogeography, Palaeoclimatology, Palaeoecology 152 (1999) 163–172

Teleconnection of climatic events between East Asia and polar, high latitude areas during the last deglaciation Weijian Zhou a,Ł , M.J. Head a,1 , Xuefeng Lu a , Zhisheng An a , A.J.T. Jull b , D. Donahue b a

State Key Laboratory of Loess and Quaternary Geology, Chinese Academy of Sciences, P.O. Box 17, Xian 710054, China b NSF–AMS Facility, University of Arizona, Tucson, AZ 85721, USA Received 18 August 1998; revised version received 1 January 1999; accepted 13 January 1999

Abstract From observations of continuous aeolian and swamp sediment sequences, augmented by a detailed 14 C chronology, δ C and organic C analyses, we can demonstrate century- to millennium-scale fluctuations in East Asian monsoon palaeoclimatic events. We also infer significant precipitation variability within the last deglaciation. The major climatic zones that have been recognised in Europe, and find counterparts in East Asia are, the Bølling–Allerød (14,750–12,800 cal yr B.P.), and the Younger Dryas (12,800–11,600 cal yr B.P.). The last deglaciation sediment sequences are characterised by frequent facies changes reflecting climate instability. These frequent, abrupt climatic events correlate well with fluctuations recorded in high latitude and polar areas, as represented by the Greenland ice core GISP 2 and a core from the North Sea. This indicates a palaeoclimate teleconnection between the high latitude and polar areas and the East Asian monsoon areas through cold air mass activity, and the related atmospheric pressure system.  1999 Elsevier Science B.V. All rights reserved. 13

Keywords: cold surge; East Asian monsoon; Younger Dryas; Chinese Loess Plateau; radiocarbon dating; precipitation variability

1. Introduction The last deglaciation was a period of intense and rapid change, which affected the global climate (Broecker et al., 1988). Evidence of these fluctuations has been well documented in deep-sea cores (Karpuz and Jansen, 1992; Rochon et al., 1998) and Greenland ice cores (Taylor et al., 1993; GRIP Project Members, 1993; Grootes et al., 1993). Zhou et al. (1996) reported a climate record from the desert–loess transitional belt of China that describes Ł Corresponding

author. E-mail: [email protected] Present address: School of Geosciences, University of Wollongong, Wollongong, N.S.W. 2500, Australia 1

the East Asian equivalent of the European Younger Dryas interval, reflecting cold-dry to cool-wet, returning to cold-dry century-scale fluctuations. Wang et al. (1996) showed that at about 13,000 14 C yr B.P. (¾15,450 cal yr B.P.), the water volume of Gu Cheng Lake, Gansu Province, expanded significantly, and later fluctuated over short time periods. However, high resolution localities in China producing comparative data have not as yet been studied in detail. These types of studies also require a reliable, detailed chronological framework augmented by suitable climate proxies. This paper is based on a detailed chronological framework and climate proxies from the Midiwan swamp–peat sequences together with the Baxie and

0031-0182/99/$ – see front matter  1999 Elsevier Science B.V. All rights reserved. PII: S 0 0 3 1 - 0 1 8 2 ( 9 9 ) 0 0 0 4 1 - 3

164

W. Zhou et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 152 (1999) 163–172

Yangtaomao aeolian sequences, and reconstructs the East Asian monsoon variation history from 13,000 to 10,000 14 C yr B.P. (15,450–11,600 cal yr B.P.). Our reconstruction is then compared with palaeoclimatic records from the high latitude areas in the Northern Hemisphere in order to improve our understanding of global-scale climatic change during the last deglaciation period. However, it must be recognised that as many as three distinct, constant radiocarbon age plateaux have been identified during this late glacial period (Lotter et al., 1992; Kitagawa and van der Plicht, 1998). This makes any absolute chronological correlation virtually impossible without extremely detailed cross-referencing. Hence correlation of event sequences within a relative time framework is important. Conventional radiocarbon ages have been calibrated to calendar ages in order to correlate the climate variations as recorded from the Chinese Loess Plateau, data from GISP 2 and a core from the North Sea. The computer software used was CALIB version 3.0.3c (Stuiver and Reimer, 1993). It has already been suggested (Porter and An, 1995) that there is a close connection between the North Atlantic Heinrich events and the Asian winter monsoon via the Mongolian high pressure system, as indicated by increased loess deposition on the Chinese Loess Plateau. Present-day meteorological observations indicate that the Mongolian high pressure system has cold air contributions from the Barents Sea 56%, the Kara Sea, 25%, and the North Atlantic, 15% (Tao and Chen, 1957). A simulation of the Northern Hemisphere winter sea level pressure and dominant wind direction for the last glacial maximum (Kutzbach et al., 1993; Fig. 5) proposes the formation of strong storm tracks by which the cold surge would have travelled through Europe into central China. In addition, there is good evidence to show that for the last deglaciation, teleconnections between Asia and the high latitude areas also existed during summer (Sirocko et al., 1996).

2. Last deglaciation stratigraphy and 14 C chronology The location of each of the three profiles used for this study is shown in Fig. 1, and site stratigraphy for each profile is shown in Fig. 2. Since this study

is only concerned with the time period covering the last deglaciation, the Holocene component of the stratigraphic sequences is not discussed in detail. The Midiwan peat deposit (37º390 , 108º370 , ca. 1400 m altitude; Figs. 1 and 2), is a continuous profile with a thickness of about 1380 cm (Zhou et al., 1996), and is located on the first terrace of the Xuanhe, which is currently a seasonal tributary of the Heiheze River, a second-order tributary of the Huang He (Yellow River), to the northeast of Jingbian County, Shaanxi Province. The late glacial stratigraphy consists of a layer of grey to black organic silty mud (1220–1240 cm), overlain by units of pale yellow to light greyish-green silts (1140–1220 cm), greyish-black silty peat (1090–1140 cm), black to grey muddy silts containing freshwater molluscs (970–1090 cm), greyish-black silty peat containing plant residues and a thin layer of carbonate nodules (900–970 cm), and light yellow aeolian sand (835– 900 cm). AMS 14 C ages from peat, charcoal and wood bracket the sequences between 13,020 š 90 and 10,025 š 80 14 C yr B.P. Cellulose was extracted from the wood using sodium chlorite, charcoal was pretreated using the acid–alkali–acid technique, and the <80 µm size fraction of the peat was used for dating (Head et al., 1989; Zhou et al., 1990; Zhou et al., 1996). The results (Fig. 2) indicate that the 14 C ages from the different material types are in good agreement with each other. Yangtaomao (38º480 , 110º270 , ca. 1400 m altitude; Figs. 1 and 2) is a loess-mantled hill consisting of sand and interstratified palaeosols exposed by wind erosion, 150 km northeast of Midiwan (Zhou et al., 1996). The profile is 350 cm thick, and the late glacial stratigraphy consists of an immature sandy palaeosol overlain by pale yellow aeolian sand containing a 38 cm thick immature sandy palaeosol (242–204 cm depth). The sequence is bracketed by two AMS 14 C ages, 11,250 š 80 14 C yr B.P. (pollen), from near the top of the lower palaeosol (ca. 320 cm depth), and 9645 š 80 14 C yr B.P. (humic acid), from near the bottom of a further sandy palaeosol (ca. 150 cm depth) (Fig. 2). Six palaeosol samples were collected from this profile, and were treated with acid, alkali and acid, and then with selected organic solvents to remove younger organic impurities. The dating was carried out on the acid and alkali insoluble residues (Head et al., 1989; Zhou et al., 1990).

W. Zhou et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 152 (1999) 163–172

165

Fig. 1. The Chinese Loess Plateau, with sites along the desert–loess boundary that are mentioned in this study. Inset map indicates the location of the study area within China.

The Baxie loess profile (35º330 , 103º350 , altitude 1800 m; Figs. 1 and 2), is located in Dongxian County, Gansu Province (Zhou et al., 1992; An et al., 1993). The profile is about 11 m thick, and the late glacial stratigraphy consists of a broad palaeosol at about 980 cm depth, overlain by loess, weakly pedogenic palaeosol and loess units to 800 cm depth, which can be correlated chronologically with that of Midiwan. The sequence is bracketed by two AMS 14 C ages of 11,980 š 75 and 10,250 š 70 14 C yr B.P. This profile contains very scarce amounts of material suitable for 14 C dating, and as a consequence has proven extremely difficult to date accurately, as the initial work was carried out using the bulk organic component and humic fractions of palaeosols. In order to improve the Baxie sequence chronology, pollen concentrates were extracted from paleosol material, and well preserved terrestrial snail shells were collected. XRD analysis indicated that the shells consisted of pure aragonite, thus making them suitable for dating. No living snail shells could

be located within the area, but organic acid and alkali insoluble 14 C ages, and 14 C ages of inorganic carbonate from the same levels agree to within one standard deviation indicating that the reservoir correction factor for the snail shells of this profile would be minimal. The new 14 C ages provide a much more detailed chronology and the shell ages fit quite nicely within the pollen age sequence (Zhou et al., 1997). Using the chronological framework provided by the 14 C ages, we are able to make a comparison between the stratigraphy of the swamp–peat sequences at Midiwan, the loess–paleosol sequences at Baxie and the loess–palaeosol sequences at Yangtaomao (Fig. 2). During the last deglaciation, frequent changes in sediment deposition occurred; these are more easily recognised in the swamp–peat sequences of Midiwan than in the loess–palaeosol sequences of Baxie and Yangtaomao. These changes illustrate the century- to millennium-scale climatic fluctuations that occurred within the Chinese Loess Plateau and the desert–loess transitional zone.

166

W. Zhou et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 152 (1999) 163–172

Fig. 2. Stratigraphic correlation of profiles. Midiwan swamp–peat; Yangtaomao loess–palaeosol; Baxie loess–palaeosol. The 14 C ages for this study were carried out in the NSF–AMS Dating Facility at the University of Arizona. Conventional 14 C ages are listed on the left-hand side of each profile, and calibrated ages (Stuiver and Reimer, 1993) are listed on the right-hand side.

3. Proxy record obtained from sites on the Loess Plateau We carried out δ13 C analyses of total organic carbon in sediment samples from Midiwan and Baxie. At Midiwan, samples were taken at 5 cm intervals within the peat units (giving a 60 year resolution), and 10 cm intervals for the aeolian sand units. At Baxie, and Yangtaomao, samples were taken at 10 cm intervals (giving a 110 year resolution). Based on work carried out by other researchers (Hakansson, 1985; Hammarlund, 1992; Sukumar et al., 1993), it can be inferred that δ13 C values from the Midiwan and Baxie sediments reflect changes in environmental conditions and hence can be used as a valuable palaeoclimate proxy indicator (more negative δ13 C values indicate an increased tree to grass ratio variation in vegetation types, hence increased precipita-

tion or humidity, while more positive δ13 C values indicate colder, dry conditions, decrease in vegetation and predominance of C4 grasses). It has been demonstrated that the organic carbon content of sediments can also be used as an indirect indicator of vegetation cover and biomass, and hence increased precipitation or humidity (Zhou et al., 1996). Correlation with % organic C analysis and detailed pollen studies at Midiwan (Zhou et al., 1996) justify the value of the δ13 C results as a proxy indicator of vegetation cover and hence moisture. The Artemisia=Chenopodiaceae ratio (El-Moslimany, 1990; Sun et al., 1994) can be regarded as a good indicator of moisture in arid to semi-arid zones. The relationship between δ13 C values and the Artemisia=Chenopodiaceae ratio for the Midiwan sequence is extremely good (Zhou et al., 1996). Calibrated 14 C ages (Stuiver and Reimer, 1993) are here preferred to conventional 14 C ages

W. Zhou et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 152 (1999) 163–172

167

Fig. 3. Climate proxy sequences for Midiwan, Baxie and Yangtaomao. δ13 C values with respect to the PDB standard are plotted against depth for Midiwan and Baxie. Percentage carbon values are plotted against depth for the Yangtaomao profile (Zhou, 1998). Calibrated 14 C ages are indicated for Midiwan against the appropriate depth, together with specific calibrated ages in parentheses. These have been calculated from 14 C ages by extrapolation and are used for discussion purposes only.

in order to correlate palaeoclimatic data with the δ18 O record from the GISP 2 ice core. In addition, estimated calibrated ages (Figs. 3 and 4, in parentheses), based on extrapolation of 14 C data, have been used to pinpoint specific peaks, purely for discussion purposes. From the data indicated in Fig. 3, it can be seen that the beginning of the last deglaciation (The European Bølling–Allerød period, shown by dotted lines) is reflected in the Midiwan δ13 C proxy record from about 1160 cm, at about 14,750 cal yr B.P. (¾12,800 14 C yr B.P.). The period preceding this will be described separately. From 14,750 cal yr B.P., the δ13 C values change rapidly from ca. 25‰ to 29.5‰, consistent with a change to slightly warmer, wetter conditions, with a peak occurring at about 14,500 cal yr B.P. A period of sharp, minor fluctuations in δ13 C values then occurs until a major negative peak appears at about 14,100 cal yr B.P. (at 1100 cm depth). This sudden negative shift in values could be con-

sidered to mark the European Bølling–Allerød transition, which in some places produces even warmer, wetter conditions, though regional differences have been documented in northwestern Europe (Coope and Lemdahl, 1995; Coope et al., 1998). The δ13 C values for the Midiwan profile then show a sharp shift towards more positive values indicating a cold, dry phase at about 1090 cm depth. The δ13 C values move to a negative peak at 1050 cm depth, at about 13,200 cal yr B.P. They then drop sharply towards more positive values at the 1020 cm level at about 12,800 cal yr B.P. (¾11,000 14 C yr B.P.). Hence, during the European Bølling–Allerød interstadial, the Midiwan δ13 C proxy curve exhibits three major negative peaks indicating warmer, wetter conditions. The δ13 C proxy curve for the Baxie profile indicates the beginning of the last deglaciation (14,750 cal yr B.P., 12,800 14 C yr B.P.) at a depth of about 975 cm. The δ13 C values become more negative until a peak is reached at about 950 cm. This corresponds

168

W. Zhou et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 152 (1999) 163–172

Fig. 4. Correlation of climate proxy sequences for Midiwan, the North Sea, and the GISP 2 ice core. The Midiwan sequence is used as a representative of the Chinese sequences for this correlation. For Midiwan, δ13 C values with respect to the PDB standard are plotted against depth. For the North Sea core, reconstructed SST values for winter and summer, together with sea ice cover (months per year) are plotted against depth (from Rochon et al., 1998). δ18 O values are plotted against calendar age for the GISP 2 ice core (from Taylor et al., 1993). All 14 C ages have been calibrated (Stuiver and Reimer, 1993) so that a direct correlation could be made with the GISP 2 ice core. Calibrated ages in parentheses are as for Fig. 3.

to a 14 C age of 11,980š75 14 C yr B.P. (Fig. 2), which fits in well with the major peak in the Midiwan data at ¾14,100 cal yr B.P. (1100 cm depth). The δ13 C values fluctuate, then become more negative until a second peak is reached at about 925 cm depth (about 13,200 cal yr B.P.), which corresponds to the last peak in the Bølling–Allerød interval shown in the Midiwan curve. The values then become sharply more positive indicating the onset of the Younger Dryas period. In this record, only two major peaks can be seen. This could be an artifact of the lower degree of sampling resolution obtained for this profile. The relative scarcity of data for the Yangtaomao plot of organic carbon levels against depth for this period caused by the limited depth of the profile, does not warrant discussion. The Bølling–Allerød interstadial generally produced warmer, wetter conditions resulting in an in-

crease in vegetation cover. This is born out by pollen analysis (Zhou et al., 1996), which shows increased tree and herb pollen concentrations. This is matched by an increase in organic carbon concentration, together with generally more negative δ13 C values. From 12,800 to ¾11,700 cal yr B.P. (11,000 to 10,000 14 C yr B.P.), as shown by the dotted lines in Fig. 3, the proxy indicators show the sudden onset of a much colder dry period corresponding to the European Younger Dryas. This is punctuated by a period between about 12,500 and 12,100 cal yr B.P. (¾10,600–10,200 14 C yr B.P.), in which δ13 C values switch rapidly back to more negative values and organic C shows a rapid rise, followed by a rapid shift to more positive δ13 C values and a lower organic C content between ¾12,100 and 11,700 cal yr B.P. (¾10,200 and 10,000 14 C yr B.P.). This reflects the existence of a virtual three-phase component of the

W. Zhou et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 152 (1999) 163–172

European Younger Dryas event within sedimentary profiles on the Chinese Loess Plateau. This is indicated more thoroughly in Fig. 2, where, for this period the Midiwan profile consists of units of silty sand overlain by a silty peat and aeolian sand. The Yangtaomao profile consists of units of aeolian sand overlain by palaeosol, then aeolian sand. The Baxie profile consists of units of loess overlain by palaeosol, and loess. From the proxy indicators, the Younger Dryas period is characterised by unstable climatic conditions, with rapid but significant fluctuations between wet and dry periods. The δ13 C data recorded in sand or loess point to a much colder and drier episode towards the end of the Younger Dryas period. This is supported by magnetic susceptibility measurements, and pollen analytical data (Zhou et al., 1996). The onset of the early Holocene is indicated by a trend towards more negative δ13 C values with rapid fluctuations in both the Midiwan and Baxie proxy curves. The Preboreal is represented generally by a trend towards more positive δ13 C values, corresponding to a unit of predominantly inorganic lacustrine silt at Midiwan. The warmest period is indicated at Midiwan at 760 cm depth, at about 10,200 cal yr B.P. The Midiwan and Baxie proxy curves then reflect a general trend towards a colder, drier climate, still with frequent fluctuations. The Yangtaomao proxy curve shows a maximum at about 130 cm depth, which is equivalent to the maximum at Midiwan, and which occurs within the same time range. The values then drop off, indicating a similar trend towards colder, drier conditions.

4. Correlation of swamp sequences with those from high latitudes Fig. 4 illustrates the depth sequences (with calibrated age markers indicated) of the δ13 C values from Midiwan, as a representative of proxy sequences from the Chinese Loess Plateau, compared with plots of reconstructed sea surface temperatures and sea ice cover in months per year based on dinoflagellate cyst data obtained from cores collected in the Norwegian channel for February and August, with calibrated age markers (Rochon et al., 1998). A plot of δ18 O values for the GISP 2 ice core (Taylor et al., 1993) is also used for correlation purposes.

169

For the monsoon-influenced area of China, the last deglaciation climate between 14,750 and 12,800 cal yr B.P. (¾13,050–11,000 14 C yr B.P.), the European Bølling–Allerød interstadial event (shown by dotted lines in Fig. 4), the Midiwan δ13 C values show three major peaks at 14,500, 14,100 and 13,200 cal yr B.P. No obvious change is indicated in the North Sea data. The sea ice cover remains minimal, and both summer and winter sea surface temperatures remain constant. Winter SST values start decreasing at the 390 cm level, with a corresponding decrease in summer SST values and an increase in the sea ice cover, reflecting the onset of the Younger Dryas period. The δ18 O data from the GISP 2 ice core records the Bølling–Allerød interstadial event from about 14,750 to 12,800 calendar years, during which the Bølling–Allerød transition is placed at 14,060 calendar years (Taylor et al., 1993). During this period the δ18 O plot shows a general trend towards more negative values, with three peaks at 14,500, 13,500 and about 13,000 calendar years. Hence the GISP 2 ice core trends are almost synchronous with those illustrated by the Midiwan data (given the limitations of the 14 C calibration data). From 12,800 to 11,600 cal yr B.P. (¾11,000 to 10,000 14 C yr B.P.), the GISP 2 ice core δ18 O indicates much more negative values, reflecting the much colder conditions associated with the onset of the European Younger Dryas period. The GISP 2 record fluctuates considerably during this period, and indicates slight changes to more positive δ18 O values at about 12,250 and 11,900 cal yr B.P. This correlates well with the Midiwan δ13 C record, at depths between 900 and 940 cm, probably reflecting increased precipitation. Both the GISP 2 and the Midiwan records fluctuate sharply until approximately 11,600 cal yr B.P., where much more positive δ18 O values in the GISP 2 record are indicated. During the Younger Dryas period, there is a striking similarity between the North Sea record and the Midiwan δ13 C proxy data. Summer SST values, together with sea ice cover in the North Sea, fluctuate in a similar manner to the Midiwan δ13 C values, while winter SST values stay low and constant for the whole of the Younger Dryas period. The increased precipitation during the mid-Younger Dryas reflected at Midiwan correlates well with the

170

W. Zhou et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 152 (1999) 163–172

high summer temperature and low sea ice cover in the North Sea. The sharp decrease in summer SST and elevated ice cover towards the last stage of the Younger Dryas exactly matches the last stage of cold and dry conditions recorded in coarse-grained aeolian sand or loess layers (Fig. 2), together with more positive δ13 C values indicated from the 875 cm level at Midiwan. This evidence is strongly supported by pollen analysis and magnetic susceptibility determinations (Zhou et al., 1996). In summary, it is interesting to note the frequent fluctuations present in the GISP 2 ice core δ18 O record for the Younger Dryas period, which correlate very well with the sharp fluctuations indicated in both the Midiwan and North Sea records. This indicates that the Younger Dryas period in the Northern Hemisphere was a period of climatic instability. The presence of other, more regional climate mechanisms affecting terrestrial northwestern Europe during the Younger Dryas period is indicated by data subdividing the same chronological period into three phases beginning with a wet, very cold period reported between 10,900 and 10,550 14 C yr B.P., followed by a slightly warmer, very dry period between 10,550 and 10,150 14 C yr B.P., then a third period of warmer, wetter conditions from 10,150 to 9850 14 C yr B.P. This warmer period has been interpreted as the beginning of the Preboreal (Bohncke et al., 1993; Vandenburghe, 1995). The onset of the early Holocene is indicated by a trend towards negative δ13 C values with rapid fluctuations in the Midiwan proxy curve. The warmest period is indicated at Midiwan at 760 cm depth (about 10,200 cal yr B.P.), and the curve then indicates a general trend towards more positive δ13 C values, reflecting a colder, drier climate, still with frequent fluctuations. The GISP 2 ice core record moves sharply towards more positive values of δ18 O, showing a general trend towards warmer temperatures, with frequent sharp fluctuations. The period of possible Preboreal conditions reflected by a silt horizon at Midiwan from a depth of 830 to 800 cm (also recorded in the proxy curve at about 11,700 to 11,000 cal yr B.P., Figs. 2 and 4), indicates a slight increase in precipitation levels, but does not follow the sharp change indicated in other records. The possible Preboreal change in the early Holocene as reflected in the Midiwan sediments needs to be further investigated.

Correlation of the sites along the desert–loess transition zone in northern China and the northern polar latitude areas during the last deglaciation indicates that fluctuations in East Asian monsoon precipitation are generally related to fluctuations in North Sea SST, and to temperature variations reflected in the δ18 O record of the GISP 2 ice core. The most likely mechanism for this is via a weakening Mongolian high pressure system and cold surge activity.

5. Possible cause of the observed effects From about 14,750 cal yr B.P., Northern Hemisphere insolation gradually increased, there was a decrease in ice cover, sea ice retreated, and ocean warming occurred. The increase of seasonality of solar insolation during this period was in response to the increase of tilt and adjustment of summer perihelion in the Northern Hemisphere (Kutzbach et al., 1993). This also led to an ‘unstable state’ in the continent–atmosphere–ocean coupling system. Changes in any one of the factors such as SST, break up of icebergs, atmospheric CO2 variations and others can cause higher amplitude changes in the earth system, which lead to changes in material and energy transportation modes within continent, ocean, and atmosphere. It would cause a change in the atmospheric circulation system and possibly induce a change in local, or even global climate. For example, the release of ice melt water and icebergs from ice sheets into the North Atlantic ocean, interrupting normal circulation and efficient operation of high density deep water currents (Broecker and Rind, 1985). There would have been a decrease in heat transfer from the ocean to the atmosphere, leading to a decrease in heat storage in summer and a rapid cooling effect in autumn, enlarging the extent of ice cover in autumn and winter. This change first influenced Europe, and subsequently influenced northern China, Japan, and the mid-high latitudes of North America. GCM results published by Kutzbach et al. (1993) (Fig. 5), indicate a storm track of cold air in winter from the North Atlantic area through Eurasia to the Loess Plateau during the last glacial maximum. It is quite feasible that during subsequent cold events, the climates of East Asia and the polar high lati-

W. Zhou et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 152 (1999) 163–172

171

Fig. 5. The Northern Hemisphere winter sea level pressure systems, dominant wind direction and the path of major storms for 18,000 yr B.P. (after Kutzbach et al., 1993). Legend: 1 D Surface storm tracks; 2 D iceberg areas; 3 D ice sheet.

tude areas were connected via this cold air activity. Porter and An (1995), through a study of loess grain size, showed that the winter monsoon record from the East Asian monsoon area correlated well with Heinrich cooling events in the North Atlantic Ocean. Consequently, the westerlies and associated Mongolian high pressure systems have an influence on the winter monsoon circulation of China. As precipitation in the East Asian monsoon area is monsoon front precipitation, it relies on the interaction of cold air from high latitude polar regions with warm, moist air from the ocean. If we take solar radiation strengthening during 15,000–10,000 cal yr B.P. as being relatively constant, the variability of monsoon precipitation depends on the changes in strengthening and weakening of the cold air mass from polar, high latitudes, and related Mongolian high pressure variations during cold periods such as the Younger Dryas. When these cold air masses are strong, the monsoon front moves southwards. The geological evidence with century- to millennial-scale fluctuations indicated for the Younger Dryas interval is essentially an amplification of climatic signals in the Northern Hemisphere connected through to the East Asian monsoon system. It is still open to question whether this teleconnection between Northern Hemisphere polar, high

latitude conditions and East Asia is also valid for the Holocene, and this will be addressed in the near future.

Acknowledgements This study was supported by the following grants: NSFC 49725308, and 49894170; U.S. National Science Foundation grant EAR 95-08413; CAS KZ951-Ai-402 & KZ-952-S1-419; SSTC 95-40. The authors would like to thank Professor J. Kutzbach for kind permission to use his GCM data for Fig. 5, Professor J. Vandenburghe and an anonymous referee for their constructive advice.

References An, Z.S., Porter, S.C., Zhou, W.J., Lu, Y., Donahue, D.J., Head, M.J., Wu, X.H., Ren, J.Z., Zheng, H.B., 1993. Episode of strengthened summer monsoon climate of Younger Dryas age on the Loess Plateau of Central China. Quat. Res. 39, 45–54. Bohncke, S., Vandenburghe, J., Huijzer, A.S., 1993. Periglacial environments during the Weichselian Late Glacial in the Maas valley, the Netherlands. Geol. Mijnbouw 72, 193–210. Broecker, W.S., Peteet, D.M., Rind, P., 1985. Does the ocean– atmosphere system have more than one stable mode of opera-

172

W. Zhou et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 152 (1999) 163–172

tion? Nature 315, 21–26. Broecker, W.S., Andree, M., Wolfli, W., Oeschger, H., Bonani, G., Kennett, J., Peteet, D., 1988. The chronology of the last deglaciation: implications to the cause of the Younger Dryas event. Paleoceanography 3, 1–19. Coope, G.R., Lemdahl, G., 1995. Regional differences in the Lateglacial climate of northern Europe based on coleopteran analysis. J. Quat. Sci. 10 (4), 391–395. Coope, G.R., Lemdahl, G., Lowe, J.J., Walkling, A., 1998. Temperature gradients in northern Europe during the last glacial– Holocene transition (14–9 C-14 kyr BP) interpreted from coleopteran assemblages. J. Quat. Sci. 13 (5), 419–433. El-Moslimany, A.P., 1990. Ecological significance of common non-arboreal pollen: examples from drylands of the Middle East. Rev. Palaeobot. Palynol. 64, 343–350. GRIP Project Members, 1993. Climate instability during the last interglacial period recorded in the GRIP ice core. Nature 364, 203–207. Grootes, P.M., Stuiver, M., White, J.W., Johnsen, S., Jouzel, J., 1993. Comparison of oxygen isotope records from the GISP 2 and GRIP Greenland ice cores. Nature 366, 552–554. Hakansson, S., 1985. A review of various factors influencing the stable carbon isotope ratio of organic lake sediments by the change from glacial to post glacial environmental conditions. Quat. Sci. Rev. 4, 135–146. Hammarlund, D., 1992. A distinct δ13 C decline in organic lake sediments at the Pleistocene–Holocene transition in southern Sweden. Boreas 22, 236–243. Head, M.J., Zhou, W.J., Zhou, M.F., 1989. Evaluation of 14 C ages of organic fractions of paleosols from loess–paleosol sequences near Xian, China. Radiocarbon 31 (3), 680–696. Karpuz, N.K., Jansen, E.A., 1992. A high-resolution diatom record of the last deglaciation from the SE Norwegian Sea: documentation of rapid climatic changes. Paleoceanography 7 (4), 499–520. Kitagawa, H., van der Plicht, J., 1998. Atmospheric radiocarbon calibration to 45,000 yr BP: late glacial fluctuations and cosmogenic isotope production. Science 279, 1187–1190. Kutzbach, J., Gutter, P., Behling, P.J., Selin, R., 1993. Simulated climatic changes: results of the COHMAP climate-model experiments. In: Wright, H.E., Kutzbach, J.E., Webb III, T., Ruddiman, W.F., Street-Perrott, F.A, Bartlein, P.J. (Eds.), Global Climates since the Last Glacial Maximum. University of Minnesota Press, London, pp. 24–93. Lotter, A.F., Amman, B., Beer, J., Hajdas, I., Sturm, M., 1992. A step towards an absolute time-scale for the Late-Glacial: annually laminated sediments from Soppensee (Switzerland). In: Bard, E., Broecker, W.S. (Eds.), The Last Deglaciation: Absolute and Radiocarbon Chronologies. Springer, Berlin, pp. 45–68. Porter, S.C., An, Z.S., 1995. Correlation between climate events in the North Atlantic and China during the last glaciation.

Nature 375, 305–308. Rochon, A., de Vernal, A., Sejrup, H.-P., Haflidason, H., 1998. Palynological evidence of climatic and oceanographic changes in the North Sea during the last deglaciation. Quat. Res. 49, 197–207. Sirocko, F., Garbe-Schonberg, D., McIntyre, A., Molfino, B., 1996. Teleconnections between the subtropical monsoons and high-latitude climates during the last deglaciation. Science 272, 526–529. Stuiver, M., Reimer, P.J., 1993. Radiocarbon calibration. Radiocarbon 35, 215–230. Sukumar, R., Ramesh, R., Pant, R.K., Rajagopalan, G., 1993. A record of late Quaternary climate change from tropical peats in southern India. Nature 364, 703–705. Sun, X.J., Du, N.Q., Weng, C.Y., Lin, R.F., Wei, K.Q., 1994. Paleovegetation and paleoenvironment of Nanaisi Lake, Xinjiang, N.W. China during the last 14,000 years. Quat. Sci. 3, 239–248. Tao, S.Y., Chen, L.X., 1957. The atmospheric structure over the Asian continent in summer (in Chinese). Bull. Meteorol. 28, 234–246. Taylor, K.C., Lamorey, G.W., Doyle, G.A., Alley, R.B., Grootes, P.M., Mayewski, P.A., White, J.W.C., Barlow, L.K., 1993. The flickering switch of Late Pleistocene climate change. Nature 361, 432–436. Vandenburghe, J., 1995. The climate of the Younger Dryas in the Netherlands. Geol. Mijnbouw 74, 245–249. Wang, S.M., Yang, X.D., Ma, Y. et al., 1996. Discussion of relationship between paleomonsoon and environmental change of Gucheng Lake, Gansu Province since 15,000 years (in Chinese). Sci. China (Ser. D) 26 (2), 137–141. Zhou, J., 1998. East Asian monsoon change in the last 250 ka and the correlation of the paleoclimatic events beween northern and southern hemispheres. Unpublished PhD Thesis, Northwest University, Xian, China, 34. Zhou, W.J., Zhou, M.F., Head, J., 1990. 14 C chronology of Bei Zhuang Cun sedimentation sequence since 30,000 years BP. Chin. Sci. Bull. 35 (7), 567–572. Zhou, W.J., An, Z.S., Lin, B.H., Xiao, J.L., Zhang, J.Z., Xie, J., Zhou, M.F., Porter, S.C., Head, M.J., Donahue, D.J., 1992. Chronology of the Baxie Loess Profile and the history of monsoon climates in China between 17,000 and 6000 years BP. Radiocarbon 34 (3), 818–825. Zhou, W.J., Donahue, D.J., Porter, S.C., Jull, A.J.T., Li, X.Q., Stuiver, M., An, Z.S., Matsumoto, E., Dong, G.G., 1996. Variability of monsoon climate in East Asia at the end of the last glaciation. Quat. Res. 46, 219–229. Zhou, W.J., Donahue, D.J., Jull, A.J.T., 1997. Radiocarbon AMS dating of pollen concentrated from eolian sediments: implications for monsoon climate change since the late Quaternary. Radiocarbon 39 (1), 19–26.