Introduction
The 4.2 ka BP event was a pronounced climate event in the Holocene that
has been widely studied in the past 20 years. It was identified as an abrupt
(mega)drought and/or cooling event in a variety of natural archives including
ice cores, speleothems, lake sediments, marine sediments and loess. This
climate episode was associated with the collapse of several ancient
civilizations and human migrations in many sites worldwide (e.g., Egypt,
Greece, the Indus Valley and the Yangtze Valley) (Weiss et al., 1993; Cullen
et al., 2000; Gasse, 2000; DeMenocal, 2001; Weiss and Bradley, 2001; Thompson
et al., 2002; Booth et al., 2005; Bar-Matthews and Ayalon, 2011; Berkelhammer
et al., 2012; Ruan et al., 2016; Railsback et al., 2018). Recently, the
4.2 ka BP event was defined as the lower boundary of the Meghalayan stage
by the International Commission on Quaternary Stratigraphy. The timing of
this geological boundary was defined at a specific level (i.e., the
transformation from calcite to aragonite accompanied by an abrupt increase in
δ18O) in a stalagmite from northeast India (Walker et al.,
2018).
Location of Shennong Cave (SN, black star) and other caves mentioned
in the paper. Panel (a) is an overview topographic map and Jiangxi
Province is framed by the green line. Red triangles show the locations of
published stalagmite records. NH: Nuanhe Cave (Tan, 2005), LH: Lianhua Cave
(Dong et al., 2015), JX: Jiuxian Cave (Cai et al., 2010), XL: Xianglong Cave
(Tan et al., 2018a), SB: Sanbao Cave (Dong et al., 2010), HS: Heshang Cave (Hu
et al., 2008), EM: E'mei Cave (Zhang et al., 2018), XS: Xiangshui Cave (Zhang
et al., 2004), DG: Dongge Cave (Wang et al., 2005), ML: Mawmluh Cave
(Berkelhammer et al., 2012). Black arrows denote the directions of the East Asian
summer monsoon (EASM), Indian summer monsoon (ISM) and westerlies, which
affect the climate in China. The green dashed line indicates the limit of the modern
East Asian summer monsoon. Panel (b) is an enlarged map showing the locations
of Shennong Cave, the Guixi meteorological station (GX) and the GNIP station
in Changsha (CS). The base map is the Natural Earth physical map at 1.24 km
per pixel for the world (data source: US National Park Service,
http://goto.arcgisonline.com/maps/World_Physical_Map; last
access: 11 February 2017). For interpretation of
the references to colors in this figure legend, the reader is referred to the
web version of this article.
The abrupt climate change associated with the 4.2 ka BP event has been
proposed to have contributed to the collapses of Neolithic cultures in China
(Jin and Liu, 2002; Huang et al., 2010, 2011; Zhang et al., 2010; Liu and
Feng, 2012; Wu et al., 2017). Most of these studies imply a temperature drop
in continental China at about 4.2 ka BP (Yao and Thompson, 1992; Jin and
Liu, 2002; Zhou et al., 2002; Zhong et al., 2017; Xu et al., 2006; Yao et al., 2017), but
changes in the spatial distribution of precipitation are also discussed (Tan
et al., 2008, 2018a, b; Huang et al., 2010, 2011; Wu et al.,
2017). For example, a grain-size record from Daihai Lake, north China,
suggests a decrease in monsoon precipitation between 4.4 and 3.1 ka BP with
a very dry interval between 4.4 and 4.2 ka BP (Peng et al., 2005). Extreme
flooding during the 4.2 ka BP event was identified by paleoflood deposits
in the middle reaches of the Yellow River (Huang et al., 2010, 2011). Wu et
al. (2017) reported evidence of two extraordinary paleoflood events in the
middle reaches of the Yangtze River at 4.9–4.6 and 4.1–3.8 ka BP, closely
related to the expansion of the Jianghan lakes. These extreme hydroclimate
events may have accelerated the collapse of the Shijiahe culture in the
middle reaches of the Yangtze River (Wu et al., 2017). Multiple proxies in
four stalagmites from Xianglong Cave, south of the Qinling Mountains,
indicate that the upper Hanjiang River region experienced a wet climate
during the 4.2 ka BP event (Tan et al., 2018a). Peat records provide a
broad picture of climate variations in southeast China (SEC) during the
Holocene (Zhou et al., 2004; Zhong et al., 2010a, b, c, 2015, 2017). The
resolution of these records, however, is not high enough to study the
detailed structure of the 4.2 ka BP event. So far, there has only been one
published stalagmite record from SEC (Xiangshui Cave; Fig. 1), indicating a
wet interval during the 4.2 ka BP event (Zhang et al., 2004).
Mean monthly temperature, precipitation and δ18O
value from two meteorological stations close to the study area and
environmental monitoring in Shennong Cave. (a) Mean monthly air
temperature (red line) and precipitation (black column) from the Guixi
meteorological station for 1951–2010. (b) Mean monthly air
temperature (red line), precipitation (black column) and δ18O
value (blue line) from the Changsha GNIP station for 1988–1992.
(c) Air temperature (red line) and relative humidity (blue line) in
Shennong Cave from October 2011 to April 2013.
The aim of this work was to obtain a high-resolution stalagmite-based record
from SEC to study the hydroclimatic variations during the 4.2 ka BP event
and to compare them to records in northern China in order to explore the
possible north–south precipitation gradient during this event.
Polished section (left) and age model (right) of stalagmite SN17.
Sampling positions for XRD analyses (red lines; A aragonite, C calcite) and
230Th datings (black lines) are shown on the slab. SN17 age model
and modeled age uncertainties using COPRA; error bars on 230Th
dates indicate 2σ analytical errors. The gray band depicts the
95 % confidence interval.
Study area and sample
Shennong Cave (28∘42′ N, 117∘15′ E; 383 m a.s.l.) is
located in the northeastern Jiangxi Province, SEC (Fig. 1), a mid-subtropical
region strongly influenced by the East Asian summer monsoon (EASM). Mean
annual precipitation and temperature at the nearest meteorological station
(Guixi station; 1951–2010 CE) are 1857 mm and 18.5 ∘C,
respectively (Fig. 2a). Shennong Cave is located in a region of spring persistent
rain. The rainy season includes both summertime monsoon rainfall
and spring persistent rain (Tian and Yasunari, 1998; Wan et al., 2008; Zhang
et al., 2018). The latter is a unique synoptic and climatic phenomenon that
occurs from March until mid-May, mostly south of the Yangtze River (about 24
to 30∘ N, 110 to 120∘ E; Tian and Yasunari, 1998; Wan and
Wu, 2007, 2009). EASM precipitation lasts from mid-May to September (Wang and
Lin, 2002). In the region of spring persistent rain, the EASM (May to
September) precipitation accounts for 54 % of the annual precipitation
and the non-summer monsoon (NSM, October to next April) precipitation
accounts for 46 % (Zhang et al., 2018). The distribution of the EASM vs. NSM
precipitation amount in this region is distinctly different from that in the
northern and southwestern part of monsoonal China, where the mean annual
percentage of EASM (65 %–90 %) is much higher than the mean annual
percentage of NSM (10 %–35 %). Data from the nearest GNIP station in
Changsha, also located in the region of spring persistent rain, indicate
that the δ18O values of EASM precipitation are lower compared
with those of NSM precipitation (Fig. 2b). A full 2 years of monitoring data
(2011–2013) in Shennong Cave indicate that the speleothem
δ18O values reflected drip water δ18O values
inherited from the amount-weighted annual precipitation δ18O
outside the cave. Therefore, different from the southwestern and northern
part of monsoonal China where the speleothem δ18O values
are mainly influenced by EASM precipitation, speleothem δ18O
from Shennong Cave is controlled by both EASM and NSM precipitation
(Zhang et al., 2018).
The cave developed in the carboniferous limestone of the Chuan-shan and
Huang-long groups, which are mainly composed of limestone and interbedded
dolostone. The thickness of the cave roof ranges from about 20 to about
80 m, with an average of ∼50 m. The overlying vegetation consists
mainly of secondary forest tree species such as Pinus,
Cunninghamia and Phyllostachys, and shrub-like
Camellia oleifera and Ilex, which are C3 plants (Zhang
et al., 2015). A full 2 years of monitoring data show that the mean temperature in
the cave is 19.1 ∘C with a standard deviation of 2.5 ∘C
(Fig. 2c), consistent with mean annual air temperature outside the cave
(Fig. 2a). The relative humidity in the interior of the cave approaches
100 % during most of the year (Fig. 2c). Abundant aragonite and
calcite speleothems are present in the cave. Their mineralogy is likely
controlled by the Mg/Ca ratio of the drip water, reflecting the
variable dolomite content of the host rock (De Choudens-Sanchez and Gonzalez,
2009; Zhang et al., 2014, 2015). All aragonite stalagmites were deposited
within ∼1.5 km of the cave entrance where the bedrock is dolomite, and
all calcite stalagmites were deposited in more distal parts of the cave where
limestone constitutes the host rock (Zhang et al., 2015).
In November 2009 stalagmite SN17 (Fig. 3), 320 mm in length, was collected
200 m behind the cave entrance where the bedrock is dolomite. X-ray
diffraction (XRD) analyses suggest that the stalagmite is composed of
aragonite, except for the bottom section below 318 mm, which is composed of
calcite (Fig. 3). The calcite section was not included in the present study.
Discussion
Test of equilibrium deposition
Speleothem δ18O can be used to indicate climatic variation
provided that the speleothem was precipitated at or close to isotopic
equilibrium. The “Hendy test” is a widely used approach to explore to what
extent calcite deposition on the stalagmite surface occurred in isotopic
equilibrium with the parent drip water. Following Hendy (1971), 21 subsamples from three growth layers were analyzed; no progressive
increase in δ18O along individual growth layers and no
significant correlation of coeval δ18O and δ13C
values were found (Fig. 5a and b). This suggests (but does not prove) that the
stalagmite SN17 was deposited close to isotopic equilibrium. On the other
hand, δ18O and δ13C values show a statistically
significant covariance along the growth axis (r=0.37, p<0.01;
Fig. 5c), suggesting that the speleothem might be effected by kinetic
fractionation (Dorale and Liu, 2009). Some studies, however, demonstrated
that stalagmites showing a significant correlation between
δ18O and δ13C can also have formed under
isotopic equilibrium if both parameters are controlled by the common factors
(Dorale et al., 1998; Dorale and Liu, 2009; Tan et al., 2018a). A more robust
test is the replication of δ18O records from different caves
(Dorale et al., 1998; Wang et al., 2001; Dorale and Liu, 2009; Cai et al.,
2010). The δ18O records of SN17 and those from Dongge (Wang et
al., 2005) and Xiangshui Cave (Zhang et al., 2004), located southwest of
Shennong Cave (Fig. 1), show remarkable similarities in the overlapping
interval (Fig. 6a and b). The replication of these records further confirms
that aragonite in stalagmite SN17 was most likely deposited close to isotopic
equilibrium; i.e., its δ18O variations primarily reflect
climatic changes.
Interpretation of δ18O and δ13C
The climatic significance of the speleothem δ18O from
monsoonal China has been intensively debated in recent years. Most scientists
agree that speleothem δ18O in monsoonal China represent
variations in EASM intensity and/or changes in spatially integrated
precipitation between different moisture sources and the cave site on
millennial timescales (Cheng et al., 2016). Some researchers, however,
suggest that the Chinese speleothem δ18O is influenced by
moisture circulation on interannual to centennial timescales (Tan, 2009,
2014, 2016). In the region of spring persistent rain, the speleothem
δ18O values from Shennong Cave are controlled by both EASM and
NSM precipitation. A 200-year speleothem δ18O record from
E'mei Cave, located 160 km northwest of Shennong Cave, shows a significantly
negative correlation with both the EASM precipitation amount (r=-0.54,
p<0.01) and the EASM / NSM ratio (r=-0.67, p<0.01) during 1951–2009 CE (Fig. 3 in Zhang et al., 2018). The EASM
precipitation amount varies in the same direction as the EASM / NSM
ratio;
i.e., an increasing (decreasing) EASM / NSM ratio corresponds to more
(less) EASM precipitation. In addition, the E'mei δ18O record
also exhibits a coherent variation with the drought–flood index during
1810–2010 CE on decadal to centennial timescales (Fig. 2 in Zhang et al.,
2018). This indicates that E'mei δ18O is likely dominated by
the EASM precipitation amount on decadal to centennial timescales. Therefore,
we suggest that the stalagmite δ18O record from Shennong Cave,
similar to E'mei Cave, might be primarily influenced by the EASM / NSM
ratio, is also affected by the EASM precipitation amount on interannual to
decadal timescales, and can be dominated by the EASM precipitation amount on
decadal to centennial timescales, i.e., lower (higher) δ18O
values corresponding to higher (lower) EASM / NSM ratios and more (less)
EASM precipitation.
Changes in speleothem δ13C are generally controlled by
vegetation density and composition in the catchment, which vary according to
the hydroclimate (Genty et al., 2001, 2003, 2006; McDermott, 2004; Baldini et
al., 2005; Cruz Jr. et al., 2006; Fairchild et al., 2006; Fleitmann et al.,
2009; Noronha et al., 2015; Wong and Breeker, 2015). In regions where the
vegetation type is predominantly C3 or C4 plants, a dry climate
will lead to a reduction of the vegetation cover, density and soil microbial
activity as well as an increase in the groundwater residence time, allowing
more δ13C-enriched bedrock to be dissolved. In addition, prior
calcite precipitation (PCP) in the vadose zone will result in higher
δ13C values accompanied by increased Mg/Ca ratios in
speleothems (Baker et al., 1997). Slow drip rates and increased evaporation
and/or ventilation inside the cave will lead to higher δ13C
values, usually accompanied by kinetic isotopic fractionation (Fairchild et
al., 2000; Oster et al., 2010; Frisia et al., 2011; Li et al., 2011; Tremaine
et al., 2011; Meyer et al., 2014). In Jiangxi Province, a region presently
occupied by mostly C3 plants, no evidence has been found for a
replacement of C3 plants by C4 plants during the middle to late
Holocene (Zhou et al., 2004; Zhong et al., 2010b). There is no significantly positive
correlation between δ13C values and Mg/Ca ratios
between 5.3 and 3.57 ka BP (Fig. 5d). In Shennong Cave ventilation is weak
and relative humidity remains close to 100 % throughout the year. Rapid
CO2 degassing is less common under these conditions. Stalagmite
SN17 was likely deposited close to isotopic equilibrium, as confirmed by the
Hendy test and the “replication test” (Sect. 5.1). The δ13C variations in this stalagmite were primarily driven by vegetation
density and soil bioproductivity associated with hydroclimatic variations but
not by PCP or rapid CO2 degassing, with lower δ13C
values corresponding to a denser vegetation cover associated with a wet
climate and vice versa (Zhang et al., 2015).
Hydroclimate between 5.3 and 3.57 ka BP
Previous speleothem studies from monsoonal China suggested a coherent trend
of decreasing precipitation from the early to the late Holocene (Wang et al.,
2005; Hu et al., 2008; Cai et al., 2010, 2012; Dong et al., 2010, 2015; Jiang
et al., 2013; Tan et al., 2018a), which follows the gradually decreasing
Northern Hemisphere summer insolation. Our δ13C record
exhibits a similarly increasing trend (Fig. 4a, b), indicating that the
climate in our study area changed from wetter to drier conditions between 5.3
and 3.57 ka BP. Although the growth rate of stalagmites is often not a
direct function of precipitation amount (Railsback, 2018), some studies
suggest that in monsoonal China changes in the growth rate of stalagmite can be
influenced by variations in monsoon precipitation (Wang et al., 2005). The
long-term decreasing trend in growth rate, broadly consistent with changes in
the δ13C record, is possible related to decreased monsoon
precipitation between 5.3 and 3.57 ka BP (Fig. 4). But it should be noted
that more monitoring data are needed to confirm this relationship between
growth rate and precipitation in our study cave. The long-term trend in
δ18O is less significant than those of δ13C and
growth rate (Fig. 4) and might be caused by changes in precipitation
seasonality since the mid-Holocene, i.e., variations in the EASM / NSM
ratio. In this paper, we focus on the timing and nature of the 4.2 ka BP
event. The long-term trend of δ18O is not discussed in detail
because it remains unclear how EASM and NSM precipitation varied during
the Holocene.
As discussed in Sect. 5.2, the SN17 δ18O might be dominated by
the EASM precipitation amount on decadal to centennial timescales, although the
EASM / NSM ratio has a significant impact on interannual to decadal
timescales. On decadal to centennial timescales the SN17 δ18O
record shows a coherent variability with the δ13C record
(Fig. 6c), which is primarily influenced by the EASM precipitation amount but not
by seasonal precipitation δ18O. In addition, the SN17
δ18O record is remarkably similar to the δ18O
record from Dongge Cave (Fig. 6a), which is dominated by summer monsoon
precipitation (Wang et al., 2005). These observations indicate that, on
decadal to centennial timescales, variations in SN17 δ18O
might primarily reflect changes in the EASM precipitation amount, although the
EASM / NSM ratio may also have an impact, with higher (lower)
δ18O values corresponding to decreased (increased) summer
monsoon precipitation. The asynchronous variations between
δ13C and δ18O during the short intervals of
4.7–4.6, 4.3–4.2 and 4.05–3.95 ka BP (Fig. 6c) may be ascribed to the
delayed response of vegetation density to variations in the EASM
precipitation amount or the EASM / NSM ratio. During 4.5–3.57 ka BP, a
wet interval between 4.2 and 3.9 ka BP can be identified in both
δ18O and δ13C records, consistent with the time
of high growth rate between 4.26 and 4.0 ka BP (Figs. 4 and 6c).
Therefore, we suggest that the climate in the study area between 5.3 and
4.5 ka BP was dominantly wet and changed to a rather dry climate between
4.5 and 3.57 ka BP, interrupted by one wet interval between 4.2 and
3.9 ka BP (Figs. 4 and 6c). It indicates that the climate in our study area
during the 4.2 ka BP event (4.2–3.9 ka BP) was predominantly wet.
Comparison of δ18O records from (a) Lianhua
Cave (Dong et al., 2015), (b) Jiuxian Cave (Cai et al., 2010),
(c) Xianglong Cave (Tan et al., 2018a), (d) Sanbao Cave
(Dong et al., 2010), (e) Heshang Cave (Hu et al., 2008),
(f) Shennong Cave (this study), (g) Dongge Cave (Wang et
al., 2005), (h) Mawmluh Cave (Berkelhammer et al., 2012) and the
ice-rafted hematite-stained grain (HSG) record from the North Atlantic (Bond
et al., 2001). 230Th dates and error bars are illustrated with
different colors for each stalagmite. The yellow bar marks the dry interval in
northern and southwestern China and the wet interval in central and
southeastern China during the 4.2 ka BP event (4.2–3.9 ka BP).
Comparison with other records in monsoonal China covering the
4.2 ka BP event
The nature and timing of the 4.2 ka BP event in southern and northern China
are still controversial because the discrepancies might also be caused by
the large dating uncertainties and the low proxy resolution in some records.
By reviewing records from the monsoon region of China, Tan et al. (2018a)
have
already proposed a “north dry, south wet” pattern during the 4.2 ka BP
event; however, more speleothem records from SEC are still needed to confirm
this pattern. In this section, we compare high-precision and
high-resolution speleothem records from monsoonal China during the
interval of 5.4–3.6 ka BP.
Map showing some locations discussed in the text. 1: Shennong Cave,
this study; 2: Xiangshui Cave (Zhang et al., 2004); 3: Heshang Cave (Hu et
al., 2008); 4: Sanbao Cave (Dong et al., 2010); 5: Jiuxian Cave (Cai et al.,
2010), 6: Xianglong Cave (Tan et al., 2018a); 7: Lianhua Cave (Dong et al.,
2015); 8: Nuanhe Cave (Tan, 2005); 9: Dongge Cave (Wang et al., 2005); 10: Dark
Cave (Jiang et al., 2013); 11: Shigao Cave (Jiang et al., 2012); 12: Xianren
Cave (Zhang et al., 2006); 13: Mawmluh Cave (Berkelhammer et al., 2012);
14: Erhai Lake (Zhou et al., 2003); 15: Tianchi Lake (Zhao et al., 2010);
16: Gonghai Lake (Chen et al., 2015); 17: Daihai Lake (Xiao et al., 2018a);
18: Dali Lake (Xiao et al., 2008); 19: Hulun Lake (Xiao et al., 2018b);
20: Daiyunshan peat (Zhao et al., 2017), 21: Dahu peat (Zhou et al., 2004),
22: Daping peat (Zhong et al., 2010a); 23: Dajiuhu peat (Ma et al., 2008);
24: Chengjiachuan site (Huang et al., 2010); 25: Huxizhuang loess–soil profile
(Huang et al., 2011); 26: Tengchongqinghai Lake (Zhang et al., 2017);
27: Gaochun profile (Yao et al., 2017); 28: Zhongqiao site (Wu et al., 2017).
The solid triangle, dot and square denote stalagmite records, lake sediment–peat
records and paleoflood sediment records, respectively. Black and blue
indicate a dry and a wet climate during the 4.2 ka BP event, respectively.
The base map is the same as that in Fig. 1.
For northern and southwestern China, the stalagmite δ18O
record from Lianhua Cave (Fig. 1) shows a long-term decreasing trend of
summer monsoon precipitation (Fig. 7a; Dong et al., 2015), consistent with
both δ18O and δ13C records of stalagmites from
Nuanhe Cave in northeastern China (Fig. 1; Tan and Cai, 2005; Wu et al.,
2011; Zhang and Wu, 2012). These records indicate that the climate in
northern China gradually varied from wet to dry between 5.4 and 3.6 ka BP,
and the climate was very dry during the interval of 4.2–3.9 ka BP. The
δ18O record from Dongge Cave, southwest of monsoonal
China, reveals a dry event between 4.4 and 3.95 ka BP (Fig. 7g; Wang et
al., 2005), which is consistent with decreased precipitation from the Indian
summer monsoon during the interval of 4.3–3.9 ka BP (Fig. 7h; Berkelhammer
et al., 2012), and dry intervals in the stalagmite records from Shigao, Dark
and Xianren Cave, southwestern China (Fig. 8 and references therein); however, a
wet climate is documented by multi-proxy data from a maar lake in southwest China
(Fig. 8; Zhang et al., 2017). The prominent drought during the 4.2 ka BP
event was also recorded by various other archives from sites in northern and
southwestern China (Fig. 8 and references therein).
For north and south of the Qinling Mountains, two stalagmite δ18O records from Jiuxian and Xianglong Cave south of the mountains
exhibit coherent variations on centennial timescales but neither of them
shows a long-term increasing trend (Fig. 7b and c; Cai et al., 2010; Tan et
al., 2018a). Both δ18O records reveal increased monsoon
precipitation between 4.3 and 3.8 ka BP, indicating that the climate south
of the Qinling Mountains was wet during the 4.2 ka BP event (Tan et al.,
2018a). It should be noted that there is an asynchronous variation during the
interval of 4.2–3.9 ka BP between the Jiuxian and Xianglong δ18O records, which might be due to either dating uncertainties in the
Jiuxian record or the spatial distribution of monsoon precipitation between
these two regions. The extraordinary flood during the 4.2 ka BP event was
also identified in the middle reaches of the Yellow River north of the
Qinling Mountains, north-central China (Fig. 8; Huang et al., 2010, 2011).
For south-central China, two stalagmite δ18O records from
Sanbao (Fig. 7d; Dong et al., 2010) and Heshang (Fig. 7e; Hu et al., 2008)
Cave in the middle reaches of the Yangtze River, south-central China, also
indicate a wet interval between 4.2 and 3.9 ka BP, which is consistent with
a δ13C peat record from the Dajiuhu basin (Ma et al., 2008)
and paleoflood sediments from Jianghan Plain (Wu et al., 2017) in the same
region (Fig. 8). δ15N and δ13C records from the
Daping swamp in Hunan Province, south-central China, also reveal a wet
interval at 4.5–4.0 ka BP (Fig. 8; Zhong et al., 2017).
For SEC, SN17 δ18O and δ13C records reveal a
wet interval between 4.2 and 3.9 ka BP on centennial timescales (Figs. 6c
and 7f), which is consistent with the speleothem δ18O record
from Xiangshui Cave (Fig. 6b). A record of total organic carbon (TOC) from
the Dahu swamp, Jiangxi Province, located 450 km south of Shennong Cave,
indicates a dry climate between 6.0 and 4.0 ka BP, with a short-lived wet
event at 4.1 ka BP (Fig. 8; Zhou et al., 2004). Subsequent multi-proxy
records of several new cores from this site also revealed a prevailingly dry
climate between 6.0 and 3.0 ka BP with a wet interval at 4.2–3.9 ka BP
(Zhong et al., 2010a, b, c). Pollen data from two sediment profiles from
Daiyunshan, a mountain in SEC, indicate a centennial-scale wet event at
4.4 ka BP (Fig. 8; Zhao et al., 2017). These published records from SEC are
consistent with the SN17 δ18O and δ13C records
within error, indicating a wet climate in SEC during the 4.2 ka BP event
(Fig. 8).
To sum up, high-resolution stalagmite records indicate that the climate
during the 4.2 ka BP event was dry in northern and southwestern China and
wet in southern China (Figs. 7 and 8), and the nature and timing of this
event were different in different regions of monsoonal China. The
remarkable dry climate lasted from ∼4.4 to 3.9 ka BP in northern and
southwestern China (Fig. 8). South-central China and SEC were dry between 4.4
and 4.2 ka BP and wet between 4.2 and 3.9 ka BP (Fig. 8). The climate
was wet between 4.4 and 4.2 ka BP but was dry between 4.2 and 3.9 ka BP
at Jiuxian Cave (Figs. 7b and 8), southeast of the Qinling Mountains, and the
climate was wet between 4.3 and 3.8 ka BP at Xianglong Cave, southwest of
this mountain range (Figs. 7c and 8). A dry and cold period during the
4.2 ka BP event was identified in a sediment profile from Gaochun, west of
Taihu Lake, eastern China (Fig. 8, Yao et al., 2017). Therefore, we suggest that
the boundary between the dry north and the wet south during the 4.2 ka BP
event was probably located along the northern rim of the Qinling Mountains
and the lower reaches of the Yangtze River (Fig. 8).
This south–north distribution of monsoon precipitation might have been caused
by a weakened EASM intensity, which could have resulted from a reduced
Atlantic Meridional Overturning Circulation (AMOC) recorded by higher
abundances of ice-rafted debris (IRD) in the North Atlantic (Fig. 7i). Strong
freshwater input into the North Atlantic derived from melting icebergs
periodically reduced the AMOC (Bond et al., 2001), causing a temperature
decrease in the high northern latitudes and intensified midlatitude
westerlies. As a consequence, the Intertropical Convergence Zone and Northern
Hemisphere westerly jet got stronger, migrated southward, and weakened
the Indian summer monsoon and the EASM (Wang et al., 2001; Chiang et al.,
2015). The stronger westerly jet and the weakened EASM delayed the westerly
jet transition from south of the Tibetan Plateau to the north in early to
middle
May and postponed the onset of the EASM (Chiang et al., 2015; Tan et al.,
2018a). The rain belt migrated southward and remained longer in southern
China than normal, which reduced rainfall in northern and southwestern China
(Fig. 7a and g) but enhanced rainfall in central and southern China (Fig. 7c,
d, e and f) during the periods of 4.7–4.5 and 4.2–3.9 ka BP.
The SN17 δ18O record exhibits a coherent variation with the
δ18O record from Dongge Cave within error on centennial
timescales but there are some differences in amplitude (Figs. 6a and 7),
which can also be found between the Jiuxian and Xianglong records and between
the Sanbao and Heshang records (Fig. 7). These discrepancies might be due to
chronology offsets because some records are constrained by two to three
dates only. Alternatively, the different amplitudes of these records might be
caused by the spatial distribution of monsoon precipitation, reflecting the
position and residence time of the rain belt associated with variations in
EASM intensity. For example, during the period of 4.2–3.9 ka BP, reduced
AMOC caused a dry climate in northern and north-central China but a wet
climate in south-central China and SEC (Figs. 7 and 8). During the period of
4.6–4.5 ka BP, however, the records show a dry climate from northern China
to south-central China but a wet climate in SEC (Figs. 7 and 8). Because the
AMOC was in a very weak stage at 4.55 ka BP this could have resulted in a
weakened EASM and a further southward shift of the monsoonal rain belt,
possibly causing a dry climate in south-central China and a wet climate in
SEC. Recently, Yan et al. (2018) used a set of long-term climate simulations
and suggested that the 4.2 ka BP event could be related to the slowdown of
the AMOC but was more likely caused by internal variability in the climate
system. Detailed modeling studies and additional high-resolution records are
needed to further investigate the possible causes and mechanisms of this
event.