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Summer monsoon precipitation variations in central China over the past 750 years
derived from a high-resolution absolute-dated stalagmite
Liangcheng Tan
a
, Yanjun Cai
a,
⁎, Hai Cheng
b
, Zhisheng An
a
, R. Lawrence Edwards
b
a
State Key Laboratory of Loess and Quaternary Geology, Institute of Earth Environment, CAS, Xi'an, 710075, China
b
Department of Geology and Geophysics, University of Minnesota, Minneapolis, 55455, USA
abstractarticle info
Article history:
Received 7 January 2009
Received in revised form 15 June 2009
Accepted 21 June 2009
Available online 27 June 2009
Keywords:
Aragonite stalagmite
Central China
Asian monsoon precipitation
Little Ice Age
Regional differences
A2–3-year resolution record of stalagmite oxygen isotope variations from the south flank of the Qinling
Mountains, central China, has revealed the Asian summer monsoon (ASM) precipitation variations in the
investigated area over the past 750 years. The summer monsoon precipitation gradually increased since
1249 AD, reaching its highest values in the period 1535–1685 AD, and then decreased with substantial
decadal- to centennial-scale fluctuations. The monsoon precipitation increased again between 1920 and
1970 AD. Three intervals of high monsoon precipitation were identified: 1535–1685 AD, 1755–1835 AD, and
1920–1970 AD. Three intervals of low precipitation were inferred in 1249–1325 AD, 1390–1420 AD, and
189 0–1915 AD. The δ
18
O composition and lithological features of the stalagmite coincidently indicate a
wetter climate during the Little Ice Age (LIA), which is also confirmed by climate records from Chinese
historical documents within this area. A comparison with other high-resolution speleothem records indicates
regional differences in monsoon precipitation variability from the south to the north of central China in the
last 750 years on decadal- to centennial-scale. Power spectrum analysis of the δ
18
O record shows significant
117.8-, 34.6-, 14-, 10.3-, and ~ 6-year periodicities. These periodicities are widely observed in the climate
records from ASM-controlled areas of China and are consistent with the Gleissburg periodicity, Brϋckner
periodicity, sunspot periodicity of solar activity, and El Nińo–Southern Oscillation (ENSO) periodicity. These
correlations suggest that both solar activity and ENSO periodicity may have had important influences on ASM
precipitation in China over the past 750 years.
© 2009 Elsevier B.V. All rights reserved.
1. Introduction
Summer monsoon precipitation is vital to the livelihood and well
being of monsoon societies because it not only provides the vital
monsoon precipitation that supports agricultural practices, but also
creates periods of drought and flood throughout the region. Under-
standing the temporal and spatial variations in monsoon precipitation is
essentialto characterizethe present monsoon precipitationpatterns and
to predict future changes. It has been suggested that monsoon
precipitation does not correlate positively at each individual record
during the Holocene, although it was mainly controlled by the intensity
of the ASM (An et al., 2000). Investigation of the meteorological data
have alsodemonstrated a spatial variability in monsoon precipitation in
both China (Qian and Lin, 2005) and India (Gregory, 1989), implying
regional differences in precipitation responses to ASM variability.
Therefore, it is expected that monsoon precipitation, which is related
to ASM intensity, might show regional differences on decadal to
centennial time scales. However, instrumental records usually begin in
the 1950s (although a few were recorded before the 1900s; Bradley,
1999), which is not longenough for the detectionof spatial and temporal
patterns in China on decadal or centennial scales. Therefore, it is
important to obtain other high-resolution proxy records of monsoon
precipitation for the last millennium from different regions.
In China, published high-resolution precipitation records for the
last thousand years mainly relate to locations on and around the
Tibetan Plateau (Yao et al., 2000; Zhang et al., 2003; Liu et al., 2006; Xu
et al., 2007; Tan et al., 2008; Zhang et al., 2008), in Guizhou Province in
southern China (He et al., 2005; Wang et al., 2005), and in the mid
Yangtze River valley (Hu et al., 2008). These records provided us with
some information on the local hydrological conditions of the
investigated areas on decadal- to centennial-scale. For example, an
abnormally dry climate was recorded in both the historical data from
Longxi, on the northeast margin of the Tibetan Plateau (Tan et al.,
2008), and the speleothem records from the Dongge Cave (He et al.,
2005; Wang et al., 2005) and Wanxiang Cave (Zhang et al., 2008)in
Guizhou and Gansu Provinces of China, respectively, during the Little
Ice Age (LIA) (Lamb, 1965; Bradley and Jones, 1992). However, the
hydrological conditions in other areas during the last thousand years
remain unclear, not to mention the spatial and temporal variability of
ASM precipitation.
The Qinling Mountains, which lie in central China (Fig. 1), are not
only the watershed of the Yangtze River valley and the Yellow River
valley, but are also a key climatic boundary in China (Chen, 1983). They
Palaeogeography, Palaeoclimatology, Palaeoecology 280 (2009) 432–439
⁎Corresponding author. Tel.: +86 29 88323194; fax: +86 29 88320456.
E-mail address: caiyj@loess.llqg.ac.cn (Y. Cai).
0031-0182/$ –see front matter © 2009 Elsevier B.V. All rights reserved.
doi:10.1016/j.palaeo.2009.06.030
Contents lists available at ScienceDirect
Palaeogeography, Palaeoclimatology, Palaeoecology
journal homepage: www.elsevier.com/locate/palaeo
block the warm, humid airflow from the south in summer, and the cold,
dry airflow from the north in winter. The climate in these areas is
sensitive to the variability of the Asian monsoon system. Paulsen et al.
(2003) have reported a 1270-year-old stalagmite record from the
Buddha Cave in the southern Qinling Mountains, and identified a
significant 33-year climatic cycle over the last 1270 years. Other cycles
of 11 and 9.6 years were also found, suggestingthat external forces (e.g.,
solar irradiance) may affect the climate in this area (Paulsen et al.,
2003). Here, we present the oxygen isotope record of an absolute-dated
aragonite stalagmite, DY-1, collected from the Dayu Cave, 300 km
southwest of the Buddha Cave on the south flank of the Qinling
Mountains. The δ
18
O values of DY-1 correlate well with local
instrumental rainfall data during the overlapping time period, implying
that the δ
18
O data can be used as a proxy index for rainfall, especially
the summer monsoon precipitation, which accounts for more than 80%
of the annual precipitation. With the high-precision
230
Th chronology,
this δ
18
O record represents the history of the summer monsoon
precipitation variations in the investigated area over the past 750 years,
with an average temporal resolution of 2–3years.
2. Geographical and geological setting
The Dayu Cave (33°08′N,106°18′E) is located on the south flank of
the Qinling Mountains, 40 km north of Ningqiang County, Shaanxi
Province, central China (Fig. 1). The cave was formed in the Upper
Proterozoic dolomite, has a small entrance of about 2 ×3 m
2
, and its
ceiling rock is about 80 m thick. The elevation of the entrance is
~870 m above sea level (asl). The main passage is more than 2 km
long, with many branched passages and small chambers. The cave has
a high humidity (N95%) and abundant modern and fossil speleothems.
At present, this area is dominated by a monsoon climate. The mean
annual temperature and precipitation (1951–2000) recorded at the
nearest meteorological station, Hanzhong station (33°02′N, 107°01′E,
508 m asl), which is ~ 65 km from the Dayu Cave, are 14.3 °C and
890 mm, respectively. Most of the rainfall (80%) occurs during the
summer monsoon months (May–Oct.).
3. Stalagmite sample, chronology, and isotopic analysis
We collected the stalagmite DY-1 in the Dayu Cave ~1 km from the
cave entrance, in 2005. The columnar-shaped stalagmite is 10.9 cm in
length, and the diameter ranges between 4.5 and 6 cm (Fig. 2). The
sample was halved along the growth axis. When the surface was
polished, the stalagmite exhibited flat and clear growth bandings. Nine
subsamples, including two replicates, were drilled parallel to the growth
plane (Fig. 2) using a hand-held carbide dental drill, and were dated with
U-series methods at the University of Minnesota, USA. The chemical
procedures were similar to those described by Edwards et al. (1987) and
Cheng et al. (2000). The measurements were performed on a Finnigan
ELEMENT inductively coupled plasma mass spectrometer (ICP-MS),
which was equipped with a double focusing sector magnet and energy
filter in reversed Nier–Johnson geometry and a single MasCom mul tiplier,
following procedures modified from Shen et al. (2002).Thecorrections
for the initial
230
Th (
230
Th/
232
Th atomic ratio of 4.4 ± 2.2 × 10
−6
)were
Fig. 1. Locations of the Dayu Cave and other sites mentioned in the paper. The black filled square denotes the Dayu Cave. The directions of the East Asian summer monsoon and Indian
summer monsoon are also illustrated.
433L. Tan et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 280 (2009) 432–439
negligible because of the high
238
Uconcentrationsof12–21 ppm and the
low
232
Th levels of 0.3–1.5 ppb in the subsamples (Table 1). Most of the
dating errors were 1–3 yr, although when the drilling thickness was
considered, the average age errors increased to 3–8yr.
Because the polished section of DY-1 shows step wise chan ges in
lithology, linear interpolation was used to establish the chronology
(Fig. 3). The growth rates of DY-1 range from 0.077 to 0.188 mm/yr and
are consistent with the lithological changes. Fast-growing portions have
micropores and are white in color, whereas slow-growing portions are
gray or grayish yellow and the slowest-growing parts are dark gray in
color. Because the stalagmite was broken at the time of collection, we
cannot confirm the youngest age by active dripping, although it shows a
relatively fresh surface. Because the lithology of the upper 2.9 cm of DY-1
is homogeneous, we extrapolated the deposition rate of 0.188 mm/yr in
the section interval at 2.8–0.3 cm to the upper section, calculating an age
of 1983± 4 AD for the top part of DY-1 (the drilling thickness, the dating
errors, and the uncertainty in the slope of the linear fitintheagemodel
were considered). Because no visible hiatus was observed on the
polished surface or in the thin sections by microscopy, DY-1 is assumed
to have grown continuously from 1249 to 1983 AD.
Two sample bars (5 mm depth ×5 mm width × total length of the
sample) were sawn along the central growth axis of DY-1. To verify the
mineralogical composition, one of the two sample bars was divided into
six subsamples and analyzed for X-ray diffraction (XRD) on an X-Pert Pro
X-ray diffractometer at the Institute of Earth Environment, Chinese
Academy of Sciences (IEECAS). The XRD measurements indicated that
DY-1 is composed of aragonite with negligible calcite, and that the
mineralogical composition is homogeneous throughout the growth
period. Subsamples were scraped from the other sample bar parallel to
successive laminations, and 2190 subsamples were collected from the
upper 9.9 cm of DY-1. We performed stable isotope analysis on every 2–3
subsamples in the segment above 0.6 cm, every 10 subsamples between
0.6 and 7.9 cm, and every five subsamples in the segment below 7.9 cm.
A total of 306 oxygen and carbon isotopic values were thus obtained on a
Finnigan MAT-252 mass spectrometer equipped with a Kiel III Carbonate
Device at the IEECAS. We added one internal laboratory standard TTB1
every 15 samples for control. The replicates showed that the precision of
the δ
18
O analyses was better than 0.1‰(2σ). The average temporal
resolution of δ
18
O was ~ 0 .5 yr f rom 1950 to 198 3 AD an d abo ut 2 –4yr
from 1500 to 1249 AD. The δ
18
O record is shown in Fig. 4A.
4. Interpretation of the δ
18
O Record
Speleothem δ
18
O values can be interpreted in terms of climate only if
the system has remained closed from water/rock interactions and/or
kinetic processes. A rigorous test for isotopic equilibrium conditions is
the replication test (Hendy and Wilson, 1968; Dorale et al., 1998; Wang
et al., 2001). This involves the comparison of two or more speleothem
records from different caves or different locations within the same cave
that grew contemporarily. We compared the DY-1 δ
18
O record with the
δ
18
O record of a calcite stalagmite SF-1 from the Buddha Cave in the
southern Qinling Mountains, 300 km northeast of the Dayu Cave
(Paulsen et al., 2003). The δ
18
O variations in the two stalagmites are
similar in general, although with different mineral compositions (Fig. 5).
Fig. 2. Section of DY-1 from the Dayu Cave; the locations of the nine dating subsamples,
including two replicates, are shown.
Table 1
230
Th dating results of stalagmite DY-1 from the Dayu Cave.
Depth
(mm)
238
U
(ppb)
232
Th
(ppt)
230
Th/
232
Th
(ppm)
δ
234
U
(measured)
230
Th/
238
U
(activity)
230
Th age (yr)
(uncorrected)
230
Th age (yr BP)
(corrected )
230
Th age (yr AD)
(corrected )
δ
234
U
Initial
(corrected)
98 20378± 195 1364± 11 4228 ± 36 1551± 9 0.01714± 0.00017 735± 8 679± 8 1271± 8 1554± 9
96.75 15556± 23 1537± 21 2841± 42 1555± 3 0.01700± 0.00012 729± 5 672± 5 1278± 5 1558± 3
78.75 21893± 46 927± 16 4527 ± 81 1570±4 0.01161± 0.00 008 494 ± 3 438± 3 1512± 3 1572± 4
72.5 12398± 86 384 ± 11 5523 ± 158 1536 ± 6 0.01036± 0.00009 447 ± 4 391± 4 1559± 4 1538± 6
54 17096± 35 612±18 3716± 113 1546 ± 4 0.00806 ± 0.000 07 346 ± 3 290± 3 1660± 3 1547± 4
35.25 17660± 128 585± 10 2528± 46 1526± 7 0.00507 ± 0.00005 219± 2 163± 2 1787 ± 2 1526 ± 7
27.75 21618± 39 400 ± 21 3504 ±186 1526± 3 0.00393 ± 0.000 05 170± 2 114 ± 2 1836 ± 2 1527 ± 3
2.5 13190± 85 309 ± 10 592 ± 24 1538±6 0.00084±0.0 0002 36 ± 1 −20±1 1970 ± 1 1538 ± 6
2.5 12429± 20 3247± 24 59 ± 4 1547±3 0.00093±0.0 0006 40 ± 2 −19±3 1969±3 1548±3
The errors are 2σerrors. Decay constant values are: λ
230
=9.1577×10
−6
y
−1
,λ
234
= 2.8263 ×10
−6
y
−1
(Cheng et al., 2000) and λ
238
=1.55125× 10
−10
y
−1
(Jaffey et al., 1971).
Corrected
230
Th ages assume the initial
230
Th/
232
Th atomic ratio of 4.4 ±2.2 ×10
−6
. Depths along the growth axis are relative to the top (youngestsurface) of the stalagmite. Year BP:
year before present (1950 AD).
Fig. 3. Plot of age verses depth for DY-1 from the Dayu Cave. Error bars indicate
230
Th
dates with 2σerrors.
434 L. Tan et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 280 (2009) 432–439
The discrepancy observed ona decadal scale may be the result of dating
errors and the different depth-age models used. The chronology of SF-1
is based on an average growth rate of0.083 mm/yr for the last 150 years
and an average growth rate of 0.0163 mm/yr before that (Paulsen et al.,
2003). The former was determined by counting the annual laminations
and
210
Pb dating, and the latter was determined from TIMS
230
Th dating
at 3100 years with a dating error of 50 years (Li et al., 2000; Paulsen
et al., 2003).
The replication test indi cated that kinetic fractionation was negligible
and that the δ
18
O signal of DY-1 is primarily of climatic origin (Dorale et
al., 1998; McDermott, 2004). Therefore, the δ
18
O of the aragonite
stalagmite was controlled simultaneously by the isotope composition of
the drip water and the temperature inside the cave (Hendy, 1971).
Although temperature is important and changes of 1–2°C (Ge et al.,
2003) in central China over the last 750 years could lead to changes of
0.25–0.5‰in our δ
18
O record (oxygen isotope fractionation of aragonite-
water is 10
3
lnα=20.44 ×10
3
/T-41.48; Zhou and Zheng, 2003), the
primary factor controlling our stalagmite δ
18
O was the changes in δ
18
Oin
the meteorological precipitation. Temperature-dependent changes are
probably masked by the variation in δ
18
O of meteorological precipitation.
Alternately, the low temperatures in the cold LIA (Yang et al., 2002; Ge
et al., 2003) would lead to a more positive δ
18
O, which is opposite the
changes in our record (Fig. 4A).
In general, the “amount effect”dominates the δ
18
O of precipitation
in Asian monsoon-affected regions (Araguás-Araguás et al., 1998;
Cheng et al., 2005), even in the northern limit of the summer
monsoon region. This was confirmed by the analysis of Johnson and
Ingram (2004), who suggested that the temperature dependence of
δ
18
O on precipitation near the northern limit of the summer monsoon
region is not likely to be more than about 0.24‰/°C, so temperature
would have a negligible effect on the DY-1 δ
18
O when the temperature
dependence of δ
18
O fractionation in aragonite-water is excluded.
Because the precipitation in the study area mainly comes from the
summer monsoon, and the δ
18
O of the summer monsoon precipitation
is distinctly more negative than that of the winter precipitation
(Cheng et al., 2005), we interpreted the δ
18
O variations in DY-1 as
reflecting changes in the ASM precipitation. More (less) summer
monsoon precipitation in this area will result in more negative
(positive) δ
18
O of amount weighted annual means, and therefore
more negative (positive) δ
18
O in the stalagmite.
We also compared the DY-1 δ
18
O record (to avoid possible
contamination, the first subsample from the top was discarded) with
the meteorological temperature and rainfall records of Hanzhong
station during the period of data overlap (1951–1983 AD). There is a
significant negative correlation (R=−0.385, N=33, Pb0.05) between
the δ
18
O of DY-1 and the annual rainfall. The correlation coefficient was
not very high, which may be ascribed to two factors. The first is the
possible “smoothing effect”of the δ
18
O in drip waters induced by the
intra/interannual mixture of “fresh water”and “old water”in the karst
aquifer. The other factor is the age error. Although the lithology of the
upper 2.9 cm of DY-1 is homogeneous, there are still small differences.
The use of an average deposition rate may cause different age errors in
given subsamples. Nevertheless, the long-term changes in the δ
18
O
record are consistent with the instrumental rainfall, with lighter δ
18
Oin
the stalagmite corresponding to higher precipitation and vice versa. In
contrast, no significant correlation exists between the δ
18
OofDY-1and
the annual temperature (Fig. 6). This comparison confirms the climatic
signal of the δ
18
OrecordofDY-1.
Fig. 4. (A) δ
18
O record of stalagmite DY-1 from the Dayu Cave. The red line is the five-
point running mean. The light gray lines indicate the average δ
18
O values for the whole
series (−7.15‰), and for the periods 1249–1510 AD (−7.0 0 ‰) and 1510–1983 AD
(−7.21 ‰). The yellow bars indicate the three intervals of inferred elevated summer
monsoon precipitation and the cyan bars indicate the three intervals of low summer
monsoon precipitation. (B) Comparison between δ
18
O of DY-1 (green line) and the
annual precipitation record in Seoul since 1780 AD (five-point running mean, wine line;
Jung et al., 2001). (For interpretation of the references to color in this figure legend, the
reader is referred to the web version of this article.)
Fig. 5. Comparison of the speleothem δ
18
O records from the Dayu Cave, Buddha Cave
(Paulsen et al., 2003), Wanxiang Cave (Zhang et al., 2008), Dongge Cave (Wang et al.,
2005), Heshang Cave (Hu et al., 2008), and Lianhua Cave (Cosford et al., 2008).
435L. Tan et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 280 (2009) 432–439
Compared with many other long-timescale speleothem records from
the Asian monsoon-dominated region (e.g. Cai et al., 2006; Wang et al.,
2008), the variations in the δ
18
O series of DY-1 are small, about 1‰.Two
factors may be responsible. The firstisthatthechangesinclimatewere
insignificant in monsoonal China over the past 750 years, so that the
climate-induced δ
18
O variations in the stalagmite are correspondingly
small. For example, the amplitude of the δ
18
O variations in stalagmite
WX42B from the Wanxiang Cave, China (33°19′N, 105°00′E), during
the past 800 years are about 1‰(Zhang et al., 2008). The δ
18
O variations
in stalagmite D15 from the Dongge Cave, China (25°17′N, 108°50′E),
over the past 800 years are about 1.2‰(He et al., 2005), and in another
two stalagmites, DA (Wang et al., 2005) and D4 (Dykoski et al., 2005),
they are about 1.5‰and 1.3‰, respectively. The other possible factor is
the influence of the epikarst zone. The thick cave roof of the Dayu Cave
(about 80 m) and possibly poor interconnectivity and permeability of
the microfissurenetworkabovestalagmiteDY-1mayhaveincreasedthe
residence time of the seepage water in the karst aquifer. This may have
induced a smoothing effect on the δ
18
O variations in the seepage water
(Genty and Deflandre, 1998; Luo and Wang, 2008), and may have
further resulted in correspondingly small variations in δ
18
OinDY-1.
However, the good correlation between the δ
18
O variability in DY-1 and
SF-1 suggests that the smoothing effect should not exceed a decade. The
significant correlation between the δ
18
O record of DY-1 and the
meteorological rainfall (Fig. 6) also supports this conclusion.
5. Comparison and discussion
5.1. Comparison of δ
18
O, lithology, and deposition rates of DY-1
The δ
18
O record of DY-1 shows two obvious stages. The first stage is
from 1249 to 1510 AD, in which the δ
18
O was relatively heavier, with
an average value of −7.0 0 ‰, indicating less precipitation during this
period. The second stage was from 1510 to 1983 AD, with an average
δ
18
O value of −7.21‰, which indicates more precipitation. From 1249
to 1550 AD, δ
18
O continued to decline, with a total fluctuation of about
0.9‰, indicating a steady increase in monsoon precipitation (Fig. 4A).
The variations in δ
18
OobservedinDY-1coincidewellwiththe
variations in its lithology and deposition rates. Heavier δ
18
O values in
the first stage correspond to lower deposition rates, whereas lighter
δ
18
O values in the second stage correspond to higher deposition rates.
The deposition rate reached 0.183 mm/yr from 1535 to 1685 AD, during
which the average δ
18
O had the lightest value of the whole series
(−7.3 2 ‰). The increased deposition rate of DY-1 may be ascribed to the
increased drip water supply, which resulted from the abundant rainfall
in this region. A correlation between the stalagmite deposition rate and
the local precipitation has been reported at many sites around theworld
(e.g. Baker et al.,1998; Burnset al., 2002; Wu et al., 2006). However, the
deposition rates of stalagmites are not only controlled by the drip rate,
but are also affected by temperature, seasonality of the surface climate,
the productivity of soil CO
2
above the cave, and the Ca
2+
content of the
drip water (Baker et al., 1998). For example, the fastest deposition rate
of DY-1 is 0.188 mm/yr (Fig. 3), but the corresponding precipitation
inferred from δ
18
O is not the highest. Consequently, factors such as
temperature and soil CO
2
productivity, together with precipitation, may
have contributed to the fastest deposition rate of this period.
5.2. Comparison of δ
18
O of DY-1 and instrumental precipitation records
in Seoul since 1780 AD
Here we compare our record with a 220-year instrumental
precipitation record (1777–1996 AD) from Seoul, Korea (Jung et al.,
2001). It may seem unrealistic to compare our record with the
precipitation record in Seoul, Korea, the longest instrumental record
in East Asia, which is more than a thousand kilometers away from the
study cave. Also, the discrepancies between meteorological records
from Hanzhong and Seoul are significant on yearly- to decadal-scale
during the last 50 years. However, it is well known that the annual
precipitation in the mid-low Yangtze River region of China and Korean
peninsula is strongly affected by the East Asian summer monsoon
(EASM) (Gao et al., 1962; Ramage, 1971; Kang et al., 1999). Monsoon
precipitation reconstructed from tree-ring and Chinese historical
documents in Baotou, Inner Mongolia show synchronous variations
with that in South Korea on a multidecadal-scale during the last
160 years (Liu et al., 2003). The analysis of a daily rainfall dataset
based on weather stations from China from 1961 to 2000 shows
similar trends of summer precipitation in Hanzhong and Baotou in the
last 40 years (Fig. 12 in Qian and Lin, 2005). Therefore, it is not
implausible that the monsoon precipitation could correlate with that
of Seoul, Korea on a multidecadal-scale.
As shown in Fig. 4B, the δ
18
O result of DY-1 correlates well with the
precipitation in Seoul since 1780 AD on multidecadal- to centennial-
scale. A remarkably dry period from 1890 to 1910 AD was recorded in
both the δ
18
O of DY-1 in central China and the instrumental
precipitation record in Seoul. Other dry intervals around 1860 AD,
1945 AD, and1980 ADalso appear in both records. Dry intervals around
1900 AD, 1945 AD, and 1980 AD were simultaneously recorded in the
δ
18
O results of the annual-layer stalagmite HS6 in the Heshang Cave in
the mid Yangtze River valley (Hu et al., 2008). The consistency between
the δ
18
O of DY-1 and the precipitation in Seoul suggests that the climate
in the Dayu Cave is strongly affected by the EASM, and that the δ
18
Oof
DY-1 has recorded the summer monsoon precipitation variations over
the past 750 years in the study region, further supporting the
conclusion that monsoon precipitation in Korea and regions in northern
and central China changed synchronously on a multidecadal- to
centennial-scale.
5.3. Summer monsoon precipitation variations recorded by δ
18
OofDY-1
The summer monsoon precipitation in the investigated area
gradually increased from 1249 AD to 1530 AD, then rose dramatically
and maintained high values till 1685 AD. After 1685 AD, the monsoon
precipitation gradually decreased, but with substantial decadal- to
centennial-scale fluctuations. It increased again from 1920 to 1970 AD.
There have been three periods of high precipitation in this area over
the past 750 years: 1535–1685 AD, 1755–1835 AD, and 1920–1970 AD.
Three periods of low precipitation were identified in 1249–1325 AD,
1390–1420 AD, and 1890–1915 AD. Our results indicate that summer
monsoon precipitation during the LIA increased in the study area. The
Fig. 6. Comparison of the δ
18
O record of DY-1 and meteorological temperature and
rainfall records from the Hanzhong station for the period 1951–1983 AD. The black line
in the upper panel represents the annual temperature record and in the lower panel
represents the annual rainfall record. The green lines in both panels represent the δ
18
O
record of DY-1. (For interpretation of the references to color in this figure legend, the
reader is referred to the web version of this article.)
436 L. Tan et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 280 (2009) 432–439
wet conditions in the LIA in this area have also been recorded in
Chinese historical documents (CAMS, 1981). The historical records
chronicled that there were severe floods or prolonged heavy rains in
the Hanzhong area for a total 30 years between 1500 and 1850 AD,
and 86 years were mildly wet. In contrast, severe droughts occurred in
only 11 years and moderate droughts in 48 years. The remaining
176 years were recorded as experiencing a normal climate (CAMS,
1981). For example, “heavy rain lasted for 60 days”was recorded in
the Hanzhong area in 1662 AD, and “sustained rain lasting for 40 days,
a big flood destroyed many houses”in 1677 AD (Wang et al., 2002).
Power spectrum analysis (Schulz and Mudelsee, 2002)oftheδ
18
O
record has revealed significant periodicities of 117.8, 34.6, 14, 10.3, and
~6 years (Fig. 7). These periodicities have been widely observed in the
climate records of ASM-controlled areas (Hughes et al., 1994; Ku and Li,
1998; Bradley, 1999;Wang and Sarnthein, 1999;Hong et al., 2000; Liu
et al., 2001, 2003; Paulsen et al., 2003; Xu et al., 2006), suggesting that
there are important periodicities in the natural variability of ASM. For
example, the 117.8-year periodicity is consistent with the quasi-100-year
periodicity exhibited in both the δ
18
O records of peat cellulose from
Hongyuan (Xu et al., 2006) and Jinchuan (Hong et al., 2000)andthe
δ
18
O records of foraminifera from the South China Sea (Wang and
Sarnthein, 1999). The approximately 35-year periodicity has been
recorded in stalagmites from the Shihua Cave in northeast China (Ku
and Li, 1998) and the Buddha Cave in central China (Paulsen et al., 2003),
in tree rings from Baotou in north-central China (Liu et al., 2001), and in
historical data for the eastern Yangtze River Basin and southwest China
(Bradley, 1999) and at the northeast margin of the Tibetan Plateau (Ta n,
2008). The 14- (double seven-year periodicity) and ~6-year periodicities
correspond to the approximately seven-year periodicity found in tree
rings (Hughes et al., 1994) and stalagmite (Paulsen et al., 2003)fromthe
Qinling Mountains, and may be linked to the El Nińo–Southern
Oscillation (ENSO). The quasi-100-year, approximately 35-year, and
10.3-year periodicities may be linked to the Gleissburg periodicity (Xu
et al., 2006, and references therein), the Brϋckner periodicity (Ku and Li,
1998, and references therein), and the sunspot periodicity (Stuiver and
Braziunas, 1993; Grootes and Stuiver, 1997) of solar activity, respectively.
This analysis suggests that both solar activity and ENSO may have had
important influences on summer monsoon precipitation in China over
the past 750 years.
5.4. Regional differences in precipitation responses to ASM variability
The analysis of a daily rainfall dataset from China from 1961 to
2000 shows a zonal pattern of precipitation variations from the south
to the north in eastern China. Decreasing trends in annual precipita-
tion and summer precipitation were observed in the south China
region and the mid-low Yellow River valley. In contrast, an increasing
trend was seen in the mid-low Yangtze River valley (refer to Fig. 12 in
Qian and Lin, 2005). Did this spatial pattern of recent precipitation
variability in China also exist during the last 750 years? We compared
our record with other high-resolution speleothem records from the
Wanxiang Cave (Zhang et al., 2008), Dongge Cave (Wang et al., 2005),
Heshang Cave (Hu et al., 2008), and Lianhua Cave (Cosford et al.,
2008), as shown in Fig. 5.
The comparison shows that the spatial variability in monsoon
precipitation in central China during the last 750 years was more
complicated than that in eastern China in recent decades. The monsoon
precipitation variations revealed by the stalagmite δ
18
O from the Dayu
Cave are consistent with those from the Buddha Cave (Paulsen et al.,
2003), as described in Section 4. Similar variations in monsoon
precipitation can also be seen in the areas of the Dayu Cave and
Heshang Cave on decadal- to centennial-scale during the well-dated
period (Hu et al., 2008). In contrast, the revealed precipitation
variations from Dayu Cave are generally antiphased with those from
the Wanxiang (Zhanget al., 2008) and Dongge Cave (Wang et al., 2005),
both on the decadal- to centennial-scale and in the long-term trends
during the last 750 years. The Wanxiang and Dongge Cave δ
18
O records
show decreasing trends in monsoon precipitation from 1250 AD to
1600 AD, and then gradually increasing trends, with a remarkably dry
LIA. The precipitation in the Longxi area, reconstructed from Chinese
historical documents, also shows a decreasing trend from 960 to
1700 AD, and an increasing trend after 1700 AD, with the driest period
in the LIA (Tan et al., 2008). Considering the small distance of about
120 km from the Dayu Cave to Wanxiang Cave and the significant
positive correlation (R=0.416, N=53, Pb0.01) between the annual
precipitation above the two caves in recent decades, it seems
implausible that the monsoon precipitation variations in the two
areas were mutually antiphase during the last 750 years. However, such
antiphase relationships in regional monsoon precipitation variability
can also be seen in other closely situated areas. For example, the
monsoon precipitation recorded in a stalagmite from the Heshang Cave
shows obvious antiphase variations with a stalagmite from the Lianhua
Cave (about 130 km to the southwest) in both long-term trends and on
the decadal- to centennial-scale (Fig. 5).
This comparison indicates obvious regional differences in the
precipitation responses to ASM variability from the south to the north
in central China during the last 750 years. However, as high-resolution
precipitation records are still sparse in other areas of China, especially
in eastern and southwestern China, it is hard to fully determine the
spatial and temporal patterns of monsoon precipitation in China in the
late Holocene. Such regional differences in monsoon precipitation in
central China on decadal- to centennial-scale may have been caused
by the complexity of the topography (e.g., plateau, basin, mountain,
and river valley) and the atmospheric circulation here (e.g., Indian
summer monsoon, East Asian summer monsoon, East Asian winter
monsoon, and Westerlies). How does regional precipitation vary with
the summer monsoon changes on decadal- to centennial-scale? What
is the underlying climate-forcing mechanism? These are still open
questions.
6. Conclusions
With precise uranium dating, the high-resolution δ
18
O record of
the aragonite stalagmite DY-1, which was collected from the Dayu
Cave on the south flank of the Qinling Mountains in central China, was
used to infer variations in the summer monsoon precipitation in the
investigated area over the past 750 years. Periods with increased
summer monsoon precipitation are characterized by depleted
stalagmite δ
18
O, attributed to a strong “amount effect”. The δ
18
O
values for DY-1 could correlate with local rainfall data during the
overlapping time period, suggesting that the δ
18
O of our aragonite
stalagmite can be used as a proxy index for the amount of summer
monsoon precipitation. The DY-1 record demonstrates that the
summer monsoon precipitation gradually increased from 1249 AD to
1530 AD, then rose dramatically and was maintained at high levels till
1685 AD. After 1685 AD, the monsoon precipitation gradually
Fig. 7. Power spectrum analysisof the δ
18
O record of DY-1. Power spectrum analysis was
performed using Redfit35(Schulz and Mudelsee, 2002). The parameters of the
software used in this study were: nsim = 1000, mctest= T, rhopre= −99.0, ofac = 2,
n50 =4, iwin = 1 (see Schulz and Mudelsee (2002) for details).
437L. Tan et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 280 (2009) 432–439
declined, with substantial decadal- to centennial-scale fluctuations,
until another rise occurred between 1920 and 1970 AD. The monsoon
precipitation variations recorded by speleothem δ
18
O in different
caves demonstrate obvious regional differences from south to north in
central China during the last 750 years. Power spectrum analysis of
the δ
18
O record of DY-1 showed significant peaks with periodicities
reflecting solar activity and ENSO. These periodicities have also been
observed in other records from ASM-controlled areas of China,
suggesting that both may affect the summer monsoon precipitation
in China.
Acknowledgments
We thank Prof. Zhongping Lai, Dr. Kathleen Johnson and an
anonymous reviewer for their constructive suggestions. This work was
supported by the National Science Foundation of China grants
40531003, 40773009 and 40403001; the National Basic Research
Program of China grant 2004CB720206; U.S. National Science
Foundation grants 0502535; and Gary Comer Science and Education
Foundation grant CC8.
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