Indian winter and summer monsoon strength over the 4.2 ka BP event in foraminifer isotope records from the Indus River delta in the Arabian Sea
The 4.2kaBP event in foraminifer isotope records from the Indus River delta
Indian winter and summer monsoon strength over the 4.2 ka BP event in foraminifer isotope records from the Indus River delta in the Arabian SeaThe 4.2kaBP event in foraminifer isotope records from the Indus River deltaAlena Giesche et al.
Alena Giesche1,Michael Staubwasser2,Cameron A. Petrie3,and David A. Hodell1Alena Giesche et al. Alena Giesche1,Michael Staubwasser2,Cameron A. Petrie3,and David A. Hodell1
1Godwin Laboratory for Palaeoclimate Research, Department of Earth Sciences, University of Cambridge, Cambridge, CB2 3EQ, UK
2Institute for Geology and Mineralogy, University of Cologne, Zülpicher Str. 49a, 50674 Cologne, Germany
3Department of Archaeology, University of Cambridge, Cambridge, CB2 3DZ, UK
Received: 15 Aug 2018 – Discussion started: 03 Sep 2018 – Accepted: 07 Dec 2018 – Published: 15 Jan 2019
The plains of northwest South Asia receive rainfall during both the Indian
summer (June–September) and winter (December–March) monsoon. Researchers
have long attempted to deconstruct the influence of these precipitation
regimes in paleoclimate records, in order to better understand regional
climatic drivers and their potential impact on human populations. The
mid–late Holocene transition between 5.3 and 3.3 ka is of particular
interest in this region because it spans the period of the Indus Civilization
from its early development, through its urbanization, and onto eventual
transformation into a rural society. An oxygen isotope record of the
surface-dwelling planktonic foraminifer Globigerinoides ruber from
the northeast Arabian Sea provided evidence for an abrupt decrease in
rainfall and reduction in Indus River discharge at 4.2 ka, which the authors
linked to the decline in the urban phase of the Indus Civilization
(Staubwasser et al., 2003). Given the importance of this study, we used the
same core (63KA) to measure the oxygen isotope profiles of two other
foraminifer species at decadal resolution over the interval from 5.4 to
3.0 ka and to replicate a larger size fraction of G. ruber than
measured previously. By selecting both thermocline-dwelling
(Neogloboquadrina dutertrei) and shallow-dwelling
(Globigerinoides sacculifer) species, we provide enhanced detail of
the climatic changes that occurred over this crucial time interval. We found
evidence for a period of increased surface water mixing, which we suggest was
related to a strengthened winter monsoon with a peak intensity over 200 years
from 4.5 to 4.3 ka. The time of greatest change occurred at 4.1 ka when
both the summer and winter monsoon weakened, resulting in a reduction in
rainfall in the Indus region. The earliest phase of the urban Mature Harappan
period coincided with the period of inferred stronger winter monsoon between
4.5 and 4.3 ka, whereas the end of the urbanized phase occurred some time
after the decrease in both the summer and winter monsoon strength by 4.1 ka.
Our findings provide evidence that the initial growth of large Indus urban
centers coincided with increased winter rainfall, whereas the contraction of
urbanism and change in subsistence strategies followed a reduction in
rainfall of both seasons.
The ∼4.2 ka BP event is considered to be a defining event of the
mid–late Holocene transition period (Mayewski et al., 2004) and is marked by
intense aridity in much of western Asia, which has been linked to cultural
transitions in Mesopotamia, Egypt, and the Indus Civilization (Staubwasser
and Weiss, 2006; Weiss, 2016). Recently, a climate reconstruction from
Mawmluh cave in northeastern India has been used to formally demarcate the
post-4.2 ka time as the Meghalayan Age (Walker et al., 2012, 2018). However, defining the
exact timing and extent of aridity at ∼4.2 ka remains an open question
(Finné et al., 2011; Wanner et al., 2008). In this special issue devoted
to the “4.2 ka BP event”, we provide new paleoclimate data from a marine
core in the northern Arabian Sea over this critical time interval to better
understand the changes that occurred in both winter and summer hydroclimate
over the Indian subcontinent.
The δ18O record of Globigerinoides ruber from
marine core 63KA, obtained from the Arabian Sea off the coast of Pakistan and
produced by Staubwasser et al. (2003), was among the first well-resolved
paleoclimate records to suggest a link between a decrease in Indus River
discharge around 4.2 ka and the decline in the urban phase of the Indus Civilization. Since the publication of
this record, several other terrestrial paleoclimate reconstructions from the
region (Berkelhammer et al., 2012; Dixit et al., 2014, 2018; Giosan et al.,
2012; Kathayat et al., 2017; Menzel et al., 2014; Nakamura et al., 2016;
Prasad and Enzel, 2006) and a number of marine reconstructions (Giosan et
al., 2018; Gupta et al., 2003; Ponton et al., 2012) have added to our
understanding of the complex relationship between the Indus Civilization and
climate change. New questions have also emerged about the relative importance
of winter rain from the Indian winter monsoon (IWM) system and summer rain
from the Indian summer monsoon (ISM) during the critical time period from 5.4
to 3.0 ka, which spans the pre-urban, urban, and post-urban phases of the
Indus Civilization (Giosan et al., 2018; Petrie et al., 2017; Prasad and
Enzel, 2006). This is because the winter rain zone partially overlaps with
the summer rain zone (Fig. 1) and provides a critical supply of rain and
snowfall for the Indus River basin. However, we currently understand much
less about the behavior of the IWM than the ISM.
Figure 1(a) Annual, (b) ISM (JJAS), and (c) IWM
(DJFM) mean precipitation (1981–2010) isohyets taken from the GPCC V7 global
gridded dataset ( resolution) (Schneider et
al., 2015); note the difference in scale for summer and winter precipitation
(0–2000 mm vs. 0–500 mm). Rainfall data overlain on GEBCO 2014 ocean
bathymetry dataset (Weatherall et al., 2015), and shaded region shows extent
of the Indus Civilization. Bold arrows show main wind directions, dashed
arrows show ocean surface currents. Other studies discussed in this paper are
indicated by letters: A – core 63KA (this study; Staubwasser et al., 2003);
B – core 16A (Ponton et al., 2012); C – core Indus 11C (Giosan et al.,
2018); D – Din Gad peat record (Phadtare, 2000); E – core 39KG and 56KA
(Doose-Rolinkski et al., 2001); F – Lake Van record (Wick et al., 2003;
Lemcke and Sturm, 1997); G – Didwana playa lake (Singh et al., 1990); H –
Sambhar playa lake (Sinha et al., 2006); I – Karsandi playa lake (Dixit et
al., 2018); J – Jeita cave speleothem (Cheng et al., 2015); K – Kotla Dahar
lake (Dixit et al., 2014); L – Lonar lake (Menzel et al., 2014); M –
Mawmluh cave speleothem (Berkelhammer et al., 2012); N – Kanod playa lake
(Deotare et al., 2004); O – Bap Malar playa lake (Deotare et al., 2004); Q
– Qunf cave speleothem (Fleitmann et al., 2003); R – Rara lake (Nakamura et
al., 2016); S – Sahiya cave speleothem (Kathayat et al., 2017); T –
Foraminifer trap EAST (Curry et al., 1992); U – Lunkaransar playa lake
(Enzel et al., 1999); V – core 723A, RC27-14, RC27-23, RC27-28 (Gupta et
al., 2003; Overpeck et al., 1996); W – Soreq cave speleothem (Bar-Matthews
et al., 2003; Bar-Matthews and Ayalon, 2011); X – core M5-422 (Cullen et
al., 2000).
At its height, the Indus Civilization spanned a considerable geographical
area with a greater extent than the other ancient civilizations of its time
(Agrawal, 2007; Possehl, 2003). Today, the region that was once occupied by
Indus populations is marked by a heterogeneous rainfall pattern, and some
locations in the central Thar desert receive as little as 100 mm yr−1,
which is only about 10 % of the amount of direct annual rainfall compared
to New Delhi. Scarce direct precipitation in the central regions around the
Thar Desert is supplemented in some cases by fluvial or groundwater sources.
In addition, the distribution of winter rain (increasing towards the
northwest) is distinct from summer rain (increasing towards the east), making
regions variably suitable for growing certain crops and grazing (Petrie et
al., 2017; Petrie and Bates, 2017). While many paleoclimate studies from
South Asia (references A–C, I, K–M, S, and U in Fig. 1) have theorized
about the overall climatic impact of drought (and in most cases identified
summer monsoon as the cause), it is important to identify changes in the
relative contributions and timing of seasonal rainfall from both the winter
and summer monsoons. Previously, it has not been possible to reliably
differentiate winter from summer rain in reconstructions from the Indus
region.
In this study, we reexamined the same marine core (63KA) used in the original
research of Staubwasser et al. (2003). We first assessed the reproducibility
of the Globigerinoides ruberδ18O record using a
larger size fraction of the same species for the time period 5.4–3.0 ka. We
also measured the δ18O of two additional foraminifer
species, G. sacculifer (Globigerinoides sacculifer) and
N. dutertrei (Neogloboquadrina dutertrei), which live
deeper than G. ruber in the water column. The different ecologies of
the three species provide additional information with which to evaluate the
multiple δ18O records and assess seasonal changes in the
paleoceanography of the northeastern Arabian Sea near the mouth of the Indus
River.
The δ18O of foraminifera has been widely applied as an
indicator of temperature and salinity changes (among others, Duplessy et al.,
1992; Maslin et al., 1995; Wang et al., 1995; Rohling, 2000). Measuring the
δ18O of species calcifying at different depths can provide
further information about upper-ocean seasonal hydrography such as surface
water mixing, depth of the thermocline, and upwelling (Ravelo and Shackleton,
1995). Similar methods have been applied by several other studies (Billups et
al., 1999; Cannariato and Ravelo, 1997; Norris, 1998; Steinke et al., 2010;
Steph et al., 2009; among others), including a reconstruction of East Asian
winter monsoon strength in the South China Sea (Tian et al., 2005). Here we
apply a comparable method to samples from core 63KA in the northeastern
Arabian Sea because surface waters at this location are influenced by
freshwater discharge from the Indus River and direct precipitation during the
summer monsoon months, whereas enhanced upper-ocean mixing occurs during the
winter monsoon. We hypothesized that our new measurements of
δ18O of G. sacculifer and N. dutertrei
would allow us to track changes in upper-ocean mixing. Weaker IWM winds are
expected to result in a shorter duration and/or less intense upper-ocean
mixing, although how this signal is ultimately related to the amount or
distribution of winter rainfall in the Indus River catchment has not been
demonstrated conclusively. Dimri (2006) studied Western Disturbances for the
time period 1958–1997 and noted that years of surplus winter precipitation
are linked to significant heat loss over the northern Arabian Sea, which is
mainly attributed to intensified westerly moisture flow and enhanced
evaporation. Such conditions would promote deeper winter mixing and provide a
basis for relating thermocline depth with IWM intensity. By comparing the
δ18O of multiple species of foraminifera, we seek to infer
variations in the relative strengths of the summer and winter monsoons, and
by comparing the 63KA record to other nearby marine and terrestrial records,
we evaluate the potential role that climate played in the cultural
transformation of the Indus Civilization.
Today, most of the annual precipitation over northwest South Asia stems from
the ISM, and occurs mainly between June and September. The pressure gradient
between the low-pressure Tibetan Plateau and high-pressure Indian Ocean is
accompanied by the ITCZ (Intertropical Convergence Zone) reaching its
northward maximum in summer, which draws in moisture over the subcontinent
via southwesterly winds from the Indian Ocean (Gadgil, 2003). The summer
rainfall gradient increases from the central Thar Desert (as little as
100 mm direct summer rainfall per year) to the Himalaya mountains in the
north (>1000 mm) and the Aravalli range to the west (>500 mm)
(Fig. 1b).
The IWM rain falls between December through March and is mainly the result of
atmospheric Western Disturbances (Dimri and Dash, 2012; Yadav et al., 2012)
originating over the Mediterranean and Black Sea (Hatwar et al., 2005) that
allow for moisture incursion from the Arabian Sea (Rangachary and
Bandyopadhyay, 1987). During the IWM, the pressure gradient is reversed from
the summer condition, allowing the passage of Western Disturbances when the
ITCZ moves southward. As winter transitions to spring, predominantly
northeasterly winds shift to westerly winds (Sirocko, 1991) that result in
peak winter rainfall over the plains of northwest India in February and
March. Anomalously cool, evaporative conditions over the northern Arabian Sea
(promoting deeper winter mixing) also correlate with increased winter
precipitation in the western Himalayas (Dimri, 2006). The winter rainfall
gradient increases from the southern Thar Desert (<10 mm yr−1) up
to the Himalayas in the northwest (>400 mm) (Fig. 1c). Overall, the IWM
contributes between roughly 10 % and 50 % of the total annual
rainfall of northwest South Asia today.
The Indus and the other rivers that make up Punjab are partly fed by winter
snowmelt and ice melt from their upper mountain catchment areas. Melting
peaks during the summer months around July–August (Yu et al., 2013), which
coincides with the peak of ISM rainfall, and Indus River discharge reaches
its maximum during August (Karim and Veizer, 2002). The proportion of winter
to summer precipitation contributing to the Indus River is not entirely
clear, although one study has estimated a 64 %–72 % contribution of
winter precipitation from the deuterium excess of Indus River water (Karim
and Veizer, 2002), whereas a previous study estimated a lower
15 %–44 % contribution of snowmelt to Indus tributaries (Ramasastri,
1999). Since the 1960s, the Indus River has seen more than a 50 %
reduction in discharge because of the construction of barrages as well as the
diversion of water for agricultural uses (Ahmad et al., 2001).
2.2 Hydrography – core site and ocean-based processes
Core 63KA was obtained by the PAKOMIN cruise in 1993 (von Rad et al., 1995).
The laminated core from the northeastern Arabian Sea (24∘37′ N,
65∘59′ E) was taken at 316 m water depth on the continental
shelf, ∼100 km west of the Indus River delta. The core has high
sedimentation rates (equivalent to a temporal resolution of around
18 yr cm−1 in the period of interest, 5.4–3.0 ka), and all
foraminifer proxies were produced from the same laminated core with no
bioturbation. An important aspect of core 63KA is that different components
of the monsoon system are co-registered in the same sediment core, thereby
permitting an explicit evaluation of the relative timing of different parts
of the climate system (e.g., ISM and IWM).
Modern hydrographic conditions in the northeastern Arabian Sea are highly
influenced by the seasonal monsoon. During summertime, highest sea surface
temperatures (SSTs) are observed along with a shallow mixed layer depth <25 m (Schulz et al., 2002) (Fig. 2a). A low-salinity plume surrounds the
Indus River delta and shoreline extending as far as the coring location
(Fig. S1 in the Supplement). The reverse occurs in winter when the lowest
SSTs are accompanied by surface water mixing to >125 m, resulting in
warming of the deeper waters (Schulz et al., 2002). Northeasterly winds
promote convection in the northeastern Arabian Sea by cooling and evaporation
of surface water (Banse, 1984; Madhupratap et al., 1996), and during the
transition from winter to spring, wind directions shift from northeasterly to
westerly (Sirocko, 1991).
Figure 2(a) Seasonal surface water mixing depth based on station
EPT-2 located nearby the coring site of 63KA (adapted from Schulz et al.,
2002, who also used data from Hastenrath and Lamb, 1979).
(b) Foraminifer depth ranges based on CTD profile.
(c) Foraminifer abundances from EAST traps (overlapping peaks
indicate data from multiple traps): G. ruber (orange), G. sacculifer (green), and N. dutertrei (blue) (adapted from Curry et
al., 1992, using Zaric, 2005). (d) World Ocean Atlas (WOA) mean
(1955–2012) temperature (red) and salinity (yellow) profiles at
24.875∘ N, 65.875∘ E, shown for summer (JAS) and winter
(JFM) seasons (Locarnini et al., 2013; Zweng et al., 2013).
The northern Arabian Sea is dominated by highly saline (up to 37 psu)
surface waters known as Arabian Sea High-Salinity Water (ASHSW), which extend
from the surface down to 100 m depth (Joseph and Freeland, 2005). The high
salinity is explained by the high evaporative rates over this region. ASHSW
forms in the winter but is prevented from reaching our coring site on the
shelf by northerly subsurface currents until the summer (Kumar and Prasad,
1999). Along coastal areas, the ASHSW is starkly contrasted by the fresh
water discharge of the Indus River, combined with direct precipitation. In
contrast, surface waters in the Bay of Bengal on the eastern side of India
have much lower surface water salinity because of overall higher
precipitation and stronger stratification from weaker winds (Shenoi et al.,
2002). The heightened evaporative conditions and highly saline surface waters
of the northeastern Arabian Sea make it a sensitive study location to observe
changes in discharge of the entire Indus River catchment area – ultimately
tracking changes in monsoon strength. Unlike individual terrestrial records,
which may be affected by local climatic processes, the marine record from
core 63KA is more likely to integrate regional changes in the large-scale
ocean–atmosphere system.
Planktonic foraminifera complete their life cycle within a few weeks (Bé
and Hutson, 1977). Peak abundances indicate the time of year when each
species tends to calcify, thereby recording the δ18O and
temperature of the seawater in their CaCO3 shells primarily during
certain seasons. Foraminifer abundances in the eastern Arabian Sea have been
studied by Curry et al. (1992) using sediment traps deployed at shallow
(∼1400 m) and deep (∼2800 m) water depths (“T” in Fig. 1a).
G. ruber and G. sacculifer have peak abundances during the
summer months (June–September), whereas N. dutertrei lives mainly
during the winter and has a secondary peak in summer months (Fig. 2c).
Preferred depth ranges for each species reflect their ecological niches,
including requirements for nutrients and tolerance for ranges of temperature
and salinity (Bé and Hutson, 1977; Hemleben et al., 2012). G. ruber lives in the upper surface waters (0–10 m), G. sacculifer
is found in slightly deeper surface waters (10–40 m), and N. dutertrei inhabits the base of the mixed layer near the thermocline
(40–140 m) (estimates based on ranges from Farmer et al., 2007, and the
local CTD – conductivity, temperature, and depth – profiles) (Fig. 2b).
The AMS (accelerator mass spectrometer) radiocarbon dates from
Staubwasser et al. (2002, 2003) were obtained from 80 samples of mainly the
foraminifer G. sacculifer and three samples of O. universa.
In the interval of interest (5.4–3.0 ka), there are 15 radiocarbon dates
with a 95 % confidence range of 30–130 years. The average sample
resolution is 18 yr cm−1. Bayesian age modeling software, BACON v2.3.3
(Blaauw and Christen, 2011), was used as an R package to update the age model
of core 63KA. No major difference exists between the old and new age models,
except for the period 13–11 ka (Fig. S5, Table S2). IntCal13 was used for
radiocarbon calibration (Reimer et al., 2013) with marine reservoir ages
provided by Staubwasser et al. (2002, 2003).
3.2 Stable isotope analysis
Oxygen and carbon isotopes were measured on three species of foraminifera
selected from washed samples at 1 cm intervals throughout 132 cm of the
core covering 5.4–3.0 ka: G. ruber (white, sensu stricto), G. sacculifer, and N. dutertrei. For G. ruber, 12±8 foraminifera were picked from the 400–500 µm
size fraction with an average weight of 21.4±2.5µg. The
400–500 µm size fraction was picked because too few specimens
remained in the size fraction 315–400 µm used by Staubwasser et
al. (2003). For G. sacculifer, 34±7 foraminifera were picked
from the 315–400 µm size fraction with an average weight of 21.9±2.6µg. For N. dutertrei, 34±4 foraminifera
were picked from the 315–400 µm size fraction with an average
weight of 25.9±2.2µg. At some depth levels in the core there
were insufficient foraminifera for measurement, along with outlier
measurements in two cases, leaving 14 gaps in the G. ruber
400–500 µm record, 4 gaps in the G. sacculifer record,
and no gaps for N. dutertrei. The published G. ruber is
from the 315–400 µm size fraction and contains 17 gaps in the
depth range examined (Staubwasser et al., 2003).
All foraminifera were weighed, crushed, and dried at 50 ∘C. Samples
were cleaned for 30 min with 3 % H2O2, followed by a few
drops of acetone, ultrasonication, and drying overnight. Where sample weights
exceeded 80 µg, oxygen and carbon isotopes were measured using a
Micromass Multicarb Sample Preparation System attached to a VG SIRA mass
spectrometer. In cases of smaller sample sizes, the Thermo Scientific Kiel
device attached to a Thermo Scientific MAT253 mass spectrometer was used in
dual-inlet mode. This method adds 100 % H3PO4 to the
CaCO3, water is removed cryogenically, and the dry CO2 is
analyzed isotopically by comparison with a laboratory reference gas. For both
measurement methods, 10 reference carbonates and 2 control samples were
included with every 30 samples. Results are reported relative to
VPDB (Vienna Pee Dee Belemnite), and long-term reproducibility of laboratory standards (e.g., Carrara marble)
is better than ±0.08 ‰ for δ18O and ±0.06 ‰ for δ13C. Reproducibility of foraminiferal
measurements was estimated by five triplicate (three separately picked)
measurements of G. ruber (400–500 µm) that yielded 1
standard deviation of ±0.12 ‰ (δ18O) and ±0.10 ‰ (δ13C). For G. sacculifer
(315–400 µm) the standard deviation of eight triplicate
measurements was ±0.07 ‰ (δ18O) and ±0.07 ‰ (δ13C), and for N. dutertrei
(315–400 µm) the standard deviation of nine triplicate
measurements was ±0.06 ‰ (δ18O) and ±0.07 ‰ (δ13C).
To calculate equilibrium values of
, we used the CTD profile from
station 11 (24.62∘ N, 66.07∘ E) taken in September 1993
during PAKOMIN Sonne cruise no. 90 (von Rad, 2013), which is nearly
identical to the location of core 63KA (24.62∘ N 65.98∘ E).
The was calculated from salinity
following Dahl and Oppo (2006), and
was further calculated using the
calcite–water equation of Kim and O'Neil (1997). We also used the equation
of Shackleton (1974) as a comparative method for calculating
.
3.3 Statistical treatment
Statistical tests were applied to the raw data from the δ18O
and δ13C time series, including the package SiZer (Chaudhuri
and Marron, 1999; Sonderegger et al., 2009) in R software (2016), which
calculates whether the derivative of a time series exhibits significant
changes given a range of time spans. A Pearson's correlation test (confidence
level 95 %) was done on paired samples from both size fractions of
G. ruber. We also conducted Welch's t test to determine if the
mean population of δ18O is significantly different before
and after 4.1 ka.
Figure 3Core 63KA δ18OG. ruber from two size
fractions shown in the context of the original record and also zoomed in over
5.4–3.0 ka: 400–500 µm (red) (this study), 315–400 µm
(orange) (Staubwasser et al., 2003). δ18O of G. sacculifer 315–400 µm (green), and δ18O and
δ13C of N. dutertrei 315–400 µm (blue)
are shown over the interval 5.4–3.0 ka. Data are shown with a 210-year
loess smoothing, and modern surface values ±1σ are plotted for
comparison. Mean values for all species are denoted by the dotted line, and
the pre- and post-4.1 ka mean values are indicated by an additional dotted
line for N. dutertrei. Individual AMS radiocarbon dates are denoted
by triangles near the timeline.
As in the original data of Staubwasser et al. (2003), the oxygen isotope
results show great variability, and distinguishing long-term trends in these
data benefits from smoothing for visualization purposes. After completing all
statistical tests and performing the differences on the raw data
(132 depths), a loess (locally weighted) smoothing function was applied to
the δ18O and δ13C data from 5.4 to 3.0 ka,
using a 210-year moving window as described by Staubwasser et al. (2003).
Loess smoothing uses weighted least squares, which places more importance on
the data points closest to the center of the smoothing interval. The
bandwidth of 210 years was considered a reasonable time window for capturing
the overall trends in the dataset (other time windows are shown for
comparison in Fig. S2).
Figure 4SiZer first derivative analysis (Chaudhuri and Marron, 1999;
Sonderegger et al., 2009) applied to δ18O of
(a)G. ruber 400–500 µm, (b)G. ruber 315–400 µm, (c)G. sacculifer
315–400 µm, and (d)N. dutertrei
315–400 µm. The red areas indicate statistically significant
increases in δ18O, the blue represent decreases, and the
purple no significant change. Black horizontal lines are the smoothing
bandwidths (h=50, 80, and 200 years). The distance between the white
lines denotes the change in smoothing bandwidth scaled to the x axis.
The new δ18O measurements of G. ruber
(400–500 µm) parallel the published record of G. ruber
(315–400 µm) (Staubwasser et al., 2003), but the
δ18O of the specimens from the larger size fraction is
offset by −0.23 ‰ on average (Fig. 3). The records from two size
fractions, produced in different laboratories by different investigators,
display a weak positive correlation for the raw data (R=0.25, p<0.01,
n=109, slope 0.26, intercept −1.36), and the 210-year smoothed records
reveal good agreement in the overall trends of the data. When comparing the
two G. ruber records, it is apparent that the increasing trend in
δ18O starts well before ∼4.2 ka – perhaps as early
as ∼4.9 ka. This trend is also observed with the SiZer analysis, which
identifies a significant increase in δ18O anywhere from 4.9
to 4.2 ka depending on which smoothing window is selected (Fig. 4). The new
δ18O record of G. ruber (400–500 µm)
shows additional detail after the ∼4.2 ka BP event – i.e.,
specifically, a double-peak maximum occurring at 4.1 and 3.95 ka that is
related to seven discrete measurements with high δ18O
values. These maxima are offset from the average δ18O value
by +0.18 ‰ (smoothed average) or up to +0.38 ‰ when
considering the maximum individual measurement at 4.1 ka. The offsets from
the average values exceed 1 standard deviation of the entire record from 5.4
to 3.0 ka, which is 0.13 ‰. Although G. ruber shows an
event at 4.1 ka, it does not show a permanent step change: Welch's t test
comparing the means of pre- and post-4.1 ka indicates that the
+0.07 ‰ shift in mean δ18O values of G. ruber (315–400 µm) is statistically significant (t value =2.9, p<0.01, n=115), but the +0.03 ‰ shift in mean
δ18O values of G. ruber (400–500 µm) is
not significant (t value =1.5, p<0.2, n=118).
Figure 5δ18O of equilibrium calcite (a) calculated
from the CTD temperature and salinity profile at station 11 (von Rad,
2013) (b) with projected depth ranges of G. ruber
400–500 µm (red), G. ruber 315–400 µm
(orange), G. sacculifer 315–400 µm (green), and
N. dutertrei 315–400 µm (blue). We show estimated values
using both the original paleotemperature equation of Shackleton (1974) (dark
teal) and Kim and O'Neil (1997) (turquoise). Horizontal ranges show the
measured δ18O values of each species between 5.4 and
3.0 ka.
The relative differences in δ18O of the planktonic species
studied (G. ruber, G. sacculifer, and N. dutertrei) reflect the temperature and salinity of their habitat in the
water column: δ18OG. ruber<δ18OG. sacculifer<δ18ON. dutertrei (Fig. 3). G. sacculifer is offset from
G. ruber (315–400 µm) by approximately
+0.57 ‰, whereas N. dutertrei is offset by
+1.14 ‰. The larger size fraction of G. ruber
(400–500 µm) is offset from G. ruber
(315–400 µm) by −0.23 ‰. The offsets among species are
maintained throughout the entire record (Fig. 3). We also measured
δ18O values near the top of the core (approximately the last
200 years) for all three species in the 315–400 µm size fraction,
which continue to show the same offsets (Fig. S3). The δ18O
of G. ruber shows the greatest variance and N. dutertrei
shows the least (Fig. S4, Table S1).
Equilibrium calcite calculations based on the salinity and temperature
measurements from the September 1993 CTD profile of station 11 of the PAKOMIN
Cruise (von Rad, 2013) show the expected depth habitats of the three
foraminifer species (Fig. 5). G. ruber is generally found at
0–30 m, G. sacculifer at 15–40 m, and N. dutertrei at
60–150 m (Farmer et al., 2007). Using the CTD profile from our core
location, we compare these depth ranges with the measured
δ18O values. The calculated depths ranges agree well with
those expected on the basis of other studies, placing G. ruber in
the upper 10 m, G. sacculifer at 10–40 m, and N. dutertrei at 40–140 m.
G. sacculiferδ18O increases around 4.1 ka, and
Welch's t test comparing the means of pre- and post-4.1 ka indicates that
the +0.08 ‰ shift in mean δ18O values is
statistically significant (t value =3.8, p<0.01, n=128). SiZer
analysis also points to a statistically significant increase at ∼4.1–3.9 ka, when considering all smoothing time windows between 20 and
500 years (Fig. 4).
Figure 6Core 63KA Δδ18O shown with a 210-year loess
smoothing. Individual AMS radiocarbon dates are denoted by triangles near the
timeline. G. ruber 315–400 µm size fraction data come
from Staubwasser et al. (2003). The green band near the timeline showing EH,
MH, and LH refers to Early Harappan (∼5.0–4.6 ka), Mature Harappan
(∼4.6–3.9 ka), and Late Harappan (∼3.9–3.6 ka) periods,
respectively.
Likewise, the dominant change in the δ18O of N. dutertrei is a mean increase at 4.1 ka (Fig. 3). SiZer analysis also
identifies a significant decrease in δ18O occurring mainly
between 4.45 and 4.35 ka, followed by a significant increase between 4.3 and
4.1 ka (Fig. 4). Welch's t test comparing the means of pre- and
post-4.1 ka indicates that the +0.08 ‰ shift in mean
δ18O values is statistically significant (t value =6.2, p<0.01, n=132), along with the +0.07 ‰ shift in mean
δ13C (t value =3.3, p<0.01, n=132).
Differencing δ18O of foraminifera (expressed as Δδ18O) in the same sample can better emphasize signals of
interest (Fig. 6). The Δδ18O of G. ruber
400–500 µm and G. ruber 315–400 µm size
fractions shows increasing similarity between ∼4.8 and 3.9 ka during
the period of overall higher δ18O. The Δδ18O of N. dutertrei and both size fractions of
G. ruber, designated ,
reveals a period of more similar values between ∼4.5 and 3.9 ka, with
two minima at 4.3 and 4.1 ka. The Δδ18O of G. sacculifer and both size fractions of G. ruber () show a period of similar values between 4.3
and 3.9 ka, with a minimum difference at 4.1 ka. In contrast, the Δδ18O of N. dutertrei and G. sacculifer
() shows the most similarity between
4.5 and 4.2 ka with a minimum at 4.3 ka, followed by the maximum
differences between 4.2 and 3.9 ka that peaks at 4.1 ka.
The trends in the original δ18O record of G. ruber
(315–400 µm) by Staubwasser et al. (2003) are reflected by our
independent δ18O measurements of G. ruber in a
larger size fraction (400–500 µm), although an important
difference exists, suggesting a decrease in freshwater discharge as early as
4.8 ka. The larger size fraction is offset by approximately
−0.2 ‰, which is similar to the size-related fractionation of
−0.3 ‰ per +100µm for G. ruber reported by
Cayre and Bassinot (1998), and this could be attributed to size-related vital
effects. Alternatively, part of the offset might be explained by
interlaboratory calibration considering the data were produced using two
different methods and mass spectrometers.
The observed 4.1 ka maximum in δ18O of G. ruber,
living near the surface during summer months, could be attributed to either
decreased SST or increased surface water salinity (Bemis et al., 1998).
Staubwasser et al. (2003) acknowledged that a decrease in SST could cause the
increase in δ18O in the G. ruber record but argued
that this explanation is unlikely because a G. ruberδ18O record from core M5-422 in the northwestern Arabian Sea
shows opposing trends over the same time period (Cullen et al., 2000), and a
local alkenone SST proxy record shows relatively higher temperatures in the
same period (Doose-Rolinski et al., 2001). If the ∼0.2 ‰
(relative to mean) increase in δ18O of G. ruber at
4.1 ka was caused by temperature change rather than salinity, a ∼1∘C cooling of surface water would be required (Kim and O'Neil,
1997).
Following Staubwasser et al. (2003), we interpret the δ18O
variations of G. ruber to be predominantly a salinity signal.
Salinity at the core site is dependent on changes in Indus River discharge,
local runoff, and direct precipitation. Although the ISM would be the main
influence on direct precipitation and runoff at the coring location, changes
in the IWM could also influence Indus River discharge.
The thermocline-dwelling foraminifera N. dutertrei show maximum
abundances during winter and are interpreted to reflect winter mixing. During
weak IWM conditions, colder unmixed water would result in higher
δ18O values of N. dutertrei, whereas enhanced
mixing and homogenization of the water column under strong IWM conditions
would decrease δ18O. The minimum of δ18O in
N. dutertrei occurs between 4.5 and 4.3 ka, pointing to a period of
strengthened IWM. We interpret the stepped increase in δ18O
of N. dutertrei at 4.1 ka to represent a decrease in IWM
wind-driven mixing. Similarly, δ13C of N. dutertrei
increases significantly after 4.1 ka (Fig. 3), which could indicate reduced
upwelling of low δ13C intermediate water (Lynch-Stieglitz,
2006; Ravelo and Hillaire-Marcel, 2007; Sautter and Thunell, 1991); however,
the interpretation of δ13C remains uncertain because of a
poor understanding of the controls on the δ13C of planktonic
foraminifera in this region. According to the δ18O signal of
N. dutertrei, the temperature pattern in the thermocline implies
surface cooling between 4.5 and 4.3 ka and surface warming after 4.1 ka
interrupted only by a period of cooling between 3.7 and 3.3 ka, which is in
broad agreement with records of alkenone sea surface temperature estimates
from cores in the northeastern Arabian Sea (“E” in Fig. 1) (Doose-Rolinski
et al., 2001; Staubwasser, 2012).
5.2 Interpretation of foraminifer Δδ18O
By using Δδ18O between foraminifer species, we can
distinguish between additional processes affecting the surface waters and
thermocline (Ravelo and Shackleton, 1995). This technique has been used
previously to infer changes in the strength of the East Asian winter monsoon
(EAWM) in the South China Sea (Tian et al., 2005), as well as mixed layer and
thermocline depth in other studies (Billups et al., 1999; Cannariato and
Ravelo, 1997; Norris, 1998). Here we use the difference in the δ18O of G. ruber and N. dutertrei () to track changes in the surface-to-deep
gradient. This gradient can be driven by either δ18O
changes in the surface-dwelling (G. ruber) and/or the
thermocline-dwelling species (N. dutertrei). During times of a
strengthened winter monsoon, will
decrease as surface waters are homogenized and the thermocline deepens.
Similarly, will also decrease during
times of a weakened summer monsoon, as decreased Indus River discharge will
increase surface water salinity and δ18O of G. ruber will become more similar to N. dutertrei.
G. sacculifer is also a surface dweller but has a slightly deeper
depth habitat than G. ruber. We thus expect G. ruber to be
more influenced by surface salinity variations than G. sacculifer
and suggest the δ18O difference between the two species
() reflects the influence of Indus
River discharge on near-surface salinity. The smallest difference in occurs at 4.1 ka, which is interpreted as an
increase in surface water salinity (Fig. 6).
The difference in δ18O between G. sacculifer and
N. dutertrei () also
reflects surface mixing and thermocline depth, but G. sacculifer is
less affected by surface salinity changes than G. ruber. Thus, the
responses of and can be used to differentiate between surface
water salinity changes and wind-driven mixing. Accordingly, simultaneously
low and indicate a period of increased surface water
mixing and increased IWM (such as the period between 4.5 and 4.3 ka), but
times of relatively low but high
and (around 5.0 ka) indicate periods of
increased Indus discharge and strength of the ISM and IWM.
The following period of low from 4.1
to 3.9 ka is likely driven by increased salinity of surface water. This
distinction becomes clearer when examining the , where increased similarity from 4.8 to 3.9 ka
(with a sharp increase at 4.1 ka) reflects the effect of increased sea
surface salinity that reduces the δ18O difference between
G. ruber and G. sacculifer. At the same time, weakened
winter mixing increases , which
occurs from 4.2 to 3.9 ka. Importantly, the proxies also indicate that
increased IWM mixing is generally positively correlated with increased Indus
discharge and vice versa. The single time period when this does not hold true
is 4.5–4.25 ka, when increased IWM mixing is coupled with decreased Indus
discharge.
In summary, our multispecies approach using δ18O of
G. ruber, G. sacculifer, and N. dutertrei allows
us to differentiate between the strength of the IWM and freshwater discharge
of the Indus River. We suggest that ISM strength decreased gradually from at
least 4.8 ka, while the IWM strength peaked around 4.5–4.3 ka and then
weakened afterwards. It is unlikely that the abrupt increase in G. ruberδ18O at 4.1 ka and low could be caused solely by the decrease in IWM
strength, even though IWM contributes to Indus River discharge. Weakening of
the ISM must have played a substantial role in the 4.1 ka shift as well,
indicated by the period 4.5–4.25 ka, when Indus discharge reflected a weak
ISM () despite a phase of strengthened
IWM.
5.3 Comparison to marine records
Other marine records from the Arabian Sea also suggest a gradual decrease in
ISM strength from ∼5 ka (Gupta et al., 2003; Overpeck et al., 1996).
Cullen et al. (2000) observed an abrupt peak in aeolian dolomite and calcite
in marine sediments in the Gulf of Oman from 4.0 to 3.6 ka, and Ponton et
al. (2012) also showed a shift to weaker ISM after 4.0 ka in the Bay of
Bengal, based on δ13C of leaf waxes. Marine IWM
reconstructions are not particularly coherent: although Doose-Rolinski et
al. (2001) find a decrease in evaporation and weakening of the ISM between
4.6 and 3.7 ka, they argue this was accompanied by a relative increase in
IWM strength. Giosan et al. (2018) inferred enhanced winter monsoon
conditions from 4.5 to 3.0 ka based on a planktic paleo-DNA and percentage
of Globigerina falconensis record close to our coring site (“C” in
Fig. 1), which disagrees with our finding of decreased upper-ocean mixing
after 4.3 ka. We suggest that the high stratigraphic (i.e., laminated) and
chronological (i.e., 15 radiocarbon dates between 5.4 and 3.0 ka) resolution
of core 63KA paired with a multispecies foraminifer δ18O
record provides a robust history of the timing of changes in IWM and ISM
strength, but additional studies are needed to resolve some of the
discrepancies among the records.
Figure 7Comparison of the δ18O record of core 63KA with
terrestrial records from the Indian subcontinent: (a, b) this study;
(c) Berkelhammer et al. (2012); (d) Dixit et al. (2018);
(e) Dixit et al. (2014); (f) this study;
(g) Nakamura et al. (2016); (h) Kathayat et al. (2017). The
mean value for each record indicated by the horizontal dashed lines is taken
for all available data between 6.0 and 2.5 ka.
The 63KA δ18O record obtained from three foraminifer species
highlights several important ocean–atmosphere changes over the 5.4–3.0 ka
time period. First, a sharp decrease occurred in both summer and winter
precipitation at 4.1 ka, which is within a broader 300-year period of
increased aridity spanning both rainfall seasons between 4.2 and 3.9 ka. In
detail, we infer a relative decrease in Indus River discharge and weakened
ISM between 4.8 and 3.9 ka, peaking at 4.1 ka, while a 200-year-long
interval of strong IWM interrupted this period from 4.5 to 4.3 ka.
Furthermore, the stepped change in δ18O of N. dutertrei suggests an enduring change in ocean–atmosphere conditions after
4.1 ka.
A relatively abrupt ∼4.2 ka BP climate event has been observed in several terrestrial records on the Indian
subcontinent, most notably Mawmluh cave (∼4.1–3.9 ka) in northeastern
India (Berkelhammer et al., 2012) and Kotla Dahar (∼4.1 ka) in
northwestern India (Dixit et al., 2014) (Fig. 7). A less abrupt yet still
arid period is documented in a peat profile (∼4.0–3.5 ka) from north
central India (Phadtare, 2000), at Lonar Lake (∼4.6–3.9 ka) in
central India (Menzel et al., 2014), and at Rara Lake (∼4.2–3.7 ka)
in western Nepal (Nakamura et al., 2016). Finally, a recent study of oxygen
and hydrogen isotopes in gypsum hydration water from Karsandi on the northern
margin of the Thar Desert showed wet conditions between 5.1 and 4.4 ka,
after which the playa lake dried out sometime between 4.4 and 3.2 ka (Dixit
et al., 2018). Considering terrestrial records can record more local climatic
conditions than marine records, it is remarkable that the records
collectively agree on a period of regional aridity between 4.2 and 3.9 ka
within the uncertainties of the age models that vary considerably among
records.
However, not all records support this finding. For example, a reconstruction
from Sahiya Cave in northwestern India shows an abrupt decrease in δ18O interpreted to reflect an increase in monsoon strength from
∼4.3–4.15 ka, followed by an arid trend after 4.15 ka (Kathayat et
al., 2017). In addition, several other Thar Desert records do not identify a
“4.2 ka BP event” sensu stricto but instead suggest that lakes
dried out several centuries earlier (Deotare et al., 2004; Enzel et al.,
1999; Singh et al., 1990) or later (Sinha et al., 2006) than 4.2 ka. This
discrepancy may relate to nonlinear climate responses of lakes, which would
not record a drought at 4.2 ka if they had already dried out earlier from
the ongoing decrease in summer rainfall. In addition, there are also
significant concerns about chronological uncertainties from the use of
radiocarbon of bulk sediment for dating in some of these records. It is also
possible that variations in the timing of climate change inferred from the
terrestrial records may be real, reflecting a different sensitivity to ISM
and IWM rain. As a marine record, core 63KA integrates large-scale
ocean–atmosphere changes and therefore can help inform the interpretation of
the more locally sensitive terrestrial records.
Figure 8Comparison of the δ18O record of core 63KA with
more distant records: (a, b) this study; (c) Bar-Matthews
et al. (2003); (d) Cheng et al. (2015); (e) Fleitmann et
al. (2003). The mean value for each record indicated by the horizontal dashed
lines is taken for all available data between 6.0 and 2.5 ka.
More distantly, several terrestrial records in the Middle East also show a
decrease in winter precipitation proxies around 4.2 ka: Jeita cave in
Lebanon records a relatively dry period between 4.4 and 3.9 ka (Cheng et
al., 2015) and Soreq cave in Israel shows a period of increased aridity
starting at ∼4.3 ka (Bar-Matthews et al., 2003; Bar-Matthews and
Ayalon, 2011) (Fig. 8). Lake Van in eastern Turkey also records reduced
spring rainfall and enhanced aridity after ∼4.0 ka (Wick et al., 2003;
Lemcke and Sturm, 1997). All of these records suggest a relatively arid
period with reduced winter precipitation after ∼4.3 ka, as inferred
from core 63KA. Qunf cave in Oman (Fleitmann et al., 2003), which is outside
the range of IWM influence, instead shows a steady mid-Holocene weakening of
the ISM that closely follows trends in summer solar insolation.
5.5 Cultural impacts
On the basis of our reconstruction of reduced IWM mixing after 4.3 ka,
accompanied by decreased freshwater discharge of the Indus River, it is worth
considering what impacts could be expected from a reduction in IWM and ISM
precipitation. A weakened IWM overlying a reduced or more variable ISM would
likely result in a distinct climate signal over the Indus River catchment,
with broad implications for seasonal river flow and water availability
throughout the year. The presence of the two rainfall systems creates a
complex and diverse range of environments and ecologies across northwest
South Asia (Petrie et al., 2017). In a situation when rainfall in both
seasons is reduced over extended periods, step shifts in the natural
environment may occur that are difficult to reverse (e.g., desertification,
lake desiccation, regional vegetation changes, decline in overbank flooding,
and shift in river avulsion patterns).
Societies reliant on IWM, ISM, or a combination of the two would have been
vulnerable to years with monsoon failure, and a shift affecting both seasons
will have challenged resilience and tested sustainability (Green et al.,
2018; Petrie et al., 2017). Archaeological research into the transition from
the urban Mature Harappan phase (∼4.6–3.9 ka) to the post-urban Late
Harappan phase (∼3.9–3.6 ka) notes progressive de-urbanization
through the abandonment of large Indus cities and a depopulation of the most
western Indus regions, concurrent with a general trend towards an increase in
concentrations of rural settlements in some areas of the eastern Indus extent
(Green and Petrie, 2018; Petrie et al., 2017; Possehl, 1997) (Fig. S6). The
relatively limited range of well-resolved available archaeo-botanical data
suggests that there was a degree of diversity in crop choice and farming
strategies in different parts of the Indus Civilization across this time span
(Petrie et al., 2016; Petrie and Bates, 2017; Weber, 1999; Weber et al.,
2010). Farmers in southerly regions appear to have focused on summer or
winter crops, while the more northern regions of Pakistan Punjab and Indian
Punjab and Haryana were capable of supporting combinations of winter and
summer crops (Petrie and Bates, 2017). Although there is evidence for diverse
cropping practices involving both summer and winter crops in the northern
areas during the urban period, agricultural strategies appeared to favor the
more intensive use of drought-resistant summer crops in the Late Harappan
period (Madella and Fuller, 2006; Petrie and Bates, 2017; Pokharia et al.,
2017; Weber, 2003; Wright, 2010). It has previously been suggested that a
weakened ISM was a major factor in these shifts (e.g., Giosan et al., 2012;
Madella and Fuller, 2006). Based on our reconstruction of a decreased IWM in
northwest South Asia after 4.3 ka with a step shift at 4.1 ka, we suggest
that both IWM and ISM climatic factors played a role in shaping the human
landscape. This includes the redistribution of population to smaller
settlements in eastern regions with more direct summer rain, as well as the
shift to increased summer-crop-dominated cropping strategies.
This study expanded on the δ18O record of planktonic
foraminifer in core 63KA of the northeastern Arabian Sea, originally
published by Staubwasser et al. (2003). Using δ18O of the
surface-dwelling foraminifera G. ruber, the original study inferred
an abrupt reduction in Indus River discharge at ∼4.2 ka. Our further
δ18O analysis of a larger size fraction of this species
corroborates maximum salinity at 4.1 and 3.95 ka. In addition, the
δ18O difference between the surface-dwelling G. ruber and slightly deeper-dwelling G. sacculifer () reveals that surface waters were more saline
than average for the period from 4.8 to 3.9 ka. By also measuring a
thermocline-dwelling planktonic foraminiferal species, N. dutertrei,
we infer an increase in the strength of the IWM between 4.5 and 4.3 ka,
followed by a reduction in IWM-driven mixing that reaches a minimum at
4.1 ka.
Assuming that weaker IWM mixing implies a reduction in IWM rainfall amount or
duration over northwest South Asia under past climatic conditions, the 63KA
core is used to infer important changes in seasonal hydrology of the Indus
River catchment. We propose that a combined weakening of the IWM and ISM at
4.1 ka led to what has been termed the “4.2 ka BP” drought over northwest South Asia. The intersection of both
a gradually weakening ISM since 4.8 ka and a maximum decrease in IWM
strength at 4.1 ka resulted in a spatially layered and heterogeneous drought
over a seasonal to annual timescale. Regions in the western part of the Indus
River basin accustomed to relying mainly on winter rainfall (also via river
runoff) would have been most severely affected by such changes. Regions in
the northeastern and eastern extents benefitted more from summer rainfall and
would have been less severely affected, particularly as the ISM appears to
recover strength by 3.9 ka.
Relatively strengthened IWM surface water mixing between 4.5 and 4.3 ka
correlates with a period of higher precipitation recorded at Karsandi on the
northern margin of the Thar Desert (Dixit et al., 2018), an area within the
summer rainfall zone that is also sensitive to small changes in winter
precipitation. This time span also represents the beginnings of the Mature
Harappan phase (Possehl, 2002; Wright, 2010), which implies that increasingly
urbanized settlements may have flourished under a strengthened IWM. With a
weakening of the IWM at ∼4.1 ka, eastern regions with more access to
ISM rainfall may have been more favorable locations for agriculture. This may
also help explain the broad shift in population towards more rural
settlements in the northeastern extent of the Indus Civilization that
occurred by ∼3.9 ka (Possehl, 1997; Petrie et al., 2017) and a shift
to more drought-tolerant kharif (summer) season crops in Gujarat (Pokharia et
al., 2017) and at Harappa (Madella and Fuller, 2006; Weber, 2003).
Given the importance of the relationships between humans and the environment
during the time of the Indus Civilization, understanding the impact of the
IWM on precipitation variability in northwest South Asia remains a critical
area of research. We especially need a better understanding of the wind
patterns and moisture pathways that controlled the IWM in the past.
Disentangling both the length and intensity of seasonal precipitation is a
crucial aspect of understanding the impact of climate change on past
societies, particularly in a diverse region relying on mixed water sources
(e.g., fluvial, ground aquifer, direct rainfall).
Data presented in the paper can be accessed by contacting
the corresponding author or online at
http://eprints.esc.cam.ac.uk/4371/ (Giesche et al., 2018).
MS supplied core 63KA material, AG prepared the material for isotopic
measurements, and AG and DAH interpreted the results. AG, DAH, and
CAP wrote the paper.
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.
This research was carried out as part of the TwoRains project, which
is supported by funding from the European Research Council (ERC) under the
European Union's Horizon 2020 research and innovation programme (grant
agreement no. 648609). The authors thank the following persons at the
University of Cambridge: Maryline Vautravers for foraminifera identification
and James Rolfe and John Nicolson for δ18O measurements. We
also thank our editor and reviewers for comments that improved the
paper.
Edited by: Harvey Weiss
Reviewed by: Ashish Sinha and two anonymous referees
Agrawal, D. P.: The Indus Civilization: an interdisciplinary perspective,
Aryan Books International, New Delhi, India, 2007.
Ahmad, N., Mohammad, A., and Khan, S. T.: Country Report on Water resources
of Pakistan, in: South Asia Water Balance Workshop, 30 April–2 May 2001, San
Diego, California, USA, Hansen Institute for World Peace, 2001.
Banse, K.: Overview of the hydrography and associated biological phenomena in
the Arabian Sea off Pakistan, in Marine Geology and Oceanography of the
Arabian Sea and Coastal Pakistan, edited by: Haq, B. U. and Milliman, J. D.,
Van Nostrand Reinhold, New York, USA, 273–301, 1984.
Bar-Matthews, M. and Ayalon, A.: Mid-Holocene climate variations revealed by
high-resolution speleothem records from Soreq Cave, Israel and their
correlation with cultural changes, Holocene, 21, 163–171, 2011.
Bar-Matthews, M., Ayalon, A., Gilmour, M., Matthews, A., and Hawkesworth, C.
J.: Sea–land oxygen isotopic relationships from planktonic foraminifera and
speleothems in the Eastern Mediterranean region and their implication for
paleorainfall during interglacial intervals, Geochim. Cosmochim. Ac., 67,
3181–3199, 2003.
Bé, A. W. and Hutson, W. H.: Ecology of planktonic foraminifera and
biogeographic patterns of life and fossil assemblages in the Indian Ocean,
Micropaleontology, 23, 369–414, 1977.
Bemis, B. E., Spero, H. J., Bijma, J., and Lea, D. W.: Reevaluation of the
oxygen isotopic composition of planktonic foraminifera: Experimental results
and revised paleotemperature equations, Paleoceanography, 13, 150–160, 1998.
Berkelhammer, M., Sinha, A., Stott, L., Cheng, H., Pausata, F. S., and
Yoshimura, K.: An abrupt shift in the Indian monsoon 4000 years ago, Geophys.
Monogr. Ser, 198, 75–87, 2012.
Billups, K., Ravelo, A. C., Zachos, J. C., and Norris, R. D.: Link between
oceanic heat transport, thermohaline circulation, and the Intertropical
Convergence Zone in the early Pliocene Atlantic, Geology, 27, 319–322, 1999.
Blaauw, M. and Christen, J. A.: Flexible paleoclimate age-depth models using
an autoregressive gamma process, Bayesian Anal., 6, 457–474, 2011.
Cannariato, K. G. and Ravelo, A. C.: Pliocene-Pleistocene evolution of
eastern tropical Pacific surface water circulation and thermocline depth,
Paleoceanography, 12, 805–820, https://doi.org/10.1029/97PA02514, 1997.
Cayre, O. and Bassinot, F.: Oxygen isotope composition of planktonic
foraminiferal shells over the Indian Ocean: calibration to modern
oceanographic data, Mineral. Mag., 62, 288–289, 1998.
Chaudhuri, P. and Marron, J. S.: SiZer for exploration of structures in
curves, J. Am. Stat. Assoc., 94, 807–823, 1999.
Cheng, H., Sinha, A., Verheyden, S., Nader, F. H., Li, X. L., Zhang, P. Z.,
Yin, J. J., Yi, L., Peng., Y. B., Rao, Z. G., Ning, Y. F., and Edwards, R.
L.: The climate variability in northern Levant over the past 20,000 years,
Geophys. Res. Lett., 42, 8641–8650, 2015.
Cullen, H. M., deMenocal, P. B., Hemming, S., Hemming, G., Brown, F. H.,
Guilderson, T., and Sirocko, F.: Climate change and the collapse of the
Akkadian empire: Evidence from the deep sea, Geology, 28, 379–382, 2000.
Curry, W. B., Ostermann, D. R., Guptha, M. V. S., and Ittekkot, V.:
Foraminiferal production and monsoonal upwelling in the Arabian Sea: evidence
from sediment traps, Geol. Soc. Spec. Publ., 64, 93–106, 1992.
Dahl, K. A. and Oppo, D. W.: Sea surface temperature pattern reconstructions
in the Arabian Sea, Paleoceanography, 21, PA1014, https://doi.org/10.1029/2005PA001162,
2006.
Deotare, B. C., Kajale, M. D., Rajaguru, S. N., Kusumgar, S., Jull, A. J. T.,
and Donahue, J. D.: Palaeoenvironmental history of Bap-Malar and Kanod playas
of western Rajasthan, Thar desert, J. Earth Syst. Sci., 113, 403–425, 2004.
Dimri, A. P.: Surface and upper air fields during extreme winter
precipitation over the western Himalayas, Pure Appl. Geophys., 163,
1679–1698, 2006.
Dimri, A. P. and Dash, S. K.: Wintertime climatic trends in the western
Himalayas, Climatic Change, 111, 775–800, 2012.
Dixit, Y., Hodell, D. A., and Petrie, C. A.: Abrupt weakening of the summer
monsoon in northwest India ∼4100 yr ago, Geology, 42, 339–342, 2014.
Dixit, Y., Hodell, D. A., Giesche, A., Tandon, S. K., Gázquez, F., Saini,
H. S., Skinner, L. C., Mujtaba, S. A. I., Pawar, V., Singh, R. N., and
Petrie, C. A.: Intensified summer monsoon and the urbanization of Indus
Civilization in northwest India, Sci. Rep., 8, 4225,
https://doi.org/10.1038/s41598-018-22504-5, 2018.
Doose-Rolinski, H., Rogalla, U., Scheeder, G., Lückge, A., and Rad, U.:
High-resolution temperature and evaporation changes during the late Holocene
in the northeastern Arabian Sea, Paleoceanography and Paleoclimatology, 16,
358–367, 2001.
Duplessy, J. C., Labeyrie, L., Arnold, M., Paterne, M., Duprat, J., and van
Weering, T. C.: Changes in surface salinity of the North Atlantic Ocean
during the last deglaciation, Nature, 358, 485–488, https://doi.org/10.1038/358485a0,
1992.
Enzel, Y., Ely, L. L., Mishra, S., Ramesh, R., Amit, R., Lazar, B., Rajaguru,
S. N., Baker, V. R., and Sandler, A.: High-resolution Holocene environmental
changes in the Thar Desert, northwestern India, Science, 284, 125–128, 1999.
Farmer, E. C., Kaplan, A., de Menocal, P. B., and Lynch-Stieglitz, J.:
Corroborating ecological depth preferences of planktonic foraminifera in the
tropical Atlantic with the stable oxygen isotope ratios of core top
specimens, Paleoceanography, 22, PA3205, https://doi.org/10.1029/2006PA001361, 2007.
Finné, M., Holmgren, K., Sundqvist, H. S., Weiberg, E., and Lindblom, M.:
Climate in the eastern Mediterranean, and adjacent regions, during the past
6000 years – A review, J. Archaeol. Sci., 38, 3153–3173, 2011.
Fleitmann, D., Burns, S. J., Mudelsee, M., Neff, U., Kramers, J., Mangini,
A., and Matter, A.: Holocene forcing of the Indian monsoon recorded in a
stalagmite from southern Oman, Science, 300, 1737–1739, 2003.
Gadgil, S.: The Indian monsoon and its variability, Annu. Rev. Earth Pl. Sc.,
31, 429–467, 2003.
Giesche, A., Staubwasser, M., Petrie, C. A., and Hodell, D. A.:
63KA_DataFile_CP.xlsx, available at:
http://eprints.esc.cam.ac.uk/4371/, last access: 22 December 2018.
Giosan, L., Clift, P. D., Macklin, M. G., Fuller, D. Q., Constantinescu, S.,
Durcan, J. A., Stevens, T., Duller, G. A. T., Tabrez, A. R., Gangal, K.,
Adhikari, R., Alizai, A., Filip, F., VanLaningham, S., and Syvitski, J. P.
M.: Fluvial landscapes of the Harappan civilization, P. Natl. Acad. Sci.,
109, E1688–E1694, 2012.
Giosan, L., Orsi, W. D., Coolen, M., Wuchter, C., Dunlea, A. G., Thirumalai,
K., Munoz, S. E., Clift, P. D., Donnelly, J. P., Galy, V., and Fuller, D. Q.:
Neoglacial climate anomalies and the Harappan metamorphosis, Clim. Past, 14,
1669–1686, https://doi.org/10.5194/cp-14-1669-2018, 2018.
Green, A. S. and Petrie, C. A.: Landscapes of Urbanization and
De-Urbanization: A Large-Scale Approach to Investigating the Indus
Civilization's Settlement Distributions in Northwest India, J. Field
Archaeol., 43, 284–299, https://doi.org/10.1080/00934690.2018.1464332, 2018.
Green, A. S., Bates, J., Acabado, S., Coutros, P., Glover, J., Miller, N.,
Sharratt, N., and Petrie, C. A.: How to Last a Millennium; Or a Global
Perspective on the Long-Term Dynamics of Human Sustainability, under review
for Nature Sustainability, 2018.
Gupta, A. K., Anderson, D. M., and Overpeck, J. T.: Abrupt changes in the
Asian southwest monsoon during the Holocene and their links to the North
Atlantic Ocean, Nature, 421, 354–357, https://doi.org/10.1038/nature01340, 2003.
Hastenrath, S. and Lamb, P. J.: Climatic atlas of the Indian Ocean. Part II:
The oceanic heat budget, Wisconsin University Press, Madison, Wisconsin, USA,
18, 97 pp., 1979.
Hatwar, H. R., Yadav, B. P., and Rao, Y. R.: Prediction of western
disturbances and associated weather over Western Himalayas, Curr. Sci., 88,
913–920, 2005.
Hemleben, C., Spindler, M., and Anderson, O. R.: Modern planktonic
foraminifera, Springer Science and Business Media, New York, NY, USA, 2012.
Joseph, S. and Freeland, H. J.: Salinity variability in the Arabian Sea,
Geophys. Res. Lett., 32, L09607, https://doi.org/10.1029/2005GL022972, 2005.
Karim, A. and Veizer, J.: Water balance of the Indus River Basin and moisture
source in the Karakoram and western Himalayas: Implications from hydrogen and
oxygen isotopes in river water, J. Geophys. Res.-Atmos., 107, 4362,
https://doi.org/10.1029/2000JD000253, 2002.
Kathayat, G., Cheng, H., Sinha, A., Yi, L., Li, X., Zhang, H., Li, H., Ning,
Y., and Edwards, R. L.: The Indian monsoon variability and civilization
changes in the Indian subcontinent, Sci. Adv., 3, e1701296,
https://doi.org/10.1126/sciadv.1701296, 2017.
Kim, S. T. and O'Neil, J. R.: Equilibrium and nonequilibrium oxygen isotope
effects in synthetic carbonates, Geochim. Cosmochim. Ac., 61, 3461–3475,
1997.
Kumar, S. P. and Prasad, T. G.: Formation and spreading of Arabian Sea
high-salinity water mass, J. Geophys. Res.-Oceans, 104, 1455–1464, 1999.
Lemcke, G. and Sturm, M.: δ18O and trace element
measurements as proxy for the reconstruction of climate changes at Lake Van
(Turkey): Preliminary results, in Third millennium BC climate change and Old
World collapse, Springer, Berlin, Heidelberg, Germany, 653–678, 1997.
Locarnini, R. A., Mishonov, A. V., Antonov, J. I., Boyer, T. P., Garcia, H.
E., Baranova, O. K., Zweng, M. M., Paver, C. R., Reagan, J. R., Johnson, D.
R., Hamilton, M., and Seidov, D.: World Ocean Atlas 2013, Volume 1:
Temperature, edited by: Levitus, S. and Mishonov, A., NOAA Atlas NESDIS 73,
Silver Spring, MD, USA, 40 pp., 2013.
Lynch-Stieglitz, J.: Tracers of past ocean circulation, in: Treatise on
geochemistry, 6, edited by: Elderfield, H., Holland, H. D., and Turekian, K.
K., Elsevier, Amsterdam, the Netherlands, 433–451, 2006.
Madella, M. and Fuller, D. Q.: Palaeoecology and the Harappan Civilisation of
South Asia: a reconsideration, Quaternary Sci. Rev., 25, 1283–1301, 2006.
Madhupratap, M., Kumar, S. P., Bhattathiri, P. M. A., Kumar, M. D.,
Raghukumar, S., Nair, K. K. C., and Ramaiah, N.: Mechanism of the biological
response to winter cooling in the northeastern Arabian Sea, Nature, 384,
549–552, 1996.
Maslin, M. A., Shackleton, N. J., and Pflaumann, U.: Surface water
temperature, salinity, and density changes in the northeast Atlantic during
the last 45,000 years: Heinrich events, deep water formation, and climatic
rebounds, Paleoceanography, 10, 527–544, 1995.
Mayewski, P. A., Rohling, E. E., Stager, J. C., Karlén, W., Maasch, K.
A., Meeker, L. D., Meyerson, E. A., Gasse, F., van Kreveld, S., Holmgren, K.,
Lee-Thorp, J., Rosqvist, G., Rack, F., Staubwasser, M., Schneider, R. R., and
Steig, E. J.: Holocene climate variability, Quaternary Res., 62, 243–255,
2004.
Menzel, P., Gaye, B., Mishra, P. K., Anoop, A., Basavaiah, N., Marwan, N.,
Plessen, B., Prasad, S., Riedel, N., Stebich, M., and Wiesner, M. G.: Linking
Holocene drying trends from Lonar Lake in monsoonal central India to North
Atlantic cooling events, Palaeogeogr. Palaeocl., 410, 164–178, 2014.
Nakamura, A., Yokoyama, Y., Maemoku, H., Yagi, H., Okamura, M., Matsuoka, H.,
Miyake, N., Osada, T., Adhikari, D. P., Dangol, V., Ikehara, M., Miyairi, Y.,
and Matsuzaki, H.: Weak monsoon event at 4.2 ka recorded in sediment from
Lake Rara, Himalayas, Quatern. Int., 397, 349–359, 2016.
Norris, R. D.: Planktonic foraminifer biostratigraphy: eastern equatorial
Atlantic, in: Proceedings of the Ocean Drilling Program: Scientific results,
159, 445–479, 1998.
Overpeck, J., Anderson, D., Trumbore, S., and Prell, W.: The southwest Indian
Monsoon over the last 18000 years, Clim. Dynam., 12, 213–225, 1996.
Petrie, C. A. and Bates, J.: `Multi-cropping', Intercropping and Adaptation
to Variable Environments in Indus South Asia, J. World Prehist., 30, 81–130,
2017.
Petrie, C. A., Bates, J., Higham, T., and Singh, R. N.: Feeding ancient
cities in South Asia: dating the adoption of rice, millet and tropical pulses
in the Indus civilisation, Antiquity, 90, 1489–1504, 2016.
Petrie, C. A., Singh, R. N., Bates, J., Dixit, Y., French, C. A., Hodell, D.
A., Pandey, A. K., Parikh, D., Pawar, V., Redhouse, D. I., and Singh, D. P.:
Adaptation to variable environments, resilience to climate change:
Investigating land, water and settlement in Indus Northwest India, Curr.
Anthropol., 58, 2017.
Phadtare, N. R.: Sharp decrease in summer monsoon strength
4000–3500 cal yr BP in the Central Higher Himalaya of India based on
pollen evidence from alpine peat, Quaternary Res., 53, 122–129, 2000.
Pokharia, A. K., Agnihotri, R., Sharma, S., Bajpai, S., Nath, J., Kumaran, R.
N., and Negi, B. C.: Altered cropping pattern and cultural continuation with
declined prosperity following abrupt and extreme arid event at yrs BP: Evidence from an Indus archaeological site Khirsara,
Gujarat, western India, PloS one, 12, e0185684,
https://doi.org/10.1371/journal.pone.0185684, 2017.
Ponton, C., Giosan, L., Eglinton, T. I., Fuller, D. Q., Johnson, J. E.,
Kumar, P., and Collett, T. S.: Holocene aridification of India, Geophys. Res.
Lett., 39, L03704, https://doi.org/10.1029/2011GL050722, 2012.
Possehl, G. L.: The transformation of the Indus civilization, J. World
Prehist., 11, 425–472, 1997.
Possehl, G. L.: The Indus Civilization: a Contemporary Perspective, Rowman
Altamira, Walnut Creek, CA, USA, 2002.
Possehl, G. L.: The Indus Civilization: an introduction to environment,
subsistence, and cultural history, Indus ethnobiology, Lexington, Lanham, MD,
USA, 1–20, 2003.
Prasad, S. and Enzel, Y.: Holocene paleoclimates of India, Quaternary Res.,
66, 442–453, 2006.
Ramasastri, K. S.: Snow melt modeling studies in India, in: The Himalayan
Environment, edited by: Dash, S. K. and Bahadur, J., New Age International,
New Delhi, India, 59–70, 1999.
Rangachary, N. and Bandyopadhyay, B. K.: An analysis of the synoptic weather
pattern associated with extensive avalanching in Western Himalaya, Int.
Assoc. Hydrol. Sci. Publ, 162, 311–316, 1987.
Ravelo, A. C. and Hillaire-Marcel, C.: Chapter Eighteen the use of oxygen and
carbon isotopes of foraminifera in Paleoceanography, Dev. Mar. Geol., 1,
735–764, 2007.
Ravelo, A. C. and Shackleton, N. J.: Evidence for surface-water circulation
changes at Site 851 in the eastern Tropical Pacific Ocean, in: Proceedings of
the Ocean Drilling Program, Scientific Results, College Station, TX, USA
(Ocean Drilling Program), edited by: Pisias, N. G., Mayer, L. A., Janecek, T.
R., Palmer-Julson, A., van Andel, T. H., 138, 503–514,
https://doi.org/10.2973/odp.proc.sr.138.126.1995, 1995.
Reimer, P. J., Bard, E., Bayliss, A., Beck, J. W., Blackwell, P. G., Ramsey,
C. B., Buck, C. E., Cheng, H., Edwards, R. L., Friedrich, M., Grootes, P. M.,
Guilderson, T. P., Haflidason, H., Hajdas, I., Hatté, C., Heaton, T. J.,
Hoffmann, D. L., Hogg, A. G., Hughen, K. A., Kaiser, K. F., Kromer, B.,
Manning, St.W., Niu, M., Reimer, R. W., Richards, D. A., Scott, E. M.,
Southon, J. R., Staff, R. A., Turney, C. S. M., and van der Plicht, J.:
IntCal13 and Marine13 radiocarbon age calibration curves 0–50,000 years cal
BP, Radiocarbon, 55, 1869–1887, 2013.
Rohling, E. J.: Paleosalinity: confidence limits and future applications,
Mar. Geol., 163, 1–11, 2000.
Sautter, L. R. and Thunell, R. C.: Seasonal variability in the δ18O and δ13C of planktonic foraminifera from an
upwelling environment: sediment trap results from the San Pedro Basin,
Southern California Bight, Paleoceanography, 6, 307–334, 1991.
Schneider, U., Becker, A., Finger, P., Meyer-Christoffer, A., Bruno, R., and
Ziese, M.: GPCC Full Data Reanalysis Version 7.0 at 0.5∘: Monthly
Land-Surface Precipitation from Rain-Gauges built on GTS-based and Historic
Data, Deutscher Wetterdienst/Global Precipitation Climatology Centre,
Offenbach, Germany, 2015.
Schulz, H., von Rad, U., and Ittekkot, V.: Planktic foraminifera, particle
flux and oceanic productivity off Pakistan, NE Arabian Sea: modern analogues
and application to the palaeoclimatic record, Geol. Soc. Spec. Publ., 195,
499–516, 2002.
Shackleton, N. J.: Attainment of isotopic equilibrium between ocean water and
the benthonic foraminifera genus Uvigerina: isotopic changes in the ocean
during the last glacial, Colloques Internationaux du C.N.R.S., Paris, France,
219, 203–209, 1974.
Shenoi, S. S. C., Shankar, D., and Shetye, S. R.: Differences in heat budgets
of the near-surface Arabian Sea and Bay of Bengal: Implications for the
summer monsoon, J. Geophys. Res.-Oceans, 107, 3052,
https://doi.org/10.1029/2000JC000679, 2002.
Singh, G., Wasson, R. J., and Agrawal, D. P.: Vegetational and seasonal
climatic changes since the last full glacial in the Thar Desert, northwestern
India, Rev. Palaeobot. Palyno., 64, 351–358, 1990.
Sinha, R., Smykatz-Kloss, W., Stüben, D., Harrison, S. P., Berner, Z.,
and Kramar, U.: Late Quaternary palaeoclimatic reconstruction from the
lacustrine sediments of the Sambhar playa core, Thar Desert margin, India,
Palaeogeogr. Palaeocl., 233, 252–270, 2006.
Sirocko, F.: Deep-sea sediments of the Arabian Sea: A paleoclimatic record of
the southwest-Asian summer monsoon, Geol. Rundsch., 80, 557–566, 1991.
Sonderegger, D. L., Wang, H., Clements, W. H., and Noon, B. R.: Using SiZer
to detect thresholds in ecological data, Front. Ecol. Environ., 7, 190–195,
2009.
Staubwasser, M.: Late Holocene drought pattern over West Asia, Climates,
Landscapes, and Civilizations, Geophys. Monogr. Ser., 198, 89–96, 2012.
Staubwasser, M. and Weiss, H.: Holocene climate and cultural evolution in
late prehistoric–early historic West Asia, Quaternary Res., 66, 372–387,
2006.
Staubwasser, M., Sirocko, F., Grootes, P. M., and Erlenkeuser, H.: South
Asian monsoon climate change and radiocarbon in the Arabian Sea during early
and middle Holocene, Paleoceanography and Paleoclimatology, 17, 1063,
https://doi.org/10.1029/2000PA000608, 2002.
Staubwasser, M., Sirocko, F., Grootes, P. M., and Segl, M.: Climate change at
the 4.2 ka BP termination of the Indus valley civilization and Holocene
south Asian monsoon variability, Geophys. Res. Lett., 30, 1425,
https://doi.org/10.1029/2002GL016822, 2003.
Steinke, S., Mohtadi, M., Groeneveld, J., Lin, L. C., Löwemark, L., Chen,
M. T., and Rendle-Bühring, R.: Reconstructing the southern South China
Sea upper water column structure since the Last Glacial Maximum: Implications
for the East Asian winter monsoon development, Paleoceanography and
Paleoclimatology, 25, PA2219, https://doi.org/10.1029/2009PA001850, 2010.
Steph, S., Regenberg, M., Tiedemann, R., Mulitza, S., and Nürnberg, D.:
Stable isotopes of planktonic foraminifera from tropical Atlantic/Caribbean
core-tops: Implications for reconstructing upper ocean stratification, Mar.
Micropaleontol., 71, 1–19, 2009.
Tian, J., Wang, P., Chen, R., and Cheng, X.: Quaternary upper ocean thermal
gradient variations in the South China Sea: Implications for east Asian
monsoon climate, Paleoceanography, 20, PA4007, https://doi.org/10.1029/2004PA001115,
2005.
Von Rad, U., Schulz, H., Khan, A. A., Ansari, M., Berner, U., Čepek, P.,
Cowie, G., Dietrich, P., Erlenkeuser, H., Geyh, M., Jennerjahn, T.,
Lückge, A., Marchig, V., Riech, V., Rösch, H., Schäfer, P.,
Schulte, S., Sirocko, F., and Tahir, M.: Sampling the oxygen minimum zone off
Pakistan: glacial-interglacial variations of anoxia and productivity
(preliminary results, SONNE 90 cruise), Mar. Geol., 125, 7–19, 1995.
Walker, M. J., Berkelhammer, M., Björck, S., Cwynar, L. C., Fisher, D.
A., Long, A. J., Lowe, J. J., Newnham, R. M., Rasmussen, S. O., and Weiss,
H.: Formal subdivision of the Holocene Series/Epoch: a Discussion Paper by a
Working Group of INTIMATE (Integration of ice-core, marine and terrestrial
records) and the Subcommission on Quaternary Stratigraphy (International
Commission on Stratigraphy), J. Quaternary Sci., 27, 649–659, 2012.
Walker, M. J., Head, J. H., Berkelhammer, M., Björck, S., Cheng, H.,
Cwynar, L., Fisher, D., Gkinis, V., Long, A., Lowe, J., Newnham, R.,
Rasmussen, S. O., and Weiss, H.: Formal ratification of the subdivision of
the Holocene Series/Epoch (Quaternary System/Period): two new Global Boundary
Stratotype Sections and Points (GSSPs) and three new stages/subseries,
Episodes, 41, 213–223, https://doi.org/10.18814/epiiugs/2018/018016, 2018.
Wang, L., Sarnthein, M., Duplessy, J. C., Erlenkeuser, H., Jung, S., and
Pflaumann, U.: Paleo sea surface salinities in the low-latitude Atlantic: The
δ18O record of Globigerinoides ruber (white),
Paleoceanography, 10, 749–761, 1995.
Wanner, H., Beer, J., Bütikofer, J., Crowley, T. J., Cubasch, U.,
Flückiger, J., Goosse, H., Grosjean, M., Joos, F., Kaplan, J. O.,
Küttel, M., Müller, S. A., Prentice, C., Solomina, O., Stocker, T.
F., Tarasov, P., Wagner, M., and Widmann, M.: Mid-to Late Holocene climate
change: an overview, Quaternary Sci. Rev., 27, 1791–1828, 2008.
Weatherall, P., Marks, K., Jakobsson, M., Schmitt, T., Tani, S., Arndt, J.
E., Rovere, M., Chayes, D., Ferrini, V., and Wigley, R.: A new digital
bathymetric model of the world's oceans, Earth and Space Science, 2,
331–345, 2015.
Weber, S. A.: Seeds of urbanism: palaeoethnobotany and the Indus
Civilization, Antiquity, 73, 813–826, 1999.
Weber, S. A.: Archaeobotany at Harappa: Indications for Change, Indus
Ethnobiology: New Perspectives from the Field, edited by: Weber, S. and
Belcher, B., Lexington Books, Lanham, MD, USA, 175–198, 2003.
Weber, S. A., Barela, T., and Lehman, H.: Ecological continuity: An
explanation for agricultural diversity in the Indus Civilization and beyond,
Man and Environment, 35, 62–75, 2010.
Weiss, H.: Global megadrought, societal collapse and resilience at
4.2–3.9 ka BP across the Mediterranean and West asia, Clim. Chang. Cult.
Evol., PAGES Mag., 24, 62–63, https://doi.org/10.22498/pages.24.2.62, 2016.
Wick, L., Lemcke, G., and Sturm, M.: Evidence of Lateglacial and Holocene
climatic change and human impact in eastern Anatolia: high-resolution pollen,
charcoal, isotopic and geochemical records from the laminated sediments of
Lake Van, Turkey, Holocene, 13, 665–675, 2003.
Wright, R. P.: The ancient Indus: Urbanism, economy and society, Case Studies
in Early Societies, Cambridge University Press, Cambridge, UK, 10, 416 pp.,
2010.
Yadav, R. K., Kumar, K. R., and Rajeevan, M.: Characteristic features of
winter precipitation and its variability over northwest India, J. Earth Syst.
Sci., 121, 611–623, 2012.
Yu, W., Yang, Y. C., Savitsky, A., Alford, D., Brown, C., Wescoat, J.,
Debowicz, D., and Robinson, S.: The Indus basin of Pakistan: The impacts of
climate risks on water and agriculture, The World Bank, Washington, D.C.,
USA, 2013.
Zweng, M. M., Reagan, J. R., Antonov, J. I., Locarnini, R. A., Mishonov, A.
V., Boyer, T. P., Garcia, H. E., Baranova, O. K., Johnson, D. R., Seidov, D.,
and Biddle, M. M.: World Ocean Atlas 2013, Volume 2: Salinity, edited by:
Levitus, S. and Mishonov, A., NOAA Atlas NESDIS 74, 39 pp., 2013.