CPClimate of the PastCPClim. Past1814-9332Copernicus PublicationsGöttingen, Germany10.5194/cp-14-1869-2018Evaluating the timing and structure of the 4.2 ka event in the Indian
summer monsoon domain from an annually resolved speleothem record from
Northeast IndiaEvaluating the timing and structure of the 4.2 ka eventKathayatGayatrikathayat@xjtu.edu.cnChengHaicheng021@xjtu.edu.cnSinhaAshishBerkelhammerMaxhttps://orcid.org/0000-0002-8924-716XZhangHaiweiDuanPengzhenLiHanyingLiXiangleiNingYoufengEdwardsR. LawrenceInstitute of Global Environmental Change, Xi'an Jiaotong
University, Xi'an, ChinaDepartment of Earth Sciences, University of Minnesota, Minneapolis,
USADepartment of Earth Science, California State University Dominguez
Hills, Carson, USADepartment of Earth and Environmental Sciences, University of
Illinois, Chicago, USAGayatri Kathayat (kathayat@xjtu.edu.cn) and Hai Cheng (cheng021@xjtu.edu.cn)30November201814121869187923July201816August201812November201817November2018This work is licensed under the Creative Commons Attribution 4.0 International License. To view a copy of this licence, visit https://creativecommons.org/licenses/by/4.0/This article is available from https://cp.copernicus.org/articles/14/1869/2018/cp-14-1869-2018.htmlThe full text article is available as a PDF file from https://cp.copernicus.org/articles/14/1869/2018/cp-14-1869-2018.pdf
A large array of proxy records
suggests that the “4.2 ka event” marks an approximately
300-year long period (∼3.9 to 4.2 ka) of
major climate change across the globe. However, the climatic manifestation of
this event, including its onset, duration, and termination, remains less
clear in the Indian summer monsoon (ISM) domain. Here, we present new oxygen
isotope (δ18O) data from a pair of speleothems (ML.1 and ML.2)
from Mawmluh Cave, Meghalaya, India, that provide a high-resolution record of
ISM variability during a period (∼3.78 and 4.44 ka) that fully
encompasses the 4.2 ka event. The sub-annually to annually resolved ML.1
δ18O record is constrained by 18 230Th dates with an
average dating error of ±13 years (2σ) and a resolution of ∼40 years, which allows us to characterize the ISM variability with
unprecedented detail. The inferred pattern of ISM variability during the
period contemporaneous with the 4.2 ka event shares broad similarities and
key differences with the previous reconstructions of ISM from the Mawmluh
Cave and other proxy records from the region. Our data suggest that the ISM
intensity, in the context of the length of our record, abruptly decreased at
∼4.0 ka (∼±13 years), marking the onset of a multi-centennial
period of relatively reduced ISM, which was punctuated by at least two
multi-decadal droughts between ∼3.9 and 4.0 ka. The latter stands out
in contrast with some previous proxy reconstructions of the ISM, in which the
4.2 ka event has been depicted as a singular multi-centennial drought.
Location map and spatial structure of mean JJAS precipitation and
low-level winds. (a) June–July–August–September (JJAS)
precipitation from the Tropical Rainfall Measuring Mission (TRMM). The
locations of Mawmluh Cave (white circle) and other proxy records mentioned in
the text (yellow circles and numbers). The numbering scheme is as follows: 1,
Sahiya Cave (Kathayat et al., 2017); 2, Lake Rara (Nakamura et al., 2016); 3,
Kotla Dhar (Dixit et al., 2014); 4, Mawmluh Cave (Berkelhammer et al., 2012);
and 5, Indus Delta (Staubwasser et al., 2003). (b)The 850 hPa
monsoon vector from zoomed Laboratoire de Meteorologie Dynamique (LMDZ)
general circulation model with telescoping zooming (figure adapted and
modified from Sabin et al., 2013). The zoom version shows a well-defined
cyclonic circulation with westerlies on the southern flanks and easterly
winds on the northern flanks of the monsoon trough. The Mawmluh Cave is
ideally located to record upstream variations in the overall strength of the
ISM (see text).
Introduction
The time interval between 4.2 and 3.9 ka (thousand of years before present,
where present is 1950 CE) constitutes an important period from both
climatological and archeological perspectives (e.g., Weiss et al., 1993;
Cullen et al., 2000; Staubwasser et al., 2003; Berkelhammer et al., 2012;
Weiss, 2016). A global suite of proxy records shows widespread climate
anomalies during the time (commonly referred to as the “4.2 ka event”; e.g.,
Cullen et al., 2000; Staubwasser et al., 2003; Arz et al., 2006; Drysdale et
al., 2006; Menounos et al., 2008; Liu and Feng, 2012; Berkelhammer et al.,
2012; Dixit et al., 2014, 2018; Cheng et al., 2015; Nakamura et al., 2016;
Railsback et al., 2018). Additionally, a number of archeological studies also
suggest that the 4.2 ka event was associated with a series of cultural
and societal changes in the Mediterranean, Middle East, Africa, and South and
East Asia (e.g., Weiss et al., 1993; Enzel et al., 1999; Cullen et al., 2000;
Staubwasser et al., 2003; Marshall et al., 2011; Liu and Feng, 2012; Dixit et
al., 2014; Weiss, 2016). For example, the 4.2 ka event has been proposed
to have contributed to collapses of the early Bronze Age civilizations,
including the Longshan culture in China (Chang, 1999; Liu and Feng, 2012),
the Egyptian Old Kingdom by the Nile River (Stanley et al., 2003), and the
Akkadian Empire in Mesopotamia (Weiss et al., 1993; Cullen et al., 2000). In
South Asia, the 4.2 ka event has been linked to a weakening of the Indian
summer monsoon (ISM) and the ensuing de-urbanization of the Indus Valley
Civilization (Staubwasser et al., 2003; Madella and Fuller, 2006; Dixit et
al., 2014, 2018; Giosan et al., 2012; Berkelhammer et al., 2012; Kathayat et
al., 2017).
230Th dating results with the 2σ analytical
error.
ML.1 Sample238U232Thd234Ua230Th/238U230Th age (yr)d234UInitialb230Th age (yr BP)cnumber(ppb)(ppt)(measured)(activity)(corrected)(corrected)(corrected )ML.1.1F6106±6164±7-274.2±0.90.0251±0.00013841±14-277±13779±14ML.1.2F6222±6123±5-272.3±0.90.0256±0.00013917±13-275±13855±13ML.1.3F6981±972±4-271.8±1.00.0257±0.00013923±12-275±13861±12ML.1.4F6378±7170±6-270.3±0.90.0258±0.00013938±14-273±13876±14ML.1.5F6674±8111±4-271.5±0.90.0259±0.00013948±11-275±13886±11ML.1.6F7702±10189±5-270.1±1.00.0260±0.00013964±11-273±13902±11ML.1.7F6455±790±6-270.7±0.90.0260±0.00013968±14-274±13906±14ML.1.8F6144±7153±5-270.3±1.00.0261±0.00013978±13-273±13916±13ML.1.9F6363±7122±6-270.0±1.00.0262±0.00013989±15-273±13927±15ML.1.10F6825±9778±16-271.5±1.00.0264±0.00014031±12-275±13969±12ML.1-9a6395±6154±5-268.7±0.90.0267±0.00014069±13-272±14007±13ML.1-10a7574±8132±4-267.8±0.90.0272±0.00014138±12-271±14076±12ML.1-11a6744±652±3-266.2±0.80.0279±0.00014240±11-269±14178±11ML.1-12a7716±873±3-265.1±0.90.0286±0.00014336±11-268±14274±11ML.1-13a7881±9127±4-262.9±1.10.0290±0.00014386±13-266±14324±13ML.1-14a6452±977±2-263.8±1.10.0294±0.00014451±14-267±14389±14ML.1-15a7392±10108±3-263.4±1.00.0294±0.00014456±11-267±14394±11ML.1-16a6970±9441±9-263.5±1.00.0297±0.00014499±12-267±14437±12ML.2 ML.2-76633±952±3-277.0±1.10.0231±0.00013541±16-280±13479±16ML.2-86173±778±3-272.6±1.00.0254±0.00013891±13-276±13829±13ML.2-97121±8134±4-266.0±1.00.0282±0.00014276±13-269±14214±13ML.2-106085±62953±59-262.1±1.00.0299±0.00014500±21-265±14438±21ML.2-9a6278±6106±4-257.5±1.00.0306±0.00014603±17-261±14541±17
aδ234U=([234U/238U]activity-1)×1000.
bδ234Uinitial was calculated based on
230Th age (T), i.e., δ234Uinitial=δ234Umeasured×el234×T. Corrected
230Th ages assume the initial 230Th/232Th atomic ratio
of 4.4±2.2×10-6. Those are the values for a material at
secular equilibrium, with the bulk earth 232Th/238U value of
3.8. The errors are arbitrarily assumed to be 50 %. c BP
stands for “before present” where the “present” is defined as the year
1950 CE.
Proxy records from the Indian subcontinent. The select proxy records
from the Indian monsoon domain from the top are Kotla Dhar (Dixit et al.,
2014), Lake Rara (Nakamura et al., 2016), Lonar Lake (Sarkar et al., 2015),
Indus Delta (Staubwasser et al., 2003), Mawmluh Cave (Berkelhammer et al.,
2012; purple) and this study (orange), and Sahiya Cave (Kathayat et al., 2017).
The yellow bar delineates the commonly accepted temporal duration of the
4.2 ka event (e.g., Weiss, 2016).
A number of proxy records from the Indian subcontinent suggest that a major
weakening of the ISM occurred around the 4.2 ka event (Staubwasser et
al., 2003; Berkelhammer et al., 2012; Dixit et al., 2014; Nakamura et al.,
2016; Kathayat et al., 2017; Figs. 1 and 2). The 4.2 ka event has been
generally described as an approximately 2- to 3-centuries-long
interval of drought (e.g., Berkelhammer et al., 2012; Dixit et al., 2014;
Nakamura et al., 2016), which was superimposed on a longer-term
insolation-induced weakening of the ISM during the Holocene (e.g., Kathayat
et al., 2017). The timing, structure, and magnitude of the 4.2 ka event in
the ISM regime, however, remain unclear because most proxy records from the
region have low temporal precision and insufficient resolution to precisely
characterize the event (e.g., Staubwasser and Weiss, 2006; Prasad and Enzel,
2006; Nakamura et al., 2016; Dixit et al., 2018). In addition, the 4.2 ka
event is notably absent in a recent high-resolution speleothem oxygen isotope
(δ18O) record from Sahiya Cave in northern India (Kathayat et
al., 2017) that exhibits a long-term drying trend from ∼4.2 to
3.5 ka.
A high-resolution (∼6 years) δ18O record (KM-A) from
Mawmluh Cave, located in the state of Meghalaya in Northeast India, has
previously provided evidence of the 4.2 ka event from the ISM domain
(Berkelhammer et al., 2012). The KM-A record was recently used to formally
ratify the post-4.2 ka time as the Meghalayan Age (Walker et al.,
2018). However, the timing and duration of the 4.2 ka event in the KM-A
record is constrained by only three 230Th dates (5048±32,
4112±30, and 3654±20 ka) and, additionally, the youngest date
defining the termination of the event and/or the δ18O values
from the top ∼30 mm of the KM-A sample that help define the event may
have been potentially affected by diagenetic changes. In this study, we
present new high-resolution δ18O data from two stalagmites
(ML.1 and ML.2) from the same cave (Figs. 1 and 3, Table 1). The ML.1 and
ML.2 δ18O records span from 4.44 to 3.78 ka and 4.53 to
3.70 ka, respectively, encompassing the 4.2 ka event completely. Our
new records are sub-annually to annually (ML.1) and sub-decadally resolved
(ML.2) and have unprecedented chronologic constraints, which allow us to
characterize the nature of ISM variability during the 4.2 ka event more
precisely than previously possible.
Samples and methodsCave location and climatology
Mawmluh Cave (25∘15′32′′ N, 91∘42′45′′ E;
1290 m a.s.l.) is located near the town of Sohra (Cherrapunji) at the
southern fringe of the Meghalayan Plateau in Northeast India (Fig. 1). The
mean annual rainfall is ∼11 000 mm in the region, 70 % of which
falls during the peak ISM months (June–September; Murata et al., 2007). The
rainfall at the cave site during the ISM period is mainly produced by
convective systems and low-level air parcels originating from the Bay of
Bengal, which propagate further northward and penetrate farther into the
Tibetan Plateau (Sengupta and Sarkar, 2006; Breitenbach et al., 2010). The
non-monsoonal component of rainfall is trivial and consists of westerly-related moisture and recycled local moisture (Breitenbach et al.,
2010, 2015; Berkelhammer et al., 2012). The cave is overlain by 30–100 m
thick and heavily karstified host rock (limestone, sandstone, and a
40–100 cm thick coal layer; Breitenbach et al., 2010). The soil layer
above the cave is rather thin (5–15 cm) and covered mainly by grasses and
bushes. Cave monitoring data (Breitenbach et al., 2010) indicate that the
relative humidity inside the cave is more than 95 % even during the dry
season (November to April). Temperature variations in the cave are small
(18.0–18.5 ∘C) and close to the mean annual temperature of the area
(Breitenbach et al., 2010, 2015). A 3-year cave monitoring result suggests
that cave drip-water δ18O signals lag corresponding local
rainfall by less than 1 month and thus preserve seasonal signals of ISM
rainfall (Breitenbach et al., 2010). Previous studies have indicated that
variations in the δ18O of speleothem calcite from Mawmluh Cave
reflect changes in the amount-weighted δ18O of precipitation
(δ18Op) values (Breitenbach et al., 2010, 2015;
Berkelhammer et al., 2012; Myers et al., 2015). The ML.1 and ML.2 samples
from Mawmluh Cave were collected in November 2015 at ∼4–5 m above the
cave floor and ∼700 m from the cave entrance. Diameters of ML.1 and
ML.2 are ∼170 and 165 mm, with lengths ∼315 and ∼311 mm,
respectively. Both stalagmite samples were cut along their growth axes using
a thin diamond blade. There are no visible changes in the texture or hiatuses
in the above sample intervals that we used for this study (Fig. 3).
Samples photograph: the total length of ML.1 and ML.2 samples is 315
and 311 mm, respectively. The arrows indicate the dating subsampling
location and the 230Th dates with the 2σ analytical error
(see also Tables 1 and S1). The centimeter scale indicates the location of isotopic
measurements, enclosing the interval of interest within both the samples.
230Th dating
We obtained 18 and 5 230Th dates for samples ML.1 and ML.2,
respectively. Subsamples for 230Th dating (∼30 mg) were
drilled from ML.1 and ML.2 by using a 0.5 mm carbide dental drill. The
230Th dating was performed at Xi'an Jiaotong University, China,
using Thermo-Finnigan Neptune Plus multi-collector inductively
coupled plasma mass spectrometers (MC-ICP-MSs). The method is described in
Cheng et al. (2000, 2013). We used standard chemistry procedures (Edwards et
al., 1987) to separate uranium and thorium. A triple-spike
(229Th–233U–236U) isotope dilution method was
used to correct instrumental fractionation and to determine U/Th
isotopic ratios and concentrations (Cheng et al., 2000, 2013). U and Th
isotopes were measured on a MasCom multiplier behind the retarding potential
quadrupole in the peak-jumping mode using the standard procedures described
in Cheng et al. (2000). Uncertainties in U/Th isotopic measurements
were calculated offline at 2σ level, including corrections for blanks,
multiplier dark noise, abundance sensitivity, and contents of the same
nuclides in spike solution. The U decay constants are reported in Cheng et
al. (2013). Corrected 230Th ages assume the initial
230Th/232Th atomic ratio of 4.4±2.2×10-6, the
values for material at secular equilibrium with the bulk earth
232Th/238U value of 3.8. The corrections are small because the
uranium concentrations of the samples are high (∼6 ppm) and detrital
232Th components are low (average <170 ppt; Tables 1 and S1 in the Supplement).
Age models of ML.1 and ML.2 records. We adopted COPRA and generated
2000 realizations of age models to account for the dating uncertainty
(2.5 % and 97.5 % quantile) confidence
limits. (a) ML-1 age models and modeled age uncertainties using
three
different age modeling algorithms: COPRA (black) (Breitenbach et al., 2012),
Bchron (purple) (Haslett and Parnell, 2008), and ISCAM (red) (Fohlmeister,
2012). The gray band depicts the 95 % confidence interval using COPRA.
Error bars on 230Th dates represent a 2σ analytical error.
(b) ML.2 age model and modeled age uncertainties using COPRA.
Age models
The ML.1 age models and associated age uncertainties were constructed using
the COPRA (Constructing Proxy Records from Age; Breitenbach et al., 2012),
Bchron (Haslett and Parnell, 2008), and ISCAM (Fohlmeister, 2012) age modeling
schemes (Fig. 4). All three modeling schemes yielded nearly identical results
and the conclusions of this study are then not sensitive to the choice of
different age models (Fig. 4). The ML.2 age model and associated
uncertainties were constructed by only using the COPRA age modeling scheme
(Breitenbach et al., 2012; Fig. 4).
Stable isotope analysis
The ML.1 and ML.2 δ18O records are established by ∼970
and ∼238 stable isotope measurements, respectively (Figs. 5, 6 and
Table S2). Subsamples for stable isotope measurements were obtained from ML.1
and ML.2 between depths of 125–250 and 182–255 mm (depth from the top),
respectively. Accordingly, we report our data with zero depths set at 125 and
182 mm from the top of stalagmites ML.1 and ML.2, respectively (Fig. 3). We
used New Wave Micromill, a digitally controlled triaxial micromill
instrument, to obtain the subsamples. The sample growth rates were determined
by sample age models, which in turn, were used to determine the subsampling
increments (typically between 50 and 100 µm) for attaining similar
temporal resolutions throughout the sample (typically ∼1 year for the
ML.1 δ18O record). The δ18O and
δ13C were measured using a Finnigan MAT-253 mass spectrometer
coupled with an online carbonate preparation system (Kiel-IV) in the Isotope
Laboratory, Xi'an Jiaotong University. Results are reported in per mil
(‰) relative to the Vienna Pee Dee Belemnite (VPDB) standard.
Duplicate measurements of standards NBS19 and TTB1 show a long-term
reproducibility of ∼0.1 ‰ (1σ) or better (Figs. 5, 6
and Table S2)
Comparison between ML.1 and ML.2 δ18O profiles over
the period of overlap. The ML.1 (a) and ML.2 (b) profiles
are on their independent age models. The circles with horizontal error bars
depict 230Th dates and errors (2σ; see also Tables 1, S1,
and S2). (c) Comparison between the ML.1 and ML.2
δ18O profiles based on the ISCAM algorithm (Fohlmeister, 2012).
The δ18O and δ13C profiles of ML.1 and
ML.2. (a) The ML.1 δ18O (orange) and
δ13C (green). (b) The ML.2 δ18O
(purple) and δ13C (blue) on their independent age models. The
Pearson correlation (r) and its 95 % confidence interval together with
actual and effective sample size (after considering autocorrelation in each
profile) are shown in the figure.
Replication and isotopic equilibrium
Excellent replication between the ML.1 and ML.2 δ18O profiles
(Fig. 5) suggests that the precipitation of speleothem calcite in Mawmluh Cave
essentially occurred at or near isotopic equilibrium conditions and the
speleothem δ18O records primarily reflects the meteoric
precipitation δ18O variations (Dorale et al., 1998; Wang et
al., 2001). A high degree of replication has been argued as a definitive test
of isotopic equilibrium. This is because if the records replicate, the effect
of additional kinetic and/or vadose-zone processes on the calcite δ18O must have been either absent or exactly the same for spatially
separated stalagmites. Principally, each speleothem–drip-water pair can have
a distinctive combination of flow path, CO2 content, residence time,
solute concentrations, and prior calcite precipitation (PCP) history in the
soil zone and epikarst above cave. Thus, the replication of different
speleothem records suggests that such additional processes are not crucial.
We assessed the degree of replication between ML.1 and ML.2 δ18O records by using the ISCAM (intra-site correlation age modeling)
algorithm (Fohlmeister, 2012). The ISCAM finds the best correlation between
proxy records within the combined age uncertainties of two records by using a
Monte Carlo approach. Significant levels were calculated against a red-noise
background from 1000 pairs of artificially simulated first-order
autoregressive time series (AR1). The ML.1 and ML.2 δ18O time
series on ISCAM-derived age models display a statistically significant
correlation (r=0.58 at 95 % confidence level) over their contemporary
growth period between ∼4.4 and 3.8 ka.
Results
The average 230Th dating
uncertainties of the ML.1 and ML.2 records are ±13 and ±16 years,
respectively (Fig. 3, Tables 1 and S1). Temporal resolutions of the ML.1
δ18O record range from ∼0.1 to ∼3 years with an
average resolution of ∼1 year. All dates are in stratigraphic order
within dating uncertainties. Of note, 9 ML.1 230Th dates were
obtained between 27 and 88 mm depths (i.e., about one date every 7 mm),
covering the interval from 4.2 to 3.9 ka, the typical time range of the
4.2 ka event. The ML.1 and ML.2 δ18O values range between
-6.6 and -4.8 ‰ with mean values of -5.80 and
-5.43 ‰, respectively (Fig. 5). The average temporal resolution of
the ML.2 record is ∼5 years (Fig. 5). The δ13C values in
ML.1 and ML.2 range between -2.8 ‰ and 1.0 ‰ with
mean values of -1.0 ‰ and -0.8 ‰, respectively (Fig. 6).
Discussion and conclusionsProxy interpretations
The temporal variability in ISM δ18Op and consequently
speleothem δ18O in the study area has been well studied
previously and attributed mainly to changes in spatially integrated upstream
rainfall at cave sites (e.g., Sinha et al., 2011a; Breitenbach et al., 2010,
2015; Berkelhammer et al., 2012; Kathayat et al., 2016; Cheng et al., 2016).
A number of model simulations with isotope-enabled general circulation models
(GCMs) also suggest a significant inverse relationship between upstream ISM
rainfall amount and the δ18Op variations over the Indian
subcontinent (e.g., Vuille et al., 2005; Pausata et al., 2011; Berkelhammer
et al., 2012; Sinha et al., 2015; Midhun and Ramesh, 2016). Following these
reasonings, we interpret the low and high δ18O values in our
records to reflect strong and weak ISM, respectively (e.g., Dayem et al.,
2010; Sinha et al., 2011b, 2015; Cheng et al., 2012; Berkelhammer et al.,
2012; Breitenbach et al., 2015; Myers et al., 2015; Kathayat et al., 2016,
2017). Climatic interpretations of the speleothem δ13C signal, however, are
more complex because the δ13C variations can be
driven by climatic changes and non-climate-related local processes
(Baker et al., 1997; Genty et al., 2003; Fairchild and Treble, 2009;
Fohlmeister et al., 2011; Deininger et al., 2012; Scholz et al., 2012). A
moderate to strong covariance between the ML.1 and ML.2 δ13C
and δ18O profile values (r=0.49 and 0.66,
respectively) suggests that both proxies reflect a common response to changes
in the local hydrology of the region; however, we cannot rule out non-climate-related factors in producing this observed relationship. Consequently, the
interpretative framework used in this study is mostly based on the speleothem
δ18O variability.
Comparison between the KM-A, ML.1, and ML2 δ18O
profiles: (a) an image of KM-A stalagmite (Berkelhammer et al.,
2012). The yellow dots indicate three 230Th dates. The black curve
marks the potential dissolution surface. The white aragonite layer above the
dissolution surface was deposited after the 1950s with the advent of
limestone mining above the Mawmluh Cave (Breitenbach et al., 2010).
(b) The dotted lines delineate the portion of the KM-A
δ18O record (red; ∼4.4 ka to 3.654 ka)
(Berkelhammer et al., 2012) discussed in the text. (c) The ML.2
δ18O profile (blue) (this study) is overlaid by a 6-year
interpolated ML.1 δ18O profile (green) (this study; see also
Table S2). The horizontal error bars (red, green, and blue) on the
230Th dates represent a 2σ analytical error. The vertical
gray bar indicates the inferred duration of weakest (driest phase) ISM as
indicated by the KM-A and ML δ18O records. The yellow bar
indicates the interval of anomalously depleted δ18O values in
the
KM-A record.
Comparisons between the KM-A and ML.1–ML.2 δ18O
records
The 4.2 ka event in the KM-A record (Berkelhammer et al., 2012)
manifests as a two-step change marked by an initial increase in the
δ18O values (∼0.6 ‰) between ∼4.31 and
4.30 ka, followed by another abrupt increase between ∼4.07 and
4.05 ka. The period between 4.05 and 3.87 ka in the KM-A profile is
characterized by the most enriched δ18O values over the entire
record (∼1.5 ‰ higher than the background values before the
event; Fig. 7), delineating ∼180 years of substantially weaker ISM.
This multi-centennial period of enriched δ18O values was
terminated abruptly by a sharp return (<20 years) to depleted
δ18O values, implying a resumption of stronger monsoon. The
ML.1 and ML.2 δ18O profiles during the contemporaneous period
with the KM-A record, however, exhibit no step-like increase around ∼4.3 ka but instead an abrupt increase in the δ18O
values at ∼4.01 ka, which are superimposed over a gradually
increasing trend over the entire length of the records. The timings and
magnitude of this abrupt increase in the δ18O values in both
the ML.1 and ML.2 profiles are comparable to those observed in the KM-A profile
(within the combined age uncertainties of both records; Fig. 7). A key
difference between the KM-A, ML.1, and ML.2 δ18O profiles,
however, is the absence of a sharp decrease in the δ18O values
at ∼3.87 ka in our records, which marks the termination of the
4.2 ka event in the KM-A record. One possible explanation for this apparent
difference is the large uncertainties of the KM-A record. Another
plausible source of the difference may stem from the dissolution of speleothem
calcite in the KM-A sample between 0 and 29 mm depths (corresponding to
∼3.65 and 5.08 ka; Fig. 7), which may have either altered the age
of the top date of the KM-A (i.e., making it younger than its true age) or
affected the δ18O values of calcite during this period.
However, without a comprehensive, petrographic examination of the KM-A
sample, we are unable to assess the aforementioned reasons for such
differences.
The inferred pattern of ISM variability during the 4.2 ka
event: the ML.1 δ18O record is shown here as z score (left
y axis) and anomalies (right y axis). The horizontal dashed lines indicate
1 standard deviation and the vertical color-saturated shaded bars denote
periods of inferred drier (yellow) pluvial (green) and variable conditions.
The vertical red bars delineate the periods of multi-decadal droughts
(z score >1). The horizonal dashed double arrows mark the commonly
accepted duration of the 4.2 ka event (see text) and the horizontal shaded
bars indicate broad hydroclimate patterns inferred from other regional proxy
records as mentioned in the text (see also Figs. 1 and 2). The circles with
2σ error bars show a subset of 230Th dates (see Tables 1
and S1 for a complete listing of 230Th dates).
The ISM variability and possible climate forcing
The z-score-transformed ML.1 δ18O profile (Fig. 8) illustrates
the ISM variability between ∼3.8 and 4.6 ka. The z score is
calculated by using the equation of the form z=(x-μ)/σ, where
x represents the individual ML.1 δ18O value, and μ and
σ are the mean and the standard deviation of the entire ML.1
δ18O record. The interval marking the onset of the 4.2 ka event
in our record (∼4.255 ka) is indicated by a transition from pluvial
(inferred by the lower δ18O values) to variable ISM (dry–wet)
conditions, with the latter superimposed by a few short-term (< decade)
droughts (Fig. 8). Subsequently, the period between 4.07 and ∼4.01 ka is marked by persistently lower δ18O values,
implying stronger ISM (Fig. 8). The latter was terminated by a rapid increase
in the δ18O values (∼1.0 ‰, Fig. 5), suggesting
an abrupt weakening of the ISM at ∼4.01 ka that occurred within a
period of ∼10 years. Notably, as discussed above, the ML.1 and ML.2
δ18O profiles show gradual increasing trends over the entire
length of the record, which was punctuated by two multi-decadal weak monsoon
events centered at ∼3.970 (∼20 years) and ∼3.915 ka (∼25 years), respectively (Fig. 8). These aspects of our ISM
reconstruction differ from previous proxy records from the ISM domain, which
typically portray the 4.2 ka event as a multi-century drought (e.g.,
Berkelhammer et al., 2012; Dixit et al., 2014). Our new data, however,
demonstrate that prominent decadal to multi-decadal variability, together with
the intermittent occurrence of multi-decadal periods of low rainfall, was the
dominant mode of ISM variability during the period coeval with the 4.2 ka
event (Figs. 5 and 8). These observations are consistent with previous
reconstructions of ISM variability from high-resolution proxy records from
the Indian subcontinent over the last 2 millennia (e.g., Sinha et al.,
2011a, 2015; Kathayat et al., 2017) as well as during the instrumental period
(e.g., Krishnamurthy and Shukla, 2000; Goswami et al., 2006). Periodic
perturbations in coupled modes of ocean–atmosphere variability, such as the
El Niño–Southern Oscillation (ENSO), and/or dynamical processes intrinsic
to the monsoon system, such as quasi-periodic episodes of intense (“active”)
and reduced (“break”) monsoon rainfall, are key processes that are known to
produce multi-decadal periods of droughts over large parts of Asia. For
instance, Sinha et al. (2011b) suggest that ISM circulation can “lock” into
decadal to multi-decadal periods of a “break-dominated” mode of ISM
circulation that promotes enhanced convection over the eastern equatorial
Indian Ocean, which in turn suppresses convection and rainfall over the
continental monsoon regions. Additionally, the source of multi-decadal
droughts may stem from the switch-on of the modern ENSO regime around the
4.2 ka event, which would presumably also weaken the ISM (e.g., Donders et
al., 2008; Conroy et al., 2008).
In conclusion, our new record from the Mawmluh Cave in Meghalaya, India,
provides a high-resolution history of ISM during a period contemporaneous
with the 4.2 ka event. While our record shares broad similarities with a
previous lower-resolution (∼6 years) reconstruction of ISM from the
same cave (Berkelhammer et al., 2012), key differences between the two
records are also evident, which are likely due to the more refined age
controls (approximately nine 230Th dates spanning the 4.2 ka event interval
and the higher (annual) temporal resolution of our record). Our reconstruction
suggests that the ISM exhibited prominent decadal to multi-centennial
variability, including sporadic but prominent multi-decadal periods of reduced
ISM rainfall (droughts), during the period spanning the 4.2 ka event. These
aspects of our reconstruction are qualitatively similar to ISM variability
during the late and middle Holocene as inferred from the previous
speleothem-based reconstructions of ISM from the Indian subcontinent (e.g.,
Kathayat et al., 2017).
All data needed to evaluate the conclusions in the paper
are presented in the paper. Additional data related to this paper may be
requested from the authors. The data will be archived at the NOAA National
Climate Data Center
(https://www.ncdc.noaa.gov/data-access/paleoclimatology-data, last
access: 26 November 2018).
Correspondence and requests for materials should be addressed to
Gayatri Kathayat (kathayatgayatrintl@gmail.com, kathayat@xjtu.edu.cn) and Hai Cheng
(cheng021@xjtu.edu.cn).
GK and HC designed the research and experiments. GK wrote the first draft
of the paper. HC, AS, and MB revised the paper. GK, HC, and XL did
the fieldwork and collected the samples. GK, HC, HZ, and RLE conducted the
230Th dating. GK, PD, and HL conducted the oxygen isotope
measurements. All authors discussed the results and provided inputs on the
paper.
The authors declare that they have no conflict of
interest.
This article is part of the special issue “The 4.2 ka BP
climatic event”. It is a result of “The 4.2 ka BP Event: An International
Workshop”, Pisa, Italy, 10–12 January 2018.
Acknowledgements
We thank Giovanni Zanchetta and the two anonymous reviewers for their comments that
helped improve the original paper. We thank Digambar Singh Chauhan,
Chetan Singh Chauhan, Aditya Singh Kathayat, Geetanjali Kathayat, Neha Pant, Sanjay Melkani, Clive Dunnai,
Ardy Dunnai, and Ceejey Dunnai for their assistance during the
fieldwork. This work is supported by grants from the Natural Science
Foundation of China to Gayatri Kathayat (NSFC 41703007), Hai Cheng (NSFC
41731174 and 4157020432), R. Lawrence Edwards and Hai Cheng (NSF
1702816), and Haiwei Zhang (NSFC 41502166).
Edited by: Giovanni Zanchetta Reviewed by: two anonymous
referees
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