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The Holocene
22(6) 705 –715
© The Author(s) 2011
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DOI: 10.1177/0959683611417740
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Research paper
Introduction
The East Asian monsoon (EAM) forms as a result of unequal
heating of the Asian landmass and the Pacific and Indian Oceans
(Figure 1a) and subsequent energy re-distribution. At the present
day the EAM plays a significant role in the global and regional
hydrological and energy cycles (Webster et al., 1998). It is sug-
gested that fluctuations during the Holocene exerted a major
influence on human societies (Mischke and Zhang, 2010; Shi et al.
1994). For example, during the stable warm and wet phase from
7200 to 6000 cal. yr BP, the Neolithic Yangshao Culture in the
Yellow River catchment and the Majiabang Culture in the lower
Yangtze plain, reached their climax (Shi et al., 1994). A sudden
dry and cold event around 4000 cal. yr BP recorded in the Dunde
ice core, Northwest China (Thompson et al., 1989, 1993), is
suggested to be responsible for the end of the Longshan and
Liangzhu Cultures in East China (Shi et al., 1994). Nearly two-
thirds of the world’s population live in areas influenced by the
EAM. A better understanding of the past variability of the EAM
and the associated links with human cultures is, therefore, crucial
to improve our understanding of potential implications of future
changes in monsoon climate.
The history of the EAM during the Holocene can be regarded
as an alternation between dominance of a cold-dry winter mon-
soon and a warm-wet summer monsoon (An, 2000). After the
Younger Dryas (11 200–10 000 cal. yr BP), a cold period which
might have included some dry/wet oscillations (Xiao et al., 1998),
the EAM switched to relatively warm and wet conditions (the
Holocene). Based on evidence from ice cores (e.g. Thompson
et al., 1989, 1993), lake water levels (e.g. Li et al., 1990; Wang
and Fen, 1991), as well as archaeological records (e.g. Zhou
et al., 1991), Shi et al. (1994) suggested that Holocene Megather-
mal phases mainly occurred during 8500–3000 cal. yr BP. They
divided the Holocene EAM into several phases: 8500–7200 cal.
yr BP characterised by unstable temperature fluctuations with an
increase in precipitation; 7200–6000 cal. yr BP was a stable,
warm and wet phase when the monsoon rainfall occurred almost
throughout China; 6000–5000 cal. yr BP was characterized by
strong climatic fluctuations and adverse climatic conditions; after
5000 cal. yr BP the climate became warm and stable – suitable
for cultural development in China. Around 4000 cal. yr BP, the
climate deteriorated and catastrophic floods occurred in East
China; from 4000 to 3000 cal. yr BP, the climate was warm and
humid. Based on geological data and numerical modelling, An
et al. (2000) suggested that the front of the EAM retreated from
417740HOLXXX10.1177/095
9683611417740Yu et al.The Holocene
1University of Durham, UK
2Nanyang Technological University, Singapore
3The University of Hong Kong, China
4British Geological Survey, UK
5City University of Hong Kong, China
6Guangzhou Institute of Geography, China
Received 30 August 2010; revised manuscript accepted 26 June 2011
Corresponding author:
Fengling Yu, Department of Geography, University of Durham, South
Road, Durham DH1 3LE, UK.
Email: fengling.yu@ntu.edu.sg
Mid-Holocene variability of the
East Asian monsoon based on bulk
organic d13C and C/N records from
the Pearl River estuary, southern China
Fengling Yu,1,2
Yongqiang Zong,3 Jeremy M Lloyd,1 Melanie J Leng,4
Adam D Switzer,2 Wyss W-S Yim3,5 and Guangqing Huang6
Abstract
Understanding the mid-Holocene dynamics of the East Asian monsoon (EAM) is integral to improving models of the Holocene development of the global
climate system. Here we reconstruct the mid-Holocene EAM history from the Pearl River estuary, southern China, using bulk organic carbon isotopes
(d13C), total carbon to total nitrogen (C/N) ratios and total organic carbon (TOC) concentration. Sedimentary d13C, C/N and TOC are potentially good
indicators of changes in monsoonal precipitation strength. Sediments buried during a period of high precipitation exhibit a high proportion of terrigenous
material, and have low d13C and high C/N, and vice versa during a period of low precipitation. Results suggest a general decreasing trend in monsoonal
precipitation from 6650 to 2150 cal. yr BP because of the weakening Northern Hemisphere insolation most likely related to the current precession circle.
Superimposed on this trend are apparent dry–wet oscillations at centennial to millennial timescales most likely in response to solar activity. Mismatches
between our d13C record and results from the Dongge Cave in southern China at millennial timescales may indicate that the d13C from the Pearl River
estuary reveals changes in precipitation over a broader area than the d18O from Dongge Cave.
Keywords
bulk organic d13C, C/N, China, East Asian monsoon history, Holocene, Pearl River estuary, solar forcing
706 The Holocene 22(6)
north to south China during the Holocene. This resulted in the
Holocene Optimum, a period of maximum precipitation, occur-
ring earlier in the north and northeast of China and later in the
south and southeast. The hydrological and thermal dynamics of
the EAM system are perturbed by external and internal factors at
a range of timescales. These factors range from solar forcing at
millennial timescales, high latitude cooling events at centennial
timescales, to low latitude El Niño events at decadal timescales
(An et al., 2000).
Variability of the Holocene EAM has been reported from a
range of archives, including natural records, such as loess (e.g.
Porter and An, 2005), tree rings (e.g. Wu, 1992; Wu et al., 1988),
ice cores (e.g. Thompson et al., 1989), and stalagmites (e.g.
Wang et al., 2005), as well as historical archives (Shi et al.,
1994). Each of these has specific advantages and disadvantages.
Loess deposits from the Loess Plateau, central China, are good
archives for investigating the EAM over timescales of millions
of years. However, it is hard to undertake high-resolution studies
based on these records because of the low sedimentation rates.
Ice cores and deep-ocean deposits also suffer from a similar
problem of low resolution. Stalagmites provide high-resolution
studies; however, the δ18O signature can be influenced by both
temperature and precipitation, which are also subject to local
scale variability. Historical records can provide valuable infor-
mation, but can be difficult to interpret and are spatially and
temporally fragmented.
This study reconstructs the EAM during the mid Holocene
from a estuarine core (core UV1) collected from Pearl River
estuary (Figure 1b). Previous studies in the Pearl River estuary
have shown that changes in the organic carbon isotopic signature
of estuarine sediments can be used as an indicator for changes in
the freshwater discharge induced by the monsoonal precipitation
(Yu et al., 2010; Zong et al., 2006).
Over centennial and millennial timescales the sources of estu-
arine sediments (Figure 2) are influenced by relative shoreline
movements, sea-level changes, and monsoonal climate change.
Zong (2004) re-examined the Holocene sea-level history for
the low latitude part of the China coast by re-assessing all the
sea-level data available from the east to south coasts. The sea-
level curve from west Guangdong and Hainan, a geologically
stable region, shows that the relative sea level reached the
present-day height by 6200 cal. yr BP and since then has been
relatively stable (Zong, 2004). The time of maximum emergence
of the sea-level highstand occurs earlier at tidal river sites than at
coastal sites (Zong, 2004). In the Pearl River Delta, the maxi-
mum emergence of the highstand was around 7000 cal. yr BP
(Li et al., 1990; Zong, 2004).
As the sea level in this area has been relatively stable since the
mid Holocene, monsoonal climate is one of the most important
controlling factors for the sediment flux into the estuary (Li et al.,
1990; Wang et al., 2005; Zong et al., 2006). For example, Zhang
et al. (2008) suggested a positive linear correlation between
Guangzhou
Hong Kong
Macau
Shenzhen
114º00' E
22º00' N 23º00' N
113º00' E
113º30' E
23º30' N
West River
North River
East River
South China Sea
b
West River
North River
East River
S o u t h C h i n a S e a
Pearl River delta
and estuary
V I E T N A M
115º E110º E105º E
25º N
20º N
a
0 200
km
C H I N A
6800 cal. yr. BP
4500 cal. yr. BP
2000 cal. yr. BP
1000 cal. yr. BP
UV1
112º30' E
23º00' N
6800 cal. yr. BP
4500 cal. yr. BP
2000 cal. yr. BP
1000 cal. yr. BP
Palaeo-shorelines
030km
Core UV1
Figure 1. Study area. (a) Location of Pearl River estuary, East Asia; (b) locations for core UV1 (this study) and palaeoshorelines (Zong et al.,
2009a)
Yu et al. 707
cumulative freshwater discharge and cumulative sediment flux in
the Pearl River. As outlined by Zong et al. (2006) the Pearl River
estuary receives sediments from three main sources: river, marine
and in situ brackish-water derived material (Figures 1b, 2); the
proportion of these three main sources of organic matter to the
estuarine sediment is controlled by the strength of the river dis-
charge due to monsoonal precipitation. Therefore, during a period
of strong summer monsoon, increased precipitation levels will
generate a large volume of freshwater discharge into the estuary,
along with high sediment flux. It follows that estuarine sediments
preserved during periods of strong summer monsoon will have a
higher proportion of terrigenous sediments compared with periods
of stronger winter monsoon (lower precipitation, lower freshwater
flux, hence lower proportion of terrigenous sediments relative to
marine/brackish sediments). Thus, the monsoon-induced fresh-
water discharge can be reconstructed by examining changes in the
relative dominance of organic matter from the three sources iden-
tified above preserved in the Pearl River estuary (Yang et al.,
2011; Yu et al., 2010; Zong et al., 2006).
This study applies this methodology to reconstruct the EAM
during the mid Holocene at decadal resolution, employing bulk
organic carbon isotopes (δ13C) and total carbon to total nitrogen
(C/N) ratios, to help understand possible driving mechanisms
behind monsoon variability. Application of δ13C and C/N anal-
yses of sedimentary organic matter for reconstructing pal-
aeoenvironmental changes have been carried out in a range of
environments (Lamb et al., 2006), including lagoons (Müller and
Mathesius, 1999; Yamamuro, 2000), isolation basins (Chivas et al.,
2001; Mackie et al., 2007; Westman and Hedenström, 2002),
fjords (Bird et al., 1991; St-Onge and Hillaire-Marcel, 2001;
Smittenberg et al., 2004), as well as estuaries (Malamud-Roam
et al., 2006; Middelburg et al., 1997; Wilson et al., 2005; Zong et al.,
2006). The advantage of δ13C and C/N analysis is that it can be
performed on most sediment, as only a small amount of organic
carbon (micrograms to milligrams) is required for analysis.
Organic δ13C and C/N ratios can be measured rapidly and analyses
are relatively inexpensive, leading to high-resolution studies
(Lamb et al., 2007). To explore driving mechanisms for oscilla-
tions of the EAM at millennial to centennial timescales, general
trends of δ13C and other proxies were removed.
Study area
The Pearl River catchment, located between 21°20′–23°30′N and
112°40′–114°50′E (Figure 1a), formed as a result of the uplift of
the Tibetan Plateau during the Tertiary and Quaternary periods
(Aitchison et al., 2007). Before the late Quaternary, sediment
from the river system bypassed the current deltaic basin and was
deposited on the continental slope and shelf (Zong et al., 2009a).
Only since the Late Pleistocene has the deltaic basin started to
receive sediments from the river system (Huang et al., 1982; Xu
et al., 1985). The lower fluvial/deltaic plains of Guangdong and
Guangxi provinces have been gradually infilled during the late
Quaternary with progradation of the delta front since 6800 cal. BP
(Zong et al., 2009a; Figure 1b).
The Pearl River catchment lies in the transitional area between
the tropical and subtropical climate zones. The West River, 2214
km in length with a catchment area of 425 700 km2, contributes
the majority of both water and sediment flux reaching the Pearl
River Estuary (Kot and Hu, 1995; Zhang et al., 2008). Under
strong influence of the monsoon climate, freshwater discharge
and sediment flux show great seasonal variability due to seasonal
changes in monsoonal precipitation. The mean annual precipita-
tion of the catchment is 1200–2000 mm, which mainly occurs
during the wet season lasting from April to September. Annual
discharge of the whole Pearl River system is 330 × 109 m3 (Hu et al.,
2006), and the annual sediment load is 89 × 109 kg (Zhang et al.,
2008). Statistics since 2002 show that more than 86.9% of the
annual suspended sediment is discharged during the wet season
and 67.6−83.5% of annual water flux occurs in this season
(Shen and Wang, 2009). Seasonal changes in monsoonal precipita-
tion also results in changes in salinity within the estuary. For
example, in summer (June–August), high precipitation generates
strong freshwater flux into the estuary, resulting in a low-salinity
estuarine environment, with water salinity lower than 1 PSU at
the river mouth. Saline water intrudes much further upstream in
winter (December–February) than in summer because of a sig-
nificant reduction in freshwater flux (Yu et al., 2010; Zong et al.,
2010a). In addition to the seasonal salinity variability, within the
estuary there is a northwest–southeast isohaline system, due to
decreasing freshwater influence seawards (Yu et al., 2010) and
the strong tidal currents in the eastern part of the estuary (Zong
et al., 2010a).
Material and methods
Vibracore UV1 was collected from 22°17′10″N, 113°51′49″E
(Figure 1b) from a water depth of 8.6 m. The top 10 m of sedi-
ment spanning the last c. 6500 years was used for this study.
The top 0.35 m of the core was washed away during the coring
process and sediments from 6.00–6.25 m were missing because
of gaps in the sampling tube. The core was sampled at 2 cm
intervals for δ13C, C/N and TOC, producing a total of 501
samples.
The chronology for UV1 used here is that of Zong et al.
(2010b) and is based on seven radiocarbon dates which are mostly
from benthic foraminifera samples collected from various depths
(Table 1). There is a small age reversal between GZ2213 at 2.6 m
and SUERC-9602 at 1.9 m, the reason for this age reversal is not
clear. The sedimentation environment was stable because parti-
cle size analysis shows no major change in sand, silt or clay
fractions (Zong et al., 2010c), and there is no bioclastic storm
bed found in this sediment section (e.g. Huang and Yim, 2001).
The only thin sandy layer found in the core is located at 1.0 m
(Zong et al., 2010c). Diatom assemblages from the section con-
taining the age inversion show no significant change either
(Zong et al., 2010b). All other dates lie in stratigraphic order
producing a relatively robust age model (Figure 3; Zong et al.,
2010b). Accordingly, these dates suggest a period of steady sedi-
mentation from c. 6650 cal. yr BP at 10 m depth to c. 3300 cal.
yr BP at 2.6 m. The average sedimentation rate in this period
is c. 0.22 mm/yr. The top 2 m section of the core was deposited
in the last 3300 years, with a reduced sedimentation rate of c.
0.08 mm/yr (Figure 3).
River discharg
eT
ide transport
Freshwater algae
Terrestrial-plant debris
Fluvial suspended sediment
Brackish-water algae
Aquatic vegetation
Mixed suspened sediment
Marine algae
Marine-plant debris
Marine-suspended sediment
Sediments
Pearl River catchment South China Sea
Monsoonal
precipitation
Sediments
Core HKUV-1
Pearl River estuary
Figure 2. Sediment input into the Pearl River estuary from three
main sources including terrestrial, brackish-water and marine areas
(adapted from Zong et al., 2006)
708 The Holocene 22(6)
Sediment samples for δ13C and C/N analyses were prepared
using 100 ml 5% HCl to remove carbonates. They were then
washed three times with deionised water using the fibreglass filter
paper (Fisher Brand 200), before being dried at 50°C overnight,
homogenised in a pestle and mortar, and weighed (25–50 mg) for
δ13C and C/N analysis. Carbon isotopes and C/N analyses were
performed by combustion in a Carlo Erba NA1500 (Series 1)
online to a VG Triple Trap and Optima dual-inlet mass spectrom-
eter, with δ13C calculated to the VPDB scale using a within-run
laboratory standards (BROC) calibrated against NBS-19 and
NBS-22. Replicate analysis of well-mixed samples indicated a
precision of ±<0.1‰ (1SD). C/N was determined by reference to
an Acetanilide standard. Replicate analysis of well-mixed sam-
ples indicated a precision of ±<0.1%. The isotopic ratios are
expressed as δ13C, in units of per mil (‰). The C/N is given using
weight ratio of total organic carbon (%TOC) to total nitrogen
(%TN). Where duplicate analyses was carried out, the data
presented are the average value (SC typically <0.1‰).
For particle size analysis, sediment (0.2−4 g) was soaked in 5%
Calgon solution overnight before analysis to allow disaggregation
of the sediment. An ultrasonic treatment was applied when the
sample did not completely disaggregate. Samples were then stirred
before being introduced to the laser granular meter (Coulter LS
13200) for the analysis of fractions of sand, clay and silt. Results
of particle size analysis are presented in the format of ‘average
value ± standard deviation (SD)’.
Exponential smoothing of δ13C and C/N data were performed
using the SPSS for Windows program. This method was chosen
over other methods such as moving average, because it does not
result in data loss, and it catches the general trend of the data
series. To compare periodicity between δ13C and other proxies,
the new δ13C data presented here was de-trended along with the
previously published 10Be (Finkel and Nishiizumi, 1997) and
Hematite-strained glass data (HSG, Bond et al., 2001). Spectral
analysis on δ13C and 10Be and de-trending of data were carried out
using the PAST program (Hammer et al., 2001).
Results
The δ13C, C/N and TOC data are presented in Figure 4. There is
a clear change in sedimentation at 10.07 m marking the bound-
ary between Holocene sedimentation above and pre-Holocene
sedimentation below (poorly sorted firm clay through to gravel
sediments below, soft silt and clay above). This is supported by
the radiocarbon date of 33 500 ± 360 cal. yr BP at 12.5 m (Zong
et al., 2009b). The age of the base of the Holocene section is
estimated at 6650 cal. yr BP based on the age model developed
by Zong et al. (2010b) (Figure 3). One of the possible reasons
for the absence of early-Holocene deposition is that the core site
was subaerially exposed at this time, when the sea level was low
(Zong, 2004; Zong et al., 2009a), and sediment only starts accu-
mulating when the sea level stabilized around 7000 cal. yr BP.
The Holocene section of the core is split into two zones based
on the lithology, grain size and δ13C values. Zone 1 from 10.07 to
1.30 m is composed of dark greenish grey silt and clay (63.6±5.2%
silt and 24.9±3.9% clay; Table 2). Zone 2 from 1.30 to 0.35 m is
composed of yellowish silt and clay with a higher proportion of
sand (17.0±12.1%; Table 2), mixed with large amounts of shell
fragments. There are clear changes in δ13C, TOC and C/N values
between Zones 1 and 2 (Figure 4); the average δ13C value for
Zone 1 is −25.2±0.3‰, compared with −23.1±2.0‰ for Zone 2
(Figure 4). The radiocarbon date from 0.52 m produced a modern
age (108 cal. yr BP), while the date from 1.33 m (immediately
below the boundary between Zones 1 and 2) provided an age of
2337–2162 cal. yr BP. Based on the chronology and clear change
in character of the sediments, the change from Zone 1 to Zone 2
most likely relates to enhanced human activity in this area as dis-
cussed by Zong et al. (2010c). In this paper we concentrate on the
mid-Holocene section preserved in Zone 1.
Sediment from Zone 1 is mainly composed of silt (average
63.5 ± 5.2%) and clay (average 24.9 ± 3.9%), while sand concen-
trations average 11.4 ± 6.6% (Figure 5; Table 2). The proportion
of sand and silt are negatively correlated through the core. Clay
Table 1. Radiocarbon dates from Core UV1
Depth (m) Material dated Conventional
age (yr BP)
Method 95% HDR
(cal. yr BP)
50% HDR
(cal. yr BP)
%C (by weight) d13C (‰,VPDB)
± 0.1
Laboratory code
0.5 Shell Modern GZ2211
1.3 Foraminifera 2254 ± 30 AMS 14C 2337-2162 2233 GZ2212
1.9 Foraminifera 3019 ± 35 AMS 14C 3333-3087 3232 9.8 -3.1 SUERC-9602
2.6 Foraminifera 2974 ± 33 AMS 14C 3292-3025 3157 GZ2213
4.5 Foraminifera 3963 ± 35 AMS 14C 4516-4299 4434 8.8 -4 SUERC-9605
7.5 Foraminifera 4847 ± 35 AMS 14C 5647-5486 5593 9.7 -4.1 SUERC-9606
9.5 Foraminifera 5633 ± 36 AMS 14C 6486-6320 6413 9 -2.6 SUERC-9607
12.5aShell 37 900 ± 320 AMS 14C 41 220-40 900 41 600 -1.4 OS-51226
Radiocarbon dates are calibrated using Calcuve Intcal 09, 95% and 50% HDR ages are estimated using Bchron (Parnell et al., 2008), HDR, highest
posterior density region; adates are cited from Zong et al. (2009b), where calibrated ages are expressed at ±2s level.
2
0
4
6
8
10
12
0 4000 6000 8000 10000
Modal age (n=86)
95% HDR
Unrestricted calibrated dates
Depth (m)
Age (cal. yr BP)
2000
Figure 3. Age model for Core UV1 (modified from Zong et al.,
2010b). The age model is established via Bchron using the method
suggested by Heegaard et al. (2005) and Parnell et al. (2008).
The 95% highest posterior density ranges (HDR) indicate the
uncertainty of the ages assigned to each sample between the dated
depths, together with the mean modelled age for each sample.
In the graph, the number of samples involved in the calibration is
shown as n, and the modelled age curve is established based on the
50% chronology
Yu et al. 709
concentration gradually decreases from approximately 30.0% at
the base to 20.0% at the top of the core.
The δ13C values of Zone 1 show a general increasing trend
from the base to the top of the zone (Figure 5). Superimposed on
this increasing trend, the δ13C values fluctuate between −26.4‰
and −24.1‰, the average value (−25.2‰) for Zone 1 is shown by
the dashed line (Figure 5). Exponential smoothing of the δ13C data
suggests two oscillations are recorded in this section. The first
complete oscillation starts from the base of this section, c. 6650
cal. yr BP to c. 4600 cal. yr BP. The second oscillation runs from
c. 4600 to c. 2100 cal. yr BP, although the last part of this oscilla-
tion is not so clearly defined and may be incomplete (Figure 5).
The mean C/N ratio in Zone 1 is 10.5 (dashed line, Figure 5). C/N
ratios in Zone 1 show a general decreasing trend, except for very
low C/N ratios (4.0−9.0) from 6650 to 6480 cal. yr BP (Figure 5).
TOC records of UV1 show a general decreasing trend from 1.4%
to 0.8% throughout Zone 1 with the average value of 1.01%
(Figure 5). The exponential-smoothed TOC record shows a
relatively close correlation with the δ13C values (Figure 5).
Discussion
An East Asian monsoon history during the mid Holocene
The δ13C and C/N signature preserved in core UV1 outlined
above can be used to infer changes in the sediment sources dur-
ing the mid Holocene. Changes in sediment source can then be
used to indicate changes in EAM precipitation-induced fresh-
water flux.
General weakening EAM during the mid Holocene. Increasing
δ13C, decreasing C/N ratios and decreasing TOC found in
sediments spanning the period 6650−2150 cal. yr BP indicate a
long-term trend of weakening precipitation-induced freshwater
discharge linked to a weakening summer monsoon (Figure 5). This
trend is highlighted in Figure 6a by higher than average δ13C values
(−25.2‰) from 4350 to 2150 cal. yr BP and generally lower than
−25.2‰ from 6650 to 4350 cal. yr BP. Possible forcing mecha-
nisms that could influence the proportion of sediments from differ-
ent sources in UV1 during this time period are sea-level change,
9
8
7
6
5
4
3
2
1
2337-2162
3292-3025
4516-4299
5647-5486
6486-6320
41220
-40900
108
3333-3087
Age
(cal yr BP)
10
12.5
0.40.81.21.6
TOC (%)
Zone 2
pre-Holocene
Zone 1
(Mid Holocene)
-28-26 -24-22 -20
13
δ C (‰)
10.08m
Depth
(m)
0
1.31m
481216
C/N
Figure 4. Results of d13C, C/N and TOC for core UV1 (10.29–0.35m)
Table 2. Measurements of Zone 1 and Zone 2 of Core UV1
Depth (m) Age (cal. yr BP) d13C (‰) TOC (%) C/N Sand (%) Clay (%) Silt (%)
0.35-1.29 1364-1990 -23.1 ± 2.0 0.1 ± 0.0 11.3 ± 1.7 17.0 ± 12.1 22.2 ± 4.4 60.8 ± 10.8
1.31-10.07 2003-6395 -25.2 ± 0.3 0.1 ± 0.0 10.5 ± 1.2 11.4 ± 6.6 24.9 ± 3.9 63.6 ± 5.2
Values are presented as ‘average values± the standard deviation’.
710 The Holocene 22(6)
7000
6000
5000
4000
3000
2000
-26-25.6 -25.2-24.8 -24.4
13
δ C (‰)
Freshwater Brackish
13 12 11 10 9
C/N
1.41.2 1 0.8
TOC (%)
Freshwater Brackish Freshwater Brackish
Raw data Exponentially smoothedMean value
Age (cal yr BP)
Sand (%) Clay (%)
30 20 10 0
10
40
30 20
Silt (%)
80
50
6070
Freshwater Brackish Freshwater Brackish FreshwaterBrackish
Figure 5. Results of d13C, C/N, TOC and particle size during the mid Holocene (10.07–1.31 m of core UV1, ranging from 6650 to 2150 cal. yr BP) plotted against calibrated ages at 95% HDR
Yu et al. 711
delta progradation and freshwater flux. Sea level reached its pres-
ent-day level around 6800 cal. yr BP and has been relatively stable
since then (Zong, 2004). Thus, the influence of sea-level changes
on the δ13C signal is assumed to be relatively minor. As a conse-
quence of the stable sea level, the deltaic shoreline advanced slowly
seawards between 6800 and 2000 cal. yr BP (Figure 1b, Zong et al.,
2009a). Delta progradation during the mid Holocene has led to core
UV1 being more proximal to terrestrial source areas through time
(Figure 1b). Based on this we would expect larger amounts of ter-
restrial sediment to be delivered to the core site through time. How-
ever, the general increase in bulk organic carbon δ13C values
recorded in UV1 during the mid Holocene indicates a decreasing
proportion of terrestrial organic matter (Figure 6a). Terrestrial/
freshwater-sourced organic material has an average δ13C value of
−25.0‰, compared with marine-sourced organic material with an
average of −21.0‰ (Hu et al., 2006; Yu et al., 2010; Zong et al.,
2006). This suggests that a decrease in terrestrial material input due
to lower freshwater flux is far greater than an assumed increase in
terrestrial material input expected because of increased proximity
associated with delta front advance. This, therefore, suggests that
the actual reduction in terrestrial material input (hence freshwater
flux) is in fact even more significant.
2000 2500 3000 3500 4000 4500 5000 5500600065007000
Age (cal yr BP)
-25.6
-25.2
-24.8
-24.4
δ
13
C
-0.4
-0.2
0
0.2
0.4
0.6
De-trended δ
13
C
20
22
24
26
28
Salinity (PSU)
-8.8
-8.4
-8
-7.6
-7.2
δ
18
O (‰)
-0.6
-0.3
0
0.3
0.6
Detrended δ
18
O
500
490
480
470
Insolation (W/m
2
)
-2
-1
0
1
2
De-trended
10
Be
0
6
12
18
24
Detrended HSG
2000 2500300035004000 4500 500055006000 6500 7000
a
.
Exponentially-smoothed δ
13
C (UV1)
b
.
Diatom-based salinity (UV1), Zong et al., 2010b
c
.
δ
18
O of stalagmites from Dongge Cave, Wang et al., 2005
d
.
Insolation at 30N, Berger and Loutre, 1991
f
.
Detrended
10
Be concentration of GISP2, Finkel and Nishiizumi, 1997
h
.
Detrended δ
18
O of stalagmites from Dongge Cave, Wang et al., 2005
e
.
Detrended δ
13
C from core UV1
g
.
Detrended Hematite-stained grains (HSG), Bond et al., 2001
Wet
Dry
Brackish
Marine
Wet
Dry
Strong
Weak
Wet
Dry
Wet
Dry
Strong
Weak
Warm
Cool
23
4
2? 3?
4?
2? 3? 4?
(‰)
Figure 6. Comparison between d13C of UV1 and other records. (a) Exponentially smoothed d13C from 6650 to 2150 cal. yr BP (this study);
(b) palaeo-salinity based on diatom-salinity transfer function from core UV1 by Zong et al. (2010b); (c) d18O of stalagemite from Dongge Cave,
southern China (Wang et al., 2005). In (a), (b) and (c), dashlines show the mean value of each proxy. (d) Insolation changes at 30°N (Berger and
Loutre, 1991); (e) detrended d13C from core UV1 (this study); (f) detrended concentration of 10Be of GISP2 ice core (Alley et al., 1995; Davis et
al., 1990; Finkel and Nishiizumi, 1997), where higher 10Be concentration is responding to stronger solar activity; (g) detrended hematite-stained
grains (HGS) of Core MC52 and VM29-191 (Bond et al., 2001), with IRD events 2, 3 and 4 (Bond et al., 1997) marked by black bars;
(h) detrended d18O of stalagemite from Dongge Cave, southern China (Wang et al., 2005). In (e) and (h), black bars show the possible position
of IRD events 2, 3, and 4. In (e), grey bars are dry events observed in this study. Detrending of these data was taken using the PAST programme
(Hammer et al., 2001)
712 The Holocene 22(6)
The general decrease in C/N ratios up core (Figure 5) also
supports the interpretation of weakening monsoon strength.
Lower C/N ratios suggest increased proportion of estuarine/
marine particulate organic matter and algae which typically have
C/N values < 10 compared with terrestrial organic matter with
typical C/N values > 10 (Yu et al., 2010). This suggests reduced
terrestrial/freshwater material input due to reduced freshwater
flux, hence reduced monsoon precipitation.
General weakening freshwater flux suggested by δ13C and C/N
is supported by the grain size results of core UV1. Modern samples
from the Pearl River estuary collected by Yu et al. (2010) suggest
that sediments from the riverine area are usually composed of
>30% sand and <50% silt, and 17−18% clay, and those from
brackish/marine areas usually have >60% silt, <20% sand, and
25−28% clay composition. Domination of clay and silt in the sedi-
ment suggests a relatively weak hydrological system under marine/
brackish-water environment at this site during 6600−2200 cal. yr
BP. Increase in silt concentrations from generally < 63.6% before
~4600 cal. yr BP, to > 63.6% after ~4400 cal. yr BP, possibly
indicates a strengthening marine influence at this site (Figure 5),
e.g. changing from brackish to marine conditions responding to
reduced freshwater flux. This is supported by the generally increas-
ing clay concentration during 6600–2200 cal. yr BP.
Diatom flora from core UV1 (Figure 6b) show a general
increase in marine species since 6650 cal. yr BP. This is inter-
preted as indicating weakening monsoon-induced freshwater dis-
charge (Zong et al., 2010b). Wang et al. (2005) examined the
δ18O of stalagmites from the Dongge Cave in southern China, and
revealed a general decreasing summer monsoon during the past
9000 years (Figure 6c). Pollen records from Jiangxi province,
southern China, investigated by Xiao et al. (1998), also show a
clear decrease of evergreen trees that prefer humid conditions
and shrubs, replaced by herb and ferns in the area, indicating
decreasing monsoonal precipitation since early Holocene.
Dry/wet fluctuations at millennial to centennial timescales.
Superimposed on the general weakening trend discussed above
are centennial- and millennial-scale fluctuations in δ13C inter-
preted as wet/dry fluctuations linked to EAM variability during the
mid Holocene (shown by the de-trended δ13C record of Figure 6e).
The δ13C data show one oscillation between 6650 and 4650 cal. yr
BP and another oscillation from 4650 to 2150 cal. yr BP (Figure 6e).
This suggests a possible periodicity for the EAM precipitation
change of around 2000 years. Shorter period, centennial scale, dry
events are also suggested by the δ13C record, for example at 6300,
5700, 4300, 3800, 3400 and 2500 cal. yr BP (shown by horizontal
bars in Figure 6e). Each of these dry events lasts for 100−300
years. Spectral analysis on de-trended δ13C data also suggests two
very significant periodicities of 2514 years and 1408 years
(significance level > 99%), and two significant ones of 765 years
and 476 years (significance level > 95%; Figure 7a).
Various studies have suggested a general weakening trend in
monsoonal climate during the Holocene (e.g. reducing precipita-
tion since the early Holocene), with fluctuations revealed by a
range of proxies such as the Loess deposit in central China (e.g.
An, 2000) and the stalagmite record from Dongge Cave, southern
China (Wang et al., 2005). Zong et al. (2006) examined the bulk
organic δ13C record from the Pearl River estuary and suggested a
strong freshwater flux during 8000–6000 cal. yr BP induced by a
strong summer monsoon. They then identified a gradual reduction
in monsoonal freshwater discharge from 6000 cal. yr BP, with
fluctuations towards the present (Zong et al., 2006). Although the
high-resolution δ18O of stalagmite from Dongge Cave (Wang et al.,
2005) also indicates the general weakening monsoon since the
mid Holocene (Figure 6c), it does not pick up centennial and
millennial fluctuations found in δ13C signal from the Pearl River
estuary (Figure 6e and h). This is probably because the δ18O of
stalagmite can be influenced by both temperature and precipitation,
and δ18O of the stalagmite is thought to mainly reflect more local
environmental changes (e.g. Kelly et al., 2006) while δ13C likely
shows broader catchment area change.
Driving mechanisms for the EAM variability during
the mid Holocene
Orbital forcing. Orbital forcing, specifically the precession
cycle, has been suggested as the most important driving mecha-
nism for EAM precipitation variability during the Holocene (An,
2000; An and Thompson, 1998; Huang et al., 2008; Kutzback,
1981; Mayewski et al., 2004). This study shows a general weak-
ening precipitation during the mid Holocene following part of the
precession cycle (Figure 6a and d). Precession influences the
monsoon through its impact on seasonal contrasts. Half a preces-
sion cycle in the past (at approximately the beginning of the Holo-
cene), perihelion occurred in the boreal summer, causing enhanced
summer insolation in the Northern Hemisphere, and aphelion fell
in the boreal winter, causing reduced winter insolation. Such sea-
sonal contrast was strongest at the beginning of the Holocene and
has weakened towards present during the current precession cycle
(Figure 6d; Berger and Loutre, 1991; Bradley, 2003). Today, in its
slightly elliptical orbit, the Earth is at perihelion around the boreal
winter solstice, and at aphelion around the boreal summer solstice
(Bradley, 2003). Berger and Loutre (1991) calculated the insola-
tion values at different latitude zones in both summer and winter,
showing weakening summer insolation at 30°N (Figure 6d). The
weakening East Asian summer monsoon during the 6650−2150
00.01 0.02 0.03 0.04 0.05 0.06
Frequency
0
5
10
15
20
25
30
Power
00.005 0.01 0.015 0.02 0.0250.03
Frequency
0
20
40
60
80
100
120
140
160
Power
2514 yrs
1408 yrs
765 yrs 476 yrs
2475 yrs
1125 yrs
α=0.01
α=0.05
α=0.01
α=0.05
ab
Figure 7. Spectral analysis of (a) d13C of Core UV1 (this study) and (b) 10Be of GISP2 ice core (Alley et al., 1995; Davis et al., 1990; Finkel
and Nishiizumi, 1997). Exponentially smoothed d13C and 10Be data have been de-trended prior to spectral analysis using the PAST program
(Hammer et al., 2001). Dashlines show significance levels of 99% (d=0.01) and 95% (d=0.05), respectively
Yu et al. 713
cal. yr BP identified here is most likely a response to the orbital-
induced weakening insolation since about 11 000 years BP (An,
2000; Wang et al., 2005).
There are many possible ways that insolation could influence
the EAM system. One possible process is that reducing insolation
leads to changes in oceanic and atmospheric circulation, resulting
in changes in temperature and precipitation (An and Thompson,
1998; Huang et al., 2008; Xiao et al., 2006). For example, reduc-
ing insolation results in cooler tropical temperature and thus
resulted in a weakened Hadley circulation (Rind and Overpeck,
1993) and reduced water vapour content of the atmosphere.
Thereby it would reduce the monsoonal precipitation and resul-
tant freshwater runoff. The influence of precession on the EAM is
also supported by a variety of other proxies, e.g. pollen records
from the desert/loess transition of north central China (e.g. Jiang
et al., 2006) and loess deposits from central China (An et al.,
2000; Kukla et al., 1990). Owing to the short record from UV1,
this study only demonstrates EAM changes following part of the
precession cycle during the mid Holocene.
Solar modulation. Solar activity is one of the possible control-
ling factors for EAM changes at the millennial timescale. One
proxy used to reflect solar modulation (cosmic ray flux) is 10Be
concentration in ice cores (Masarik and Beer, 1999; Raisbeck et al.,
1990). Studies show that high 10Be concentrations can be attrib-
uted to lower solar shielding and consequently to a lower solar
activity (Finkel and Nishiizumi, 1997; Muscheler et al., 2004).
Horiuchi et al. (2008) analysed an ice core at the Dome Fuji sta-
tion, inland Antarctica, and found that both the concentration and
the flux of 10Be increased at the known solar-activity minimums
during the past millennia, e.g. Oort (1010−1050 yr CE) and Wolf
(1280−1340 yr CE). Finkel and Nishiizumi (1997) suggest that
oscillations in the 10Be concentration of the GISP2 ice core indi-
cate changes in cosmic ray flux and thus (activity of) solar output,
as production rates at geomagnetic latitudes above 60° are almost
completely insensitive to changes in geomagnetic field intensity
(Lal and Peters, 1967). Spectral analysis on 10Be (GISP2 ice core)
indicates two significant periodicities of 2475 years and 1125
years (significance level > 99%; Figure 7b), which is similar to
the two periodicities (2514 and 1408 years) derived from δ13C
(Figure 7a). The similar periodicities at millennial scale, together
with a good correlation between de-trended δ13C signal (core
UV1) and de-trended 10Be concentration (GISP2, Finkel and
Nishiizumi, 1997) (Figure 6e and f) indicate that solar irradiation
might be a driving mechanism of EAM variability at millennial
and centennial timescales (e.g. Chen et al., 2008).
Although studies have hypothesized several possible correla-
tions between solar activity and the Earth’s climate, the exact pro-
cess controlling how the EAM responds to insolation change still
remains unclear. Haigh (1996) suggested that strong solar output
increases UV radiation, resulting in an increase in stratospheric O3.
The stratospheric O3 absorbs radiation and heats up the strato-
sphere, and the heat propagates down into the troposphere. How-
ever, whether the stratospheric heat is strong enough to pass down
through the atmosphere far below still needs further investigation.
Carslaw et al. (2002) proposed the ‘cosmic ray-cloud effect’ to
explain the link between solar activity and atmospheric processes.
Cosmic rays are charged particles with very high energy. Charged
particles can move freely along with the magnetic field lines, but
not when its track is transverse to the field lines. On their way to
the Earth, cosmic rays must cross through interplanetary (IP) mag-
netic fields, and thus, the stronger IP magnetic field will result in
less cosmic rays arriving at the Earth. The strength of the IP mag-
netic field at solar maxima is almost 1.5 times that at solar minima
(Yuming Wang and Tengfei Zhang, personal communication,
2011). Therefore, stronger solar activity will result in fewer cosmic
rays reaching the Earth, which is then correlated with a reduction
in cloud cover (Parker, 1958). One possible mechanism for the
cosmic ray-cloud effect is suggested by Yu and Turco (2001). Cos-
mic rays, with their high energy, are able to ionize the atmospheric
molecules during the collision, which results in charged molecular
clusters. These clusters then work as the condensation nuclei for
the cloud formation. Thus shortage of cosmic rays during periods
of strong solar activity will discourage cloud formation, and result
in higher solar radiation reaching the Earth, resulting in higher
Earth surface temperature (Carslaw et al., 2002). As a high 10Be
concentration has been found to be related to periods of low solar
activity during past millennia at centennial timescales (e.g. Finkel
and Nishiizumi, 1997; Horiuchi et al., 2008; Muscheler et al.,
2004), the Earth surface temperature would be low during periods
of low solar activity indicated by high concentration of 10Be in the
GISP2 ice core (Figure 6e; Finkel and Nishiizumi, 1997). Thus, it
is possible that at the millennial timescale, solar activity might
influence the EAM through changing the Earth surface tempera-
ture, and resultant changes in strength of the Siberian High. The
Siberian High becomes strong during periods of low solar activity,
resulting in strong winter monsoons of low precipitation (e.g. An,
2000). For example, dry and wet periods indicated by de-trended
δ13C signal (Figure 6e) are coherent with low and high solar
activity indicated by 10Be concentration (Figure 6f).
Solar forcing seems to explain most of the fluctuations
observed in this study, however, correlation between EAM and
insolation breaks down from 4500 cal. yr BP to 3500 cal. yr BP.
Strengthened solar activity around 3930 cal. yr BP indicated by
low 10Be concentration (Figure 6f), might result in the wet event
around 3990 cal. yr BP indicated by high δ13C (Figure 6e). Solar
activity the weakens gradually until c. 3300 cal. yr BP (indicated
by increasing 10Be concentration, Figure 6f), and correspondingly,
the EAM weakens generally.
Other forcing. Decadal to centennial fluctuations in EAM might
be due to other forcing, such as high-latitude IRD events (Figure 6g
and h). Changes in solar activity have been suggested as a pos-
sible forcing mechanism for millennial-scale cooling events in the
Northern Atlantic area, such as the North Atlantic ice-rafted
debris (IRD) events (Bond et al., 2001). Such correlation is sup-
ported by comparison of de-trended 10Be (Figure 6f) and de-
trended hematite-stained grain concentration (HSG, Bond et al.,
2001) in the North Atlantic sediments (Figure 6g). These changes
in the North Atlantic lead to enhanced strength of the Siberian
High (Bianchi and McCave, 1999; Rohling et al., 2002) which, in
turn, can lead to reduced EAM precipitation. Connection between
the IRD events and the EAM is supported by good consistence
between de-trended HSG (Figure 6g, Bond et al., 2001) and de-
trended δ13C (Figure 6e). For example, three key IRD events from
the North Atlantic during the mid Holocene, at 5900−5800,
4300−4100 and 2800−2700 cal. yr BP (Bond et al., 1997; black
bars numbered 2, 3 and 4 in Figure 6g) appear to correlate with
dry events observed in the Pearl River area, at around 5700, 4300
and 2700 cal. yr BP (black bars in Figure 6e). Dry events observed
in UV1 are slightly delayed compared to IRD events suggested by
Bond et al (2001). This might be due to the time lag between
vegetation change and climate change (likely to be 100–200
years), and the UV1 record is based on vegetation signal.
However, in addition to the dry events addressed above (black
bars in Figure 6e), there are other dry events lasting 100−300
years suggested by δ13C (grey bars in Figure 6e). These addi-
tional dry events seem to be consistent with low solar activity
(Figure 6f), while none of them is correlated with IRD events.
These additional dry events might be linked to some regional
forcing such as ENSO and possibly high-frequency and
low-amplitude changes in the sea level.
714 The Holocene 22(6)
Correlation between high-latitude cooling events and the
EAM at centennial timescales is supported by the δ18O record
from stalagmites (Figure 6h, Wang et al., 2005). Wang et al.
(2005) suggested good correlation between EAM and IRD events
during the Holocene, suggesting a possible connection between
EAM and high-latitude cooling events. However, comparison
between the Pearl River estuary δ13C record, and the δ18O from
the Dongge Cave, reveals some obvious mismatches. The δ13C
from UV1 indicates two significant dry periods around 5700 and
3500 cal. yr BP, respectively (Figure 6a), while δ18O from Dongge
Cave does seem to suggest a cool period around 5700 cal. yr BP
and no obvious change around 3500 cal. yr BP (Figure 6c). It is
particularly noticeable that the correlation between our Pearl
River estuary record (Figure 6e) and the Dongge Cave stalagmite
record (Figure 6h) at a millennial timescale seems to be weak
compared with the correlation between the Pearl River estuary,
10Be concentration and HSG records (Figure 6e, f and g). Com-
parison between δ13C and 10Be suggests the two dry periods
(~5700 and ~3500 cal. yr BP) might be due to low solar activity
(Figure 6e and f), while the δ18O stalagmite record might be more
sensitive to local scale variability and thus does not catch up these
changes (Figure 6h).
Conclusion
Results from this study suggests that the East Asian Monsoon
experienced a general weakening between 6650 and 2150 cal. yr
BP. Superimposed on this trend are wet/dry oscillations at millen-
nial to centennial timescales, with major millennial periodicities
of about 2500 years and 1400 years, and centennial periodicities
of 770 years and 480 years. Our data support the view that
orbital-induced precession forcing is the primary controlling
mechanism for the monsoonal variability during the Holocene,
while solar modulation is a significant forcing for wet/dry
oscillations at millennial timescales. Solar activity influences the
EAM by changing the strength of Siberian High at millennial
timescales (also influencing the North Atlantic through IRD
events). Results of this study are supported by other records,
especially the high-resolution stalagmite record from Dongge
Cave (Wang et al., 2005). However, mismatches between millen-
nial cycles from the δ13C record presented here and the Dongge
Cave δ18O stalagmite record suggest the Dongge Cave might
reflect more local conditions, while δ13C presented here reflects
broader climate conditions.
Acknowledgements
We thank two anonymous reviewers for valuable comments and
suggestions for improvement of this manuscript. We acknowl-
edge staff at the RCL and NIGL for their analytical support.
Funding
This research was part of the PhD project sponsored by NERC/
EPSRC (UK) through the Dorothy Hodgkin Postgraduate Award
(to FY). This research was also supported by the University of
Durham through a special research grant (to YZ), the NERC Ra-
diocarbon Laboratory Steering committee (1150.1005) (to YZ)
and the NERC Isotope Geosciences Facilities Steering Commit-
tee (IP/883/1105) (to YZ) and a research grant from the Re-
search Grant Council Hong Kong (HKU707109P) (to YZ). This
research was also supported by the Research Grants Council of
the Hong Kong SAR through research grants HKU7058/06P and
HKU7052/08P (to W.W.-S. Yim). We also acknowledge support
from the University College (Durham University) to complete
fieldwork and laboratory visits.
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