Ecology, 89(4), 2008, pp. 1079–1088
Ó 2008 by the Ecological Society of America
ECOTONE SHIFT AND MAJOR DROUGHTS DURING THE MID–LATE
HOLOCENE IN THE CENTRAL TIBETAN PLATEAU
CAIMING SHEN,1,2,3,6 KAM-BIU LIU,2 CARRIE MORRILL,4 JONATHAN. T. OVERPECK,5 JINLAN PENG,3
AND
LINGYU TANG3
1
Atmospheric Sciences Research Center, State University of New York, Albany, New York 12203 USA
Department of Oceanography and Coastal Sciences, School of the Coast and Environment, Louisiana State University,
Baton Rouge, Louisiana 70803 USA
3
Nanjing Institute of Geology and Paleontology, Academia Sinica, Nanjing 210008 China
4
NOAA Paleoclimatology Program and Cooperative Institute for Research in Environmental Sciences, Boulder, Colorado 80305 USA
5
Institute for the Study of Planet Earth and Department of Geosciences, University of Arizona, Tucson, Arizona 88721 USA
2
Abstract. A well-dated pollen record from a large lake located on the meadow–steppe
ecotone provides a history of ecotone shift in response to monsoonal climate changes over the
last 6000 years in the central Tibetan Plateau. The pollen record indicates that the ecotone
shifted eastward during 6000–4900, 4400–3900, and 2800–1600 cal. yr BP when steppes
occupied this region, whereas it shifted westward during the other intervals when the steppes
were replaced by meadows. The quantitative reconstruction of paleoclimate derived from the
pollen record shows that monsoon precipitation fluctuated around the present level over the
last 6000 years in the central Tibetan Plateau. Three major drought episodes of 5600–4900,
4400–3900, and 2800–2400 cal. yr BP are detected by pollen signals and lake sediments.
Comparison of our record with other climatic proxy data from the Tibetan Plateau and other
monsoonal regions shows that these episodes are three major centennial-scale monsoon
weakening events.
Key words: ecotone; Holocene; lake; meadow; monsoon; pollen; precipitation; steppe; Tibetan Plateau.
INTRODUCTION
An ecotone is defined as a transitional zone between
adjacent ecological systems (Holland and Risser 1991).
Because many ecotone species are at the limits of
distributional ranges, it has been suggested that ecotones
are areas sensitive to environmental changes, and that
changes in the location of ecotones serve as detectors of
climatic change (di Castri et al. 1988, Neilson 1993,
Allen and Breshears 1998, Kullman 1998). Hence, the
history of ecotone shift can provide information about
paleoclimate (e.g., Liu 1990, Crumley 1993, Kullman
1995, Reasoner and Jodry 2000, Baker et al. 2002).
Some studies have revealed many strong hints of
centennial-scale variability in the southwest Asian
monsoon (Gasse and Van Campo 1994, Guo et al.
2000, Zhou et al. 2001, Gupta et al. 2003, Morrill et al.
2003, 2006, Overpeck et al. 2005). However, more welldated and high-resolution monsoon proxy records are
needed to develop a detailed picture of regional
monsoon variability on the century timescale so that
we may understand the temporal and spatial patterns
and mechanisms for this variability (Morrill et al. 2003,
Overpeck et al. 2005).
On the Tibetan Plateau, vegetation distribution is
controlled by the summer monsoon precipitation. Along
Manuscript received 8 December 2006; revised 26 July 2006;
accepted 30 July 2006. Corresponding Editor: C. C.
Labandeira.
6 E-mail: [email protected]
a precipitation gradient from southeast to northwest,
vegetation changes from forest to meadow, to steppe,
and to desert (Fig. 1a). The meadow–steppe ecotone
(MSE) roughly follows the 400 mm-isohyet close to the
northern boundary of the modern summer monsoon in
central Tibet (Tibetan Investigation Group 1988).
Therefore, a pollen record from Co (Co means lake)
Ngion located on the MSE would yield a history of the
MSE shift and a sensitive record of monsoonal climate
changes.
LOCATION
AND
0
SETTING
Co Ngion (31824 –31832 N, 91828 0 –91833 0 E, 4515 m)
is situated in the central Tibetan Plateau (Fig. 1a). It is a
large calcareous fresh lake, about 14.8 km in length, 4.1
km in width, and 61.3 km2 in area. Its maximum water
depth is 5 m, and average depth about 3.5 m. Its
catchment area is 1019.7 km2. Fourteen streams flow
into the lake. Quaternary lacustrine plains and hills of
Cretaceous rocks surround this tectonic lake (Wang and
Dou 1998).
The study region has a cold semiarid climate. The
mean annual temperature is 1.4 to 0.98C, with mean
January temperature of 13.1 to 11.28C and mean July
temperature (Tjuly) of 8.5–9.08C. The mean annual
precipitation (MAP) is 300–400 mm, mostly originating
from the southwest summer monsoon. About 90% of the
annual precipitation is concentrated in the months of
May to September (Fig. 1b; Tibetan Investigation
Group 1988).
1079
0
1080
CAIMING SHEN ET AL.
Ecology, Vol. 89, No. 4
FIG. 1. (a) Map of the Tibetan Plateau showing the vegetation, location of surface pollen samples, and Co Ngion. Note that the
map is slightly skewed because of the projection angle. (b) A summary of regional climate; lines are temperature, and bars are
precipitation. (c) Regional vegetation in the central Tibetan Plateau.
The Stipa purpurea community is the typical steppe
growing in the areas west of the ecotone (Fig. 1c). Its
coverage is about 20–40%. Artemisia wellbyi, A. minor,
Carex montis-everestii, Astragalus heydei, Dracocephalum heterophyllum, Festuca ovina, Kobresia macrantha,
Leontopodium pusillum, and Poa litwinowiana are
frequently accompanying species (Wang 1988). The
Kobresia pygmaea community, the most widely distributed alpine meadow in the Tibetan Plateau, occurs in the
areas east of the ecotone. Its coverage is about 60–90%.
Besides the dominant species Kobresia pygmaea, commonly found species include Cyperaceae species such as
Carex moorcroftii, K. humilis, K. royleana, and K.
prainii; Poaceae species such as Poa calliopsis, Stipa
purpurea, Festuca nitidula, Deyeuxia tibetica, and
Roegneria brevipes; and other family species such as
Leontopodium nanum, Potentilla bifurca, Thalictrum
alpinum, and Polygonum macrophyllum. In the lowlands
April 2008
ECOTONE SHIFT IN THE TIBETAN PLATEAU
of lake basins are marsh meadows, which mainly consist
of Kobresia littledalei, K. royleana, K. pygmaea, Carex
satakeana, C. moorcrofti, and Blysmus sinocompressus
(Tibetan Investigation Group 1988).
MATERIALS
AND
METHODS
A 110-cm piston core was taken from the center of Co
Ngion (30828 0 N, 91830 0 E) under 4.8 m of water (see
Plate 1). Loss-on-ignition analysis (LOI) of sediment at
1-cm intervals was conducted at 1058C, 5508C, and
10008C to determine the contents of water, organic,
carbonate, and minerogenic matter in the sediments
(Dean 1974). Subsamples of 0.9 cm3 for pollen analysis
were extracted at 2-cm intervals. Samples were processed
using standard methods involving HCl, KOH, HF, and
acetolysis treatment (Faegri and Iversen 1975). Tablets
containing a known quantity of Lycopodium spores were
added as exotic markers into the samples to determine
pollen concentrations and pollen influx values (Stockmarr 1971). Thirty-five pollen types were identified.
More than 500 (500–785) pollen grains, excluding spores
and algal colonies, were counted at each level. Pollen
percentages were calculated based on a sum of all
terrestrial plant pollen including Cyperaceae pollen.
A pollen database consisting of 227 surface samples
from the Tibetan Plateau (Fig. 1a) was used to develop
discriminant functions for vegetation reconstruction
through discriminant analysis (Liu and Lam 1985, Shen
2003). The results of discriminant analysis were summarized in two indices (for details, see Liu and Lam
1985). The probability of modern analogue is determined by the measurement of similarity between a
sample and other samples of the same group. The
vegetation zonal index, which is converted from
probabilities of group membership between the predicted group (e.g., meadow) and the second most probable
group (e.g., steppe), assigns a fossil pollen spectrum to a
vegetation type. A subset of 80 surface samples from
meadows, steppes, and deserts was selected for developing modern pollen–climate transfer functions. The
surface samples were collected from moss polsters, top
soils, and lake bottom sediments. Transfer functions
were built using an inverse linear (IL) regression model
(Webb and Clark 1977, Birks 1995) and a weightedaveraging partial least squares (WA-PLS) regression
model (ter Braak and Juggins 1993, Birks 1995). These
two models were selected because the former is a linearbased method and the latter is a nonlinear, unimodalbased technique; hence they complement each other.
The MAP equation, which was derived from the IL
regression model, accounts for 88% of the total variance
and has a standard error of 39 mm, whereas the Tjuly
equation accounts for 81% of the total variance and has
a standard error of 0.98C (Shen et al. 2006). The
performance of WA-PLS regression models is shown in
Table 1, where the root mean square error of prediction
(RMSEP), the maximum bias, and the coefficient of
determination (r2) between observed and predicted
1081
TABLE 1. Performance statistics for weighted-averaging partial
least squares (WA-PLS).
WA-PLS
component
RMSEP
r2
Maximum bias
MAP
1
2 3
4
5
6
51.91
41.05
41.34
43.28
48.25
49.59
0.75
0.84
0.84
0.83
0.79
0.79
138.80
81.25
94.08
108.36
103.66
106.69
Tjuly
1
2
3 4
5
6
1.24
1.11
1.06
1.01
1.01
1.03
0.59
0.67
0.70
0.73
0.73
0.72
2.56
1.85
1.86
1.90
1.79
1.73
Note: Key to abbreviations: RMSEP, root mean square error
of prediction; MAP, mean annual precipitation; Tjuly, mean
July temperature.
Selected models.
values are based on jack-knifing (Birks 1995). According
to Birks (1998), the models with low RMSEP, low
maximum bias, and smaller number of components are
ideal candidates. Two- and three-component WA-PLS
models are thus selected for MAP and Tjuly, respectively.
STRATIGRAPHY
AND
CHRONOLOGY
The sediments of the core are mainly composed of
clayey silt. But at 91.5–93 cm, a unit of peat with 30%
organic matter is mixed with clay (Fig. 2). A unit
composed of clay and very thin peat layers with 22–30%
organic matter is found at 59–61 cm. The peat layers
consist of Potamogeton, an aquatic plant living in
shallow water of lakes in central Tibet. Results of LOI
show three periods of high organic matter contents
(more than 25%). The sediments of the core generally
contain about 15% carbonate, though somewhat less in
some units rich in organic matter. Minerogenic matter
shows a trend of nonlinear increase and three cycles
from low (;60%) to high values (;70–80%). The
sediments of the upper 35 cm contain more minerogenic
matter than those of the other parts except the interval
of 61–70 cm.
Radiocarbon dates for the Co Ngion core were
obtained from the aquatic macrophyte Potamogeton
and bulk organic sediments (Table 2). As expected, Co
Ngion has a hard-water effect like other lakes (e.g.,
Sumxi Co, Selin Co, and Lake Zigetang) in the Tibetan
Plateau (Fontes et al. 1993, Sun et al. 1993, Herzschuh et
al. 2006). The date of the surface sediments (the
uppermost 1.5 cm) indicates that the lake has a hardwater effect as long ago as ;2000 yr, which is close to
the date of 2010 6 50 14C yr BP determined from bulk
organic sediment in the uppermost 2-cm in Lake
Zigetang (Herzschuh et al. 2006), about 100 km
northwest of Co Ngion (Fig. 1c). Twelve radiocarbon
dates from this core suggest that the hard-water effect
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CAIMING SHEN ET AL.
FIG. 2.
Ecology, Vol. 89, No. 4
Lithology of sediment in the Co Ngion core, results of loss-on-ignition analysis (LOI), and age model.
was not constant through time, but relatively constant
during certain periods as shown by the depth–age
relationships (Fig. 2). The upper 35 cm of sediments
has the same hard-water effect, whereas the lower part
has an additional relatively constant effect. A hardwater correction of 2000 yr, as suggested by the date of
the surface sediments, was applied to the dates of the
upper 35 cm sediments. Corrected and calibrated dates
show that the sediments of the upper 35 cm were
deposited after ;3000 cal. yr BP (calendar years before
present). The dates of terrestrial charcoal and Potamogeton samples from Ahung Co (Fig. 1c), about 60 km
east of Co Ngion, indicate that there is relatively
constant hard-water effect of no greater than 600 yr
for the Potamogeton dates during the interval of 3000–
9000 cal. yr BP (Morrill et al. 2006). Considering the
TABLE 2. Radiocarbon dates for the Co Ngion core.
Depth
(cm)
Lab no.
0–1.5
15–16
31–32
45–46
58–59
60–61
60–61
71–72
82–83
91–92
91–92
106–107
AA51121
AA51122
AA51123
AA51124
AA38070
AA38069
AA38069
AA51125
AA51126
AA38068
AA38068
AA38067
14
C age
(14C yr BP)
2140
3750
4530
4400
4580
5450
4310
4730
5030
5290
4910
6090
6
6
6
6
6
6
6
6
6
6
6
6
50à
70à
50à
60§
50§
50§
50}
80§
60§
50§
50}
70§
Calibrated using CALIB4.3 (Stuiver et al. 1998).
à Approximately 2000-yr hard-water correction.
§ Approximately 1000-yr hard-water correction.
} Approximately 600-yr hard-water correction.
# Excluded from age model.
Calendar age
(cal. yr BP) 2r error range
(cal. yr BP)
Material
140
1690
2560
3660
3890
5090#
4070
4130
4560
4850
4890
5820
0–290
1520–1870
2370–2750
3480–3830
3720–4060
4880–5290
3910–4230
3870–4380
4300–4810
4670–5030
4730–5040
5660–5990
bulk organic
bulk organic
bulk organic
bulk organic
bulk organic
bulk organic
Potamogeton
bulk organic
bulk organic
bulk organic
Potamogeton
bulk organic
April 2008
ECOTONE SHIFT IN THE TIBETAN PLATEAU
1083
FIG. 3. (a) Pollen percentage diagram and (b) pollen influx diagram for the Co Ngion core.
same regional background of bedrock for Co Ngion and
Ahung Co, it is reasonable to recalculate Potamogeton
dates from the sediments below 35 cm by this hard-water
correction. Two dates from bulk organics and Potamogeton at 91–92 cm have a ;400-yr difference (older for
bulk organic). The dates of Potamogeton and bulk
organics from the sediments below 35 cm are thus
corrected for the hard-water effect by subtracting 600 yr
and 1000 yr, respectively. The final age–depth model
suggests that the lower part of the core has a higher
sedimentation rate than the upper part.
RESULTS
Pollen record
The pollen spectra from Co Ngion are dominated by
herbaceous pollen (91–99%) such as Poaceae, Artemisia,
Cyperaceae, Ranunculaceae, Thalictrum, Chenopodiaceae, Asteraceae, Caryophyllaceae, Fabaceae, and Polygonum. The arboreal pollen contains mainly grains of
Hippophae, Betula, Pinus, and Alnus along with isolated
grains of Abies, Picea, Quercus, Ulmus, Salix, Potentilla,
and Caragana. The occurrence of tree pollen is
apparently due to long-distance transport. The shrub
pollen such as Hippophae, Potentilla, and Caragana may
come from local vegetation. No aquatic pollen is present
in the pollen spectra. The pollen record from Co Ngion
is divided into six pollen zones (Fig. 3).
Zone CN-6 (110–91 cm; 6000–4900 cal. yr BP).—
Poaceae pollen is consistently and abundantly represented throughout this zone. Its percentages vary
between 23% and 29%. Artemisia pollen fluctuates
between 20% and 29% with two peaks and two troughs.
Cyperaceae pollen declines nonlinearly from 45.5% at
the base level to 30.9% at the top level. This zone has
relatively high pollen influx values.
Zone CN-5 (91–79 cm; 4900–4400 cal. yr BP).—
Cyperaceae pollen abruptly increases from 30.9% to
56%, whereas Poaceae and Artemisia pollen decreases
markedly. Pollen influx values decrease slightly.
Zone CN-4 (79–59 cm; 4400–3900 cal. yr BP).—
Cyperaceae pollen steadily decreases to its lowest
percentages (29.7%). Artemisia pollen increases first,
then Poaceae pollen dramatically increases to its highest
level (35%). Pollen influx values show a generally
increasing trend. There are no significant changes for
other pollen types.
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CAIMING SHEN ET AL.
Ecology, Vol. 89, No. 4
FIG. 4. Vegetation types reconstructed using discriminant functions, vegetation cover inferred from the content of minerogenic
matter, climate estimates derived from transfer functions (inverse linear [IL] and weighted-averaging partial least squares [WAPLS] models) over the last 6000 years, and a comparison with other proxy data in the southwest monsoon domain, including d18O
of stalagmite Q5 from Qunf cave in Southern Oman (a monsoon precipitation proxy record; Fleitman et al. 2003), and Globigerina
bulloides percentage (a proxy of Arabian Sea upwelling and monsoon strength calculated from an aliquot of ;300 specimens of
planktonic foraminifera in Hole 723A from the Arabian Sea; Gupta et al. 2003).
Zone CN-3 (59–31 cm; 3900–2800 cal. yr BP).—This
zone is marked by a steady decrease in the pollen influx
values and percentages of Artemisia and Poaceae.
Cyperaceae pollen increases at the expense of Artemisia
and Poaceae pollen. Other pollen types are still
represented at low quantities.
Zone CN-2 (31–15 cm; 2800–1600 cal. yr BP).—This
zone is characterized by a dramatic increase in Poaceae
pollen, and by a decrease in Cyperaceae pollen.
Artemisia pollen remains relatively high, but a marked
decline occurs in the middle of the zone. Pollen influx
values in this zone are significantly lower than that in
zone CN-3.
Zone CN-1 (15–0 cm; 1600 cal. yr BP to present).—
This zone is dominated by Cyperaceae pollen. It
accounts for about 50% of pollen sum, and reaches its
highest percentages. However, its influx values are lower
than that of other zones. Pollen percentages of both
Poaceae and Artemisia significantly decrease.
Meadow–steppe ecotone shift and climatic change
Co Ngion is a large lake located on the MSE. Its
pollen record thus reflects regional vegetation changes.
Fig. 4 shows the quantitative reconstruction of vegetation and climate based on this pollen record. The low
probabilities (,0.5) of modern analogue suggest that
most fossil samples do not have good modern analogues, indicating the transitional nature of ecotones.
The reconstructed vegetation reflects alternations be-
tween steppe and meadow by a MSE shift, as indicated
by the vegetation zonal index. Lacking good modern
analogues and a possible Cyperaceae pollen source
problem whereby a small amount of Cyperaceae pollen
is probably derived from marsh meadows growing in
lake edge habitat, may produce extra sources of error in
climate estimates. In addition, the surface pollen sample
gap in the vast high-elevation areas west of the ecotone
(Fig. 1a), where MAP is ,300 mm, may cause the
overestimate of MAP. Nevertheless, the estimates
derived from linear and unimodal models (IL and
WA-PLS) are consistent with each other. Modern values
of MAP and Tjuly are within the 95% confidence
intervals of reconstructed MAP and Tjuly of the surface
sample. These two facts indicate that the reconstruction
is still reliable for reflecting climatic variability despite
these problems.
The reconstructed vegetation and climate reveals a
detailed history of the MSE shift and climatic variations.
The MSE was east of Co Ngion during the interval from
6000 to 4900 cal. yr BP (pollen zone CN-6), when
regional vegetation was dominated by the steppe. The
reconstructed Tjuly from 6000 to 5400 cal. yr BP was
similar to that of today. It gradually rose to the highest
point at about 18C higher than today around ca. 5000
cal. yr BP. MAP decreased from 6000 to 4900 cal. yr BP,
causing a decline in effective moisture and lake level and
the formation of a peat layer in the lake. These facts
suggest the occurrence of a major centennial-scale
April 2008
ECOTONE SHIFT IN THE TIBETAN PLATEAU
FIG. 4.
Continued.
drought during 5600–4900 cal. yr BP. This period is
followed by a westward shift of the MSE at 4900–4400
cal. yr BP (pollen zone CN-5), resulting in the regional
vegetation changing from steppe to meadow. MAP
began to increase toward the present level, and remained
relatively stable between 4900 and 4400 cal. yr BP,
whereas Tjuly declined to 0.58C lower than present. After
4400 cal. yr BP, the MSE shifted eastward, resulting in
the establishment of a steppe in the region again. An
abrupt decline in MAP and a marked rise in Tjuly
occurred around 4200–4000 cal. yr BP, causing a
decrease of effective moisture and a remarkable drop
of lake level. Consequently, the shallower lake allowed
Potamogeton to grow at the coring site, eventually
forming a peat layer that marks the culmination of
another major drought. The vegetation changed several
times between steppe and meadow during the period of
3900–2800 cal. yr BP (pollen zone CN-3). However, no
typical steppe occurred, whereas typical meadow appeared only near the end of this period, indicating
frequent shifts of the ecotone near its present location.
During this period, inferred MAP increased to the
present value, and Tjuly fell and fluctuated around the
present value. Between 2800 and 2400 cal. yr BP, the
MSE moved eastward markedly, and regional vegetation changed from meadow to steppe. Tjuly only changed
slightly after 2800 cal. yr BP, whereas inferred MAP fell
abruptly to about 30 mm below the present level,
suggesting the occurrence of another major centennialscale drought. This drought is followed by a short wet
period from 2400 to 2200 cal. yr BP when the MSE
shifted westward. The MSE moved eastward again from
2200 to 1800 cal. yr BP, probably indicating another
1085
centennial drought, albeit of lower magnitude than
before. After 1800 cal. yr BP, the ecotone gradually
migrated to its modern position, and regional vegetation
was dominated by meadow. The MAP was gradually
rising to the present level, and Tjuly decreased slightly.
However, Tjuly dropped by ;0.88C during the interval of
700–300 cal. yr BP, which may indicate the occurrence
of the Little Ice Age cooling event in the central Tibetan
Plateau.
Generally an increase in minerogenic matter of
sediments may represent a lowering in lake level and/or
loss of vegetation cover, resulting from increased soil
erosion (Digerfeldt 1986, Clark et al. 2002, Serge et al.
2003). In our case, the coring site is located in the center
of Co Ngion, and the sedimentation rate of the lake is
very low (Fig. 2). Furthermore, the study region
experiences more than 100 days of strong wind
(.17m/s; Dai 1979). Therefore, the minerogenic sediments would have been mainly delivered to the lake by
wind, although there are 14 streams flowing into Co
Ngion. As the lake dried up, the mid-lake core site was
closer to the sources of minerogenic sediments. On the
other hand, loss of vegetation cover (increased soil
erosion) results in more minerogenic sediments transported into the lake by wind. Variations in the content
of minerogenic matter (Fig. 4) are consistent with our
interpretation of pollen data. Marked increases in
minerogenic matter occurred during the early stages or
at the onset of three major droughts, and were marked
during the two warm droughts. This fact indicates a
drop of lake level and/or decrease in vegetation cover.
Following the high influx of minerogenic matter, a peat
layer formed from aquatic macrophytes during the two
earlier (and warmer) droughts. This caused organic
matter to increase in the sediments, diluting and
decreasing the percentage of minerogenic matter. The
abundance of aquatic macrophytes suggests that the
lake reached its lowest level for each drought at that
time (Morrill et al. 2006). High levels of minerogenic
matter found over the last 1500 years are attributed to
the loss of vegetation cover, as suggested by low pollen
influx values. However, this decrease in vegetation cover
probably is not caused by the dry conditions but by the
decline in temperature. Overall, pollen data and
sediment evidence suggest that the MSE shifts in two
steps during a drought. First, the meadow components
decrease to cause a loss of vegetation cover, and then
steppe components dominate the regional vegetation.
DISCUSSION
AND
CONCLUSIONS
A number of studies have revealed the occurrence of
strong southwest summer monsoon during the early–
mid Holocene and a gradual weakening of monsoon
intensity in response to variation of summer insolation
(e.g., Overpeck et al. 1996, Neff et al. 2001, Fleitmann et
al. 2003, Gupta et al. 2003; Fig. 4). Our data show that
monsoon precipitation as inferred from reconstructed
MAP fluctuated around the present level over the last
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CAIMING SHEN ET AL.
Ecology, Vol. 89, No. 4
PLATE 1. An overall view of Co Ngion from the east. The flag in the photo indicates the coring location. Photo credit: C. Shen.
6000 years in the central Tibetan Plateau. This result is
supported by the water-balance history of Ahung Co
inferred from sediment geochemistry, particularly the
occurrence of a drought between 5600 and 4900 cal. yr
BP (Morrill et al. 2006). It is also consistent with multiproxy data from Sumxi Co in the northwestern Tibetan
Plateau, which suggest that summer monsoon influence
on this region ended around 6500 cal. yr BP (Van
Campo and Gasse 1993, Maxwell and Liu 2002).
Therefore, the environmental conditions similar to those
occurring today were probably established around 6500
cal. yr BP in the central and northwest part of the
Tibetan Plateau.
Our pollen data show that the MSE shifted several
times in response to the variations of monsoon
precipitation over the past 6000 years. The history of
the MSE shift indicates that the mid–late Holocene was
punctuated by three major centennial-scale drought
episodes in the central Tibetan Plateau. The major
droughts of 5600–4900 and 4400–3900 cal. yr BP
developed from a relatively dry and warm climate
background. The dry or cool events in these two
episodes are also found in the records from lake
sediments (e.g., Van Campo and Gasse 1993, Wei and
Gasse 1999, Morrill et al. 2006), peat (Zhou et al. 2001,
Xu et al. 2002), and ice cores (Thompson et al. 1989) in
the Tibetan Plateau. In the southwest monsoon domain,
the d18O isotopic record of stalagmite Q5 from Qunf
cave in Southern Oman shows marked decreases in
monsoon precipitation during these two episodes
(Fleitmann et al. 2003; Fig. 4). And the record of the
percentage contribution of Globigerina bulloides, a proxy
of Arabian Sea upwelling and monsoon strength in Hole
723A from the Arabian Sea indicates two intervals of
weak summer monsoon occurring at 6100–5300 and
4500–4200 cal. yr BP (Gupta et al. 2003; Fig. 4). Having
considered the radiocarbon measurement uncertainty in
both records, these two weak summer monsoon events
could be comparable with our identified drought
episodes, although the match is not perfect. Beyond
the southwest monsoon domain, evidence for dry or
cool events around 5500 cal. yr BP is well documented in
Africa, China, and other regions of the world (Guo et al.
2000, Magny and Haas 2004, Mayewski et al. 2004). The
first systematic look at abrupt century-scale variability
in the Asian monsoon since the last glacial period shows
that a monsoon weakening around 4200 cal. yr BP
appears to coincide with abrupt climate changes in
North Africa, the Middle East, the Indian subcontinent,
and China (Morrill et al. 2003). The coherent pattern
among different proxy data suggests that these two
major centennial-scale droughts are two in a series of
non-orbital monsoon weakening events during the
Holocene. Furthermore, these two events may have
been global in extent (Bond et al. 2001, Thompson et al.
2002, Mayewski et al. 2004, Booth et al. 2005). The
drought of 2800–2400 cal. yr BP is another major
monsoon weakening event. Its hints also are found in
other records in monsoon regions (e.g., Zhou et al. 2001,
Xu et al. 2002, Gupta et al. 2003, Mayewski et al. 2004),
although it is not as severe as other two major droughts
in the central Tibetan Plateau.
Hypothesized causes of monsoon droughts include
internal variability in the climate system such as landsurface albedo changes (Gasse et al. 1991, deMenocal et
al. 2000) and cold temperature anomalies in the North
Atlantic (Gupta et al. 2003), as well as external forcing
such as variations in solar activity (Neff et al. 2001,
April 2008
ECOTONE SHIFT IN THE TIBETAN PLATEAU
Fleitmann et al. 2003). Our results and those of previous
studies (e.g., Gasse et al. 1991) indicate significant
changes in vegetation cover, and possibly albedo, in the
central and northwestern Tibetan Plateau during these
major droughts. Since our identified events appear to
have spanned a large area, as mentioned above, local
albedo changes may not have been the primary or sole
driver of these events, but could have been one
important player. Gupta et al. (2005) suggest that the
correspondence between cold periods in the North
Atlantic and weaker monsoon conditions is not one of
causation, but due rather to solar variations affecting
each of these regions individually. In this case, the
monsoon–solar link can be explained by the direct
influence of solar forcing on the Intertropical Convergence Zone that controls the monsoonal precipitation
(Gupta et al. 2005). Alternatively, and at least for the
two earlier warmer drought periods, the influence of
solar variation on surface warming and the monsoon
may have been dominant (Shindell et al. 2006). The
three major centennial-scale droughts identified from the
Co Ngion record coincide with periods of weak solar
activity as indicated by the atmospheric 14C content
(Stuiver et al. 1998) and 10Be values from Greenland
(Finkel and Nishiizumi 1997). Therefore, we infer that
the major centennial-scale droughts may have been
triggered by variations in solar activity and amplified by
albedo changes.
ACKNOWLEDGMENTS
This research was supported by grants from the U.S.
National Science Foundation (NSF grants ATM-9410491 and
ATM-0081941), the Chinese National Science Foundation
(grants 49371068 and 49871078), and dissertation research
grants from the Geological Society of America (GSA),
Association of American Geographers (AAG), and the Robert
C. West Field Research Award (LSU Department of Geography and Anthropology).
LITERATURE CITED
Allen, C. D., and D. D. Breshears. 1998. Drought-induced shift
of a forest-woodland ecotone: rapid landscape response to
climate variation. Proceedings of the National Academy of
Sciences (USA) 95:14839–14842.
Baker, R. G., E. A. Bettis III, R. F. Denniston, L. A. Gonzalez,
L. E. Strickland, and J. R. Krieg. 2002. Holocene paleoenvironments in southeastern Minnesota: chasing the prairie–forest ecotone. Paleogeography, Paleoclimatology,
Paleoecology 177:103–122.
Birks, H. J. B. 1995. Quantitative palaeoenvironmental
reconstructions. Pages 161–254 in D. Maddy and J. S. Brew,
editors, Statistical modeling of quaternary science data.
Quaternary Research Association, Cambridge, UK.
Birks, H. J. B. 1998. Numerical tools in quantitative
paleolimnology: progress, potentialities, and problems. Journal of Paleolimnology 20:301–332.
Bond, G., B. Kromer, J. Beer, R. Muscheler, M. N. Evans, W.
Showers, S. Hoffmann, R. Lotti-Bond, I. Hajdas, and G.
Bonani. 2001. Persistent solar influence on North Atlantic
climate during the Holocene. Science 294:2130–2136.
Booth, R. K., S. T. Jackson, S. L. Forman, J. E. Kutzbach,
E. A. Bettis III, J. Kreig, and D. K. Wright. 2005. A severe
centennial-scale drought in mid-continental North America
1087
4200 years ago and apparent global linkage. Holocene 15:
321–328.
Clark, J. S., E. C. Grimm, J. J. Donovan, S. C. Fritz, D. R.
Engstrom, and J. E. Almendinger. 2002. Drought cycles and
landscape response to past aridity on prairies of the northern
Great Plains, USA. Ecology 83:595–601.
Crumley, C. L. 1993. Analyzing historic ecotonal shifts.
Ecological Applications 3:377–384.
Dai, J. 1979. Climate of the Tibetan Plateau. Meteorological
Press, Beijing, China.
Dean, W. W. 1974. Determination of carbonate and organic
matter in calcareous sediments and sedimentary rocks by loss
on ignition: comparison with other methods. Journal of
Sedimentary Petrology 44:242–248.
deMenocal, P., J. Ortiz, T. Guilderson, J. Adkins, M.
Sarnthein, L. Baker, and M. Yarusinsky. 2000. Abrupt onset
and termination of the African humid period: rapid climate
responses to gradual insolation forcing. Quaternary Science
Reviews 19:347–61.
di Castri, F., A. J. Hansen, and M. M. Holland, editors. 1988.
A new look at ecotone emerging international projects on
landscape boundaries. Biology International (Special Issue)
17:1–163.
Digerfeldt, G. 1986. Studies on past lake-level fluctuations.
Pages 127–143 in B. E. Berglund, editor. Handbook of
holocene palaeoecology and palaeohydrology. Wiley, Chichester, UK.
Faegri, K., and J. Iversen. 1975. Textbook of pollen analysis.
Macmillan, New York, New York, USA.
Finkel, R. C., and K. Nishiizumi. 1997. Beryllium 10
concentrations in the Greenland Ice Sheet Project 2 ice core
from 3–40 ka. Journal of Geophysical Research 102:26699–
26706.
Fleitmann, D. S., J. Burns, M. Mudelsee, U. Neff, J. Kramers,
A. Mangini, and A. Matter. 2003. Holocene forcing of the
Indian monsoon recorded in a stalagmite from southern
Oman. Science 300:1737–1739.
Fontes, J.-C., F. Melieres, E. Gibert, Q. Liu, and F. Gasse.
1993. Stable isotope and radiocarbon balances of two
Tibetan lakes (Sumxi Co, Longmu Co) from 13,000 B.P.
Quaternary Science Reviews 12:875–887.
Gasse, F., M. Arnold, J. C. Fontes, M. Fort, A. Huc, B. Y. Li,
Y. F. Li, Q. Liu, F. Melieres, E. Van Campo, F. B. Wang,
and Q. S. Zhang. 1991. A 13 000 year climate record from
western Tibet. Nature 353:742–745.
Gasse, F., and E. Van Campo. 1994. Abrupt post-glacial
climate events in West Asia and North Africa monsoon
domains. Earth and Planetary Science Letters 126:435–456.
Guo, Z., N. Petit-Maire, and S. Kropelin. 2000. Holocene noorbital climatic events in present-day arid areas of northern
Africa and China. Global and Planetary Change 26:97–103.
Gupta, A. K., D. M. Anderson, and J. T. Overpeck. 2003.
Abrupt changes in the Asian southwest monsoon during the
Holocene and their links to the North Atlantic Ocean.
Nature 421:354–356.
Gupta, A. K., M. Das, and D. M. Anderson. 2005. Solar
influence on the Indian summer monsoon during the
Holocene. Geophysical Research Letter 32:L17703. [doi: 10.
1029/2005GL022685]
Herzschuh, U., K. Winter, B. Wunnemann, and S. Li. 2006. A
general cooling trend on the central Tibetan Plateau
throughout the Holocene recorded by the Lake Zigetang
pollen spectra. Quaternary International 154–155:113–121.
Holland, M. M., and P. G. Risser. 1991. Introduction: the role
of landscape boundaries in the management and restoration
of changing environments. Pages 1–7 in M. M. Holland,
P. G. Risser, and R. J. Naiman, editors. Ecotones: the role of
landscape boundaries in the management and restoration of
changing environments. Chapman and Hall, New York, New
York, USA.
1088
CAIMING SHEN ET AL.
Kullman, L. 1995. Holocene tree-limit and climate history from
the Scandes mountains, Sweden. Ecology 76:2490–2502.
Kullman, L. 1998. Tree-limits and montane forests in the
Swedish Scandes: sensitive biomonitors of climate change
and variability. Ambio 27:312–321.
Liu, K. B. 1990. Holocene paleoecology of the boreal forest and
Great Lakes–St. Lawrence forest in northern Ontario.
Ecological Monographs 60:179–212.
Liu, K. B., and N. S.-N. Lam. 1985. Paleovegetational
reconstruction based on modern and fossil pollen data: an
application of discriminant analysis. Annals of the Association of American Geographers 75:115–130.
Magny, M., and J. N. Haas. 2004. A major widespread climatic
change around 5300 cal. yr BP at the time of the Alpine
iceman. Journal of Quaternary Science 19:423–430.
Maxwell, A. L., and K.-B. Liu. 2002. Late Quaternary pollen
and associated records from the monsoonal areas of
continental South and SE Asia. Pages 189–228 in P.
Kershaw, B. David, N. Tapper, D. Penny, and J. Brown,
editors. The environmental and cultural history and dynamics of the SE-Asian-Australia region. Advances in Geoecology 34. Catena Verlag, Reiskirchen, Germany.
Mayewski, P. A., et al. 2004. Holocene climate variability.
Quaternary Research 62:243–255.
Morrill, C., J. T. Overpeck, and J. E. Cole. 2003. A synthesis of
abrupt changes in the Asian summer monsoon since the last
deglaciation. Holocene 13:465–476.
Morrill, C., J. T. Overpeck, J. E. Cole, K. B. Liu, C. Shen, and
L. Tang. 2006. Holocene variations in the Asian monsoon
inferred from the geochemistry of lake sediments in central
Tibet. Quaternary Research 65:232–243.
Neff, U., S. J. Burns, A. Mangini, M. Mudelsee, S. Fleitmann,
and A. Matter. 2001. Strong coherence between solar
variability and the monsoon in Oman between 9 and 6 kyr
ago. Nature 411:290–293.
Neilson, R. P. 1993. Transient ecotone response to climatic
change: some conceptual modeling approaches. Ecological
Applications 3:385–395.
Overpeck, J. T., D. Anderson, S. Trumbore, and W. Prell. 1996.
The southwest Indian Monsoon over the last 18,000 years.
Climate Dynamics 12:213–225.
Overpeck, J., K. Liu, C. Morrill, J. Cole, C. Shen, D. Anderson,
and L. Tang. 2005. Holocene environmental change in the
Himalayan–Tibetan Plateau region: lake sediments and the
future. Pages 83–92 in U. M. Huber, H. K. M. Bugmann, and
M. A. Reasoner, editors. Global change and mountain
regions: an overview of current knowledge. Springer, Boston,
Massachusetts, USA.
Reasoner, M. A., and M. A. Jodry. 2000. Rapid response of
alpine timberline vegetation to the Younger Dryas climate
oscillation in the Colorado Rocky Mountains, USA.
Geology 22:439–442.
Serge, D. M., P. J. H. Richard, J. Guiot, J. de Beaulieu, and D.
Fortin. 2003. Postglacial climate in the St. Lawrence
lowlands, southern Québec: pollen and lake-level evidence.
Paleogeography, Paleoclimatology, Paleoecology 193:147–
159.
Shen, C. 2003. Millennial-scale variations and centennial-scale
events in the southwest Asian monsoon: pollen evidence from
Ecology, Vol. 89, No. 4
Tibet. Dissertation. Louisiana State University, Baton
Rouge, Louisiana, USA.
Shen, C., K.-B. Liu, L. Tang, and J. T. Overpeck. 2006.
Quantitative relationships between modern pollen rain and
climate in the Tibetan Plateau. Review of Palaeobotany and
Palynology 140:61–77.
Shindell, D. T., G. Faluvegi, R. L. Miller, G. A. Schmidt, J. E.
Hansen, and S. Sun. 2006. Solar and anthropogenic forcing
of tropical hydrology. Geophysical Research Letter 33:
L24706. [doi: 10.1029/2006GL027468]
Stockmarr, J. 1971. Tablets with spores used in absolute pollen
analysis. Pollen et Spores 13:615–621.
Stuiver, M., P. J. Reimer, E. Bard, J. W. Beck, G. S. Burr, K. A.
Hughen, B. Kromer, G. McCormac, J. Van der Plicht, and
M. Spurk. 1998. INTCAL98 radiocarbon age calibration,
24,000–0 cal BP. Radiocarbon 35:215–230.
Sun, X., N. Du, Y. Chen, Z. Gu, J. Liu, and B. Yuan. 1993.
Pollen analyses of lacustrine sediment in Seling Co, Tibet.
Acta Botanica Sinica 35:943–950.
ter Braak, C. J. F., and S. Juggins. 1993. Weighted averaging
partial least squares regression (WA-PLS): an improved
method for reconstructing environmental variables from
species assemblages. Hydrobiologia 269/270:485–502.
Thompson, L. G., E. Mosley-Thompson, M. E. Davis, J. F.
Bolzan, J. Dai, T. Yao, N. Gundestrup, X. Wu, L. Klein, and
Z. Xie. 1989. Holocene-Late Pleistocene climatic ice core
records from Qinghai–Tibetan Plateau. Science 246:474–477.
Thompson, L. G., E. Mosley-Thompson, M. E. Davis, K. A.
Henderson, H. H. Brecher, V. S. Zagorodnov, T. A.
Mashiotta, P.-N. Lin, V. N. Mikhalenko, D. R. Hardy,
and J. Beer. 2002. Kilimanjaro ice core records: evidence of
Holocene climate change in tropical Africa. Science 298:589–
593.
Tibetan Investigation Group. 1988. Vegetation of Xizang
(Tibet). Science Press, Beijing, China.
Van Campo, E., and F. Gasse. 1993. Pollen- and diatominferred climatic and hydrological changes in Sumxi Co Basin
(Western Tibet) since 13,000 yr B.P. Quaternary Research 39:
300–313.
Wang, J. 1988. The steppes and deserts of the Xizang Plateau
(Tibet). Vegetatio 75:135–142.
Wang, S., and H. Dou, editors. 1998. Lakes in China. Science
Press, Bejing, China.
Webb, T., III, and D. R. Clark. 1977. Calibrating micropaleontological data in climatic terms: a critical review. Annals of
the New York Academy of Science 288:93–118.
Wei, K., and F. Gasse. 1999. Oxygen isotopes in lacustrine
carbonates of West China revisited: Implications for post
glacial changes in summer monsoon circulation. Quaternary
Science Reviews 18:1315–1334.
Xu, H., Y. Hong, Q. Lin, B. Hong, H. Jiang, and Y. Zhu. 2002.
Temperature changes over the 6500 years indicated by
oxygen isotope of plant cellulose from peat in Hongyuan.
Chinese Science Bulletin 47:1181–1186.
Zhou, W., X. Luo, Z. Wu, L. Duan, A. J. T. Jull, D. Donahue,
and W. Beck. 2001. Peat record of Holocene climatic changes
in the Zoige Plateau and AMS radiocarbon dating. Chinese
Science Bulletin 46:1040–1044.