www.elsevier.com/locate/gloplacha

Global and Planetary Chang

Past 100 ky surface salinity-gradient response in the Eastern

Arabian Sea to the summer monsoon variation recorded

by y18O of G. sacculifer

Anjali R. Chodankara, Virupaxa K. Banakara,*, Tadamichi Obab

aNational Institute of Oceanography, Geological Oceanography Division, Donalpaula, Goa-403 004, IndiabEnvironmental Earth Science Division, Hokkaido University, Sapporo, 060-810, Japan

Received 21 January 2004; accepted 29 October 2004

Abstract

Northward flowing coastal currents along the western margin of India during winter–spring advect low-salinity Bay of

Bengal water in to the Eastern Arabian Sea producing a distinct low-salinity tongue, the strength of which is largely

governed by the freshwater flux to the bay during summer monsoons. Utilizing the sedimentary records of y18OG.

sacculifer, we reconstructed the past salinity-gradient within that low-salinity tongue, which serves as a proxy for the

variation in freshwater flux to the Bay of Bengal and hence summer monsoon intensity.

The north–south contrast in the sea level corrected (residual)-y18OG. sacculifer can be interpreted as a measure of surface

salinity-contrast between those two locations because the modern sea surface temperature and its past variation in the study

region is nearly uniform. The core-top residual-y18OG. sacculifer contrast of 0.45x between the two cores is assumed to reflect

the modern surface salinity difference of 1 psu and serves as a calibration for past variations.

The residual-y18OG. sacculifer contrast varies between ~0.2x at ~75 ky B.P. (i.e., late-Marine Isotope Stage 5) and ~0.7x at

~20 ky B.P. (i.e., Last Glacial Maximum), suggesting that the overall salinity difference between the northern- and southern-end

of the low-salinity tongue has varied between ~0.6 and ~1.6 psu. Considerably reduced difference during the former period than

the modern suggests substantially intensified and northward-extended low-salinity tongue due to intense summer monsoons

than today. On the other hand, larger difference (~1.6 psu) during the latter period indicates that the low-salinity tongue was

significantly weakened or withdrawn due to weaker summer monsoons. Thus, the salinity-gradient in the eastern Arabian Sea

low-salinity tongue can be used to understand the past variations in the Indian summer monsoons.

D 2004 Elsevier B.V. All rights reserved.

Keywords: Northern Indian Ocean; Eastern Arabian Sea; planktonic foraminifera; stable isotopes; palaeomonsoons; surface circulation

0921-8181/$ - s

doi:10.1016/j.gl

* Correspondi

E-mail addr

e 47 (2005) 135–142

ee front matter D 2004 Elsevier B.V. All rights reserved.

oplacha.2004.10.008

ng author. Tel.: +91 832 2450361; fax: +91 832 2450602.

ess: banakar@darya.nio.org (V.K. Banakar).

A.R. Chodankar et al. / Global and Planetary Change 47 (2005) 135–142136

1. Introduction

The northern Indian Ocean is characterised by

distinct surface salinity (henceforth dsalinityT) differ-ence between its two major basins viz., Arabian Sea

(AS) and Bay of Bengal (BOB; Wyrtki, 1973). The

net loss (gain) of freshwater from the western basin

(to the eastern basin) gives rise to high salinity in the

AS and low salinity in the BOB. The formation of

Arabian Sea high-salinity water (ASHSW) in the

northern AS (Prasannakumar and Prasad, 1999;

Shenoi et al., 1993) and the presence of low-salinity

BOB water in the southern AS result in northward-

increasing east–west-trending isohalines in the basin

(Fig. 1). This general east–west trend of the isoha-

lines sharply deviates towards north in the eastern-AS

(EAS) due to the presence of low-salinity tongue

along the western margin of India (Fig. 1; see Levitus

and Boyer, 1994; annual mean salinity from

www.nodc.noaa.gov). The northward-flowing pole-

ward coastal current (PCC) along the western margin

(Shetye et al., 1991), which is a part of the winter

monsoon circulation (WMC), transports on an aver-

age 6 Sv of low-salinity BOB water in to the EAS,

beginning with the decay of summer monsoons

Fig. 1. Locations of the GC-8 (northern) and CR-4 (southern) sediment c

MD900 is the sediment core (MD900963: 5 804V N, 73 853V E) studied by

sediment core (128 30V N, 718 15V E) having AMS 14C chronology (Sar

contours with corresponding values in psu. Dotted curved arrows are s

monsoon circulations. AS—Arabian Sea, BOB—Bay of Bengal, EC—Eq

South equatorial currents, SC—Somali currents.

through the winter–spring (Shankar et al., 2002).

Thus, the PCC is responsible for the existence of the

low-salinity tongue in the EAS. The presence of low-

salinity tongue and its dependence on the summer

monsoon freshwater flux to the BOB is largely

unambiguous (Shankar et al., 2002; Shetye et al.,

1991). The summer monsoon rainfall in the Hima-

layan region that drains in to the BOB through one of

the world’s largest fluvial systems viz., Ganges–

Brahmaputra, together discharging nearly 1000 km3/y

freshwater (Rao, 1979) and excess overhead precip-

itation determines the freshness of the BOB. It has

been shown that the alongshore pressure gradient

dominates the winter-wind stress in the EAS region

because of the strong density gradient developed due

to the presence of high-salinity water in the north and

low-salinity water in the south, which largely drives

the PCC (Shetye et al., 1991). However, the strength

of PCC depends on several other parameters such as

details in the decay of summer monsoon and

evolution of winter monsoons, seasonal wind patterns

and sea level changes along the coasts of India, and

the wind gap between Sri Lanka and India (D.

Shankar, personal communication, 2004). Therefore,

the strength and northward extension of the low-

ores used in the present study (open stars). Open circle-labelled as

Rostek et al. (1993) for palaeo-salinity and -labelled as 3104G is the

kar et al., 2000). Continuous lines are the annual average salinity

ummer monsoon circulations, and solid curved arrows are winter

uatorial currents, SECC—South-equatorial counter currents, SEC—

A.R. Chodankar et al. / Global and Planetary Change 47 (2005) 135–142 137

salinity tongue need not covary with the PCC but

certainly depends upon the freshness of BOB-water

entering in to the EAS and influence of the ASHSW

on the northern-EAS. Further, in addition to overhead

precipitation, the freshwater discharge in to the

northern-AS by Indus River (~210 km3/y: Rao,

1979) and by several medium seasonal rivers in to

the EAS (~100 km3/y) during the summer monsoon

determine the evaporation minus precipitation bal-

ance (E�P), in turn intensity of the ASHSW. Thus,

the interplay between the southerly low-salinity

tongue and the northerly ASHSW appears to be

responsible for the salinity-gradient in the EAS

(Shenoi et al., 1993). The changes in salinity-gradient

between two well-separated locations within the low-

salinity tongue may therefore provide a potential tool

for understanding the past variation in the intensity of

Indian summer monsoons.

In response to past fluctuations in the summer-

and winter-monsoon intensity, the salinities in both

the basins have oscillated (Duplessy, 1982; Kudrass

et al., 2001; Rostek et al., 1993), and hence, the

salinity-gradient in EAS must have fluctuated accord-

ingly. If we could track the changes in the past

salinity-gradient between northern- and southern-

regions of the low-salinity tongue, then it should be

possible to understand past changes in the strength of

low-salinity tongue and ASHSW, in turn the summer

monsoon intensity. The oxygen isotopes of the

planktonic foraminifera from sediment cores provide

an opportunity to monitor the past changes in the

salinity (Duplessy et al., 1981; Rostek et al., 1993)

because both of them covary in response to E�P in

the region (Delaygue et al., 2001). Few previous

studies from the EAS have utilized the calcite-y18Ofor reconstructing the past variations in E�P and the

summer-monsoons (Cayre and Bard, 1999; Duplessy,

1982; Rostek et al., 1997; Sarkar et al., 2000;

Thamban et al., 2001). But past variation in summer

monsoon-dependant low-salinity tongue, which has

considerable contribution to the freshening of the AS,

has not been looked into. In this study, utilizing the

calcite-y18O of the upper mixed layer dwelling

foraminifera Globigerinodes (G.) sacculifer (Hem-

bleben et al., 1989) from sediment cores collected

from the region of the low-salinity tongue in EAS, we

reconstruct the past variations in the summer monsoon

intensity.

2. Material and methods

Two sediment cores from differing salinity (by ~1

psu) regions of the low-salinity tongue in the EAS

are used for the present study. One core (SK-117/GC-

08: 15829V N and 71801V E, water depth=2500 m) is

from the northern high-salinity region (hereafter

dnorthern coreT), and the other core (SK-129/CR-04:

6830V N and 75859V E, water depth=2000 m) is from

the southern low-salinity region (hereafter dsoutherncoreT) of the modern low-salinity tongue (Fig. 1). The

salinity at the northern location is ~35.8 psu and at

the southern location is ~34.8 psu. Around 30 clean

and intact tests of G. sacculifer (without terminal-

sacs) from 250–350 Am size-range were hand picked

under microscope from 4-cm interval subsections of

the cores. The crushed and cleaned tests were reacted

with 100% phosphoric acid at 60F1 8C in a reaction

vessel to release the CO2 gas, which was then

subjected to a three-step purification in a glass-line

maintained under vacuum. The 18O/16O of the

purified CO2 was measured in a Finnigan MAT-251

mass-spectrometer against PDB calibrated NBS-20

standard. The measured isotopic ratios are expressed

as y18O in per mil (x) unit with reference to PDB.

The precision obtained by six-replicate measurement

was within F0.01x. The chronology for both the

cores was established by tuning the down-core

y18OG. sacculifer profiles (Fig. 2) to the tropical stack

of the SPECMAP marine isotope stages (MIS;

Bassinot et al., 1994), which is a modified SPEC-

MAP of Imbrie et al. (1984). Further, to evaluate the

accuracy of our age models, we compared: (a)

Holocene sedimentation rate in the northern-core

with a published record of AMS 14C-based sedimen-

tation rate for a nearby core 3104G (Fig. 1) from

nearly similar water depth (Sarkar et al., 2000). Both

the cores yield similar Holocene sedimentation rate of

~2.4 cm/ky; (b) LGM sedimentation rate in the

southern-core with nearby MD900963 core (Fig. 1)

from nearly similar water depth (Rostek et al., 1997).

Both these cores yield closely comparable sedimen-

tation rates of ~4.5 cm/ky; and (c) correspondence of

MIS5.1 and MIS5.3 interstadials between the north-

ern- and southern-cores, which superimpose fairly

well (Fig. 2). These comparisons clearly indicate that

the age-models derived for the present cores are

accurate, and hence, time-series variations in the

A.R. Chodankar et al. / Global and Planetary Change 47 (2005) 135–142138

y18OG. sacculifer of both the records exhibit good

correspondence. Nevertheless, AMS 14C-dating

would provide most accurate age models for the

studied cores particularly for the Holocene-LGM

sections, which presently not available to us.

The global ice-volume (sea level) and local-SST

and E�P together determine the seawater-y18O

composition (Emiliani, 1955; Imbrie et al., 1984;

Shackleton, 1967; 2000; Shackleton and Opdyke,

1973). As the geographical SST variation is small in

the northern Indian Ocean (within ~1 8C: www.nodc.noaa, 2002), the E�P variation leads to significant

changes in the local salinity. Therefore, the salinity

and seawater-y18O exhibit covariation in this region

(Delaygue et al., 2001). The modern SST is uniform

~28 8C in the EAS, and the last glacial cooling also

appears to be nearly uniform ~2 8C for the study

region (see Cayre and Bard, 1999; Rostek et al., 1993;

Sonzogni et al., 1998). Therefore, the SST-effect on

the surface water-y18O would be similar at both

locations. The down-core y18OG. sacculifer records of

both the cores were interpolated to 1-ky interval for

obtaining the variations on a time-scale similar to the

sea level curve (see Fig. 3; Shackleton, 2000) to

facilitate ice-volume correction to our y18OG. sacculifer.

Although this correction does not make any differ-

ence, the structure of the ice volume corrected

(residual)-y18OG. sacculifer helps appreciate the salinity

variations. A careful evaluation of the available SST

reconstructions for the EAS region (Cayre and Bard,

1999; Rostek et al., 1993; Uk37-SST data for LGM

section of the northern core: Banakar et al., in press)

suggests that, in general, the fluctuations of ~F0.5 8Coccur within each climate period. A 0.5 8C SST

variation is equivalent to ~0.1x change in the y18OG.

sacculifer (Duplessy et al., 1981). Further, the Uk37-SST

estimates contain certain uncertainty depending upon

the calibration used (see Cayre and Bard, 1999;

Sonzogni et al., 1998), and preservation of the surface

temperature signals in the sediment (Conte et al.,

2001). However, to minimise short time-scale

fluctuations, we smoothed the residual-y18OG. sacculifer

variations by 3-ky moving averages and assume that

Fig. 2. Depth vs. age relationship derived from tuning the down-

core y18OG. sacculifer of GC-8 (upper panel) and CR-4 (lower panel)

with the tropical-SPECMAP. The SPECMAP-based marine isotope

stage boundaries identified in the cores are shown with vertica

lines, while their corresponding ages are shown with horizonta

lines. Marine isotope stages are labelled with numbers. Vertica

broken line represents the depth at which the Youngest Toba Tuff

(YTT: ~72 Ka) occurs in the cores. Vertical dotted lines labelled as

5.1 and 5.3 are interstadials of the last warm period. Two downward

arrows on top of the upper panel are the AMS 14C-ages (6.7 and

15.2 ky B.P.) obtained for nearby sediment core (3104G, see Fig. 1)

by Sarkar et al. (2000).

l

l

l

Fig. 3. The relative sea level (or global ice volume) curve for the last

100 ky (bold line) and corresponding y18O correction (thin line)

required for subtracting the ice volume effect (Shackleton, 2000)

from the y18OG. sacculifer of GC-8 and CR-4 sediment cores to obtain

residual-y18O.

A.R. Chodankar et al. / Global and Planetary Change 47 (2005) 135–142 139

those smoothed variations reasonably reflect average

changes in the salinity. Therefore, we interpret the

time-series residual-y18OG. sacculifer contrast between

the two cores in terms of salinity-gradient between the

northern- and southern-region of the low-salinity

tongue through time.

Fig. 4. Time-series residual-y18OG.sacculifer records of GC-8 (bold

curve) and CR4 (thin curve). Shaded vertical bars are the glacia

stages numbered at the top. Curves are the result of 3-ky moving

averages of 1-ky step interpolated y18OG. sacculifer to minimise the

internal variability in the raw data sets. YTT—Youngest Toba Tuff

3. Results and discussion

The MIS1/2 boundary (11 ky B.P.) is identified at

25 and 40 cm; MIS 2/3 (24 ky B.P.) at 105 and 96 cm;

MIS 3/4 (57 ky B.P.) at 227 and 180 cm; and MIS 4/5

(71 ky B.P.) at 290 and 220 cm in the northern- and

southern-core, respectively (Fig. 2). The peak-LGM

(~18 ky B.P.) exhibiting most recent heaviest y18O at

75 and 70 cm; the occurrence of Youngest Toba Tuff

(YTT: ~72 ky B.P.; Ninkovich et al., 1978) identified

by the presence of abundant YTT-characteristic glass-

shards (bubble-wall-junction type: Pattan et al., 1999)

at ~298 and 228 cm, respectively, were used as the tie-

points for the age models of the cores (Fig. 2).

The time-series residual-y18OG. sacculifer in the

northern core are heavier throughout the last 100 ky

than the southern core (Fig. 4). Therefore, these two

records unambiguously suggest that the salinity in the

former core location has remained always higher than

in the latter location irrespective of monsoon varia-

tions. The residual-y18OG. sacculifer profiles of both the

cores are comparable with the palaeo-salinity record of

the southern-EAS (Rostek et al., 1993), indicating the

robustness of the present records. The full-Holocene

residual-y18OG. sacculifer in the northern core is �1.4xand in the southern core is �2x, while during the

LGM, it is �0.5x and �1.1x, respectively (Fig. 4),

i.e., Holocene-LGM contrast of ~0.9x. This contrast,

in addition to local salinity change, contains the LGM-

SST drop of ~2 8C, particularly in the vicinity of our

core locations (Sonzogni et al., 1998; Rostek et al.,

1993). Due to nonavailability of SST data, we are

unable to provide the actual salinity at the core

locations hence emphasise only relative changes. A 2

8C SST drop would increase the G. sacculifer calcite-

y18O by 0.4x (Duplessy et al., 1981). Subtracting the

SST-effect from the residual-y18OG. sacculifer contrasts

yields a remainder of ~0.5x accountable for the

Holocene-LGM local salinity change. Considering

the y18OG. sacculifer contrast of 0.45x between the

two core tops as a measure of modern salinity differ-

ence of ~1 psu between those two locations, the

increase of 0.5x in the residual-y18OG. sacculifer during

the LGM would translate in to ~1 psu salinity increase

in the EAS.

The average residual-y18OG. sacculifer contrast

between the two cores is highest (~0.7x) for the

MIS2 (24–11 ky B.P.), moderate (~0.5x) for the

MIS4 (71–57 ky B.P.), MIS3 (57–24 ky B.P.), and

MIS1 (12 ky B.P.–Present), and low (~0.3x) for the

l

.

A.R. Chodankar et al. / Global and Planetary Change 47 (2005) 135–142140

later half of the MIS5 (100–72 ky B.P.; Fig. 5). The

residual-y18OG. sacculifer contrast of 0.3x to 0.7xwould translate in to ~0.6 to 1.6 psu salinity-contrast

variation (i.e., ~F50% of the modern gradient) on

glacial–interglacial time-scale. However, the time-

series residual-y18OG. sacculifer contrast curve contains

short-term higher amplitude oscillations (0.2x to

0.9x; Fig. 5), suggesting rapid and larger changes

in the salinity-contrast (~0.4–2 psu), indicating

salinity-gradient change from ~60% less to almost

double than the modern gradient between northern

and southern locations of the low-salinity tongue.

The fluctuations in the salinity-gradient between

the northern- and southern-region of the low-salinity

tongue reflect the relative changes in the salinity of

those regions. The salinity of the northern location

depends upon the intensity of ASHSWand strength of

the Deccan Rivers discharging in to the EAS.

Whereas, the salinity in the southern location depends

mainly on the freshness of the BOB because of greatly

reduced influence of ASHSW in the upper mixed

Fig. 5. Time-series variation in the residual-y18OG. sacculifer contrast

between the northern- and southern-core (referred as salinity-

contrast in the text for clarity). The contrast was obtained for 1-

ky step interpolated residual-y18OG. sacculifer, and then the resultant

curve is smoothed by 3-ky moving average. Horizontal broken line

is the modern salinity-contrast of 1 psu assumed to be equivalent to

0.45x of residual-y18OG. sacculifer contrast between the two core-

tops. This line facilitates ready assessment of the changes in palaeo-

salinity-gradient. Grey shaded vertical bars are the glacial stages

numbered at the top. (A) Sector representing lowered salinity-

contrast than the modern contrast. Contrast tending towards d0Tindicates diminishing salinity-gradient; (B) Sector with increased

salinity-gradient than the modern. E�P=evaporation minus precip-

itation: higher E�P indicates reduced freshwater flux to the region

and vice versa.

layer (Prasannakumar and Prasad, 1999) and the river

input. The intense summer monsoons associating with

increased freshwater flux to the northern-AS, EAS,

and BOB not only intensify the freshening of the BOB

but also weaken the ASHSW due to reduced E�P.

The cumulative effect of the above should result in

strengthened low-salinity tongue and its further north-

ward extension in the EAS leading to a reduced

salinity-gradient. On the other hand, the ASHSW

intensifies with increased E�P due to increased winter

evaporation and weakened summer precipitation

(Shenoi et al., 1993). The latter also results in reduced

freshwater flux to the BOB and EAS by decreased

freshwater discharge from the Himalayan and Deccan

rivers respectively, in addition to reduced overhead

precipitation. In such circumstance, the northern core

location comes under increased influence of ASHSW

and reduced influence of low-salinity tongue. The net

effect is to increase the salinity-gradient between

northern and southern locations. This conceptual

scenario essentially suggests that the intensified

summer monsoons reduce the salinity-gradient within

the low-salinity tongue, resulting in decreased resid-

ual-y18OG. sacculifer contrast between the present two

core locations.

The high north–south salinity-contrast (up to ~1.6

psu) in the EAS during the LGM (Fig. 5) therefore

suggests weakened low-salinity tongue; that is, the

northward deviation of the isohalines was signifi-

cantly decreased during that time. This is clearly

reflected in the previously reported LGM-salinity

structure of the AS (see Fig. 3 in Duplessy, 1982).

The diminished low-salinity tongue might have been

the result of nearly collapsed communication between

the EAS and BOB due to significant reduction in the

freshwater flux to the latter and intensified northeast-

erly winter monsoon winds enhancing the alongshore

component opposing the PCC in the EAS. There are

several evidences to suggest that the LGM-monsoon

system was dominated by the winter winds. For

example: (a) the mean position of the summer

monsoons shifted southward (Sirocko et al., 1991),

and summer precipitation and southerly wind velocity

were greatly reduced (Niitsuma et al., 1991; Prell and

Kutzbach, 1992); (b) the minimum surface water-y18Ocontrast between the AS and BOB (Duplessy, 1982)

suggested significant reduction in freshwater transfer

from the AS to the Indian subcontinent; (c) enhanced

A.R. Chodankar et al. / Global and Planetary Change 47 (2005) 135–142 141

abundance of high-salinity fauna in LGM sections of

the sedimentary records (Cullen, 1981); (d) increased

salinity in the southern-EAS during the LGM (Rostek

et al., 1993); (e) dominant Mediterranean type pollens

characteristic of intense arid climate in the LGM

sediment of the AS (Van Campo et al., 1982); and (f)

enhanced deep-water nutrient injection in to the

photic-zone of the EAS, leading to high productivity

(Cayre and Bard, 1999; Rostek et al., 1997; Banakar et

al., in press). Large-scale climate models have

suggested that intense winter and extended snow cover

over the Tibet and Himalayas during the spring delays

the onset and weakens the following summer monsoon

(Yang, 1996); that is, the extreme long winters of the

LGM might have forced the summer monsoons to

decline. Thus, the weakened low-salinity tongue in the

LGM–EAS is in accordance with then existed ocean-

atmosphere coupling. Further, the earlier suggested

enhanced low-salinity water transport from the BOB to

the AS during the LGM due to intensified winter

circulation (Sarkar et al., 1990) does not appear to be

feasible, considering the circulation dynamics of the

northern Indian Ocean (Shankar et al., 2002; Shetye et

al., 1991) and present record of the salinity-gradient.

There are very few events during the last 100 ky

exhibiting lower salinity-gradient than the Holocene

(Fig. 5). They are evident particularly in the MIS5, in

that, closing of the MIS5 (~75 Ka) has recorded lowest

contrast (residual-y18OG. sacculifer contrast of ~0.2xequating to ~0.4 psu salinity-contrast), suggesting

significant (~60%) reduction in the salinity-gradient.

Thus, the last warm period (MIS5) appears to have

witnessed more intense summer monsoons than the

present warm period (Holocene). Whereas, the LGM–

EAS appears to have witnessed weakest summer

monsoons of the last 100 ky as evident in significantly

increased (nearly doubled) salinity-gradient than the

modern (Fig. 5). The pre-Holocene high-amplitude (up

to four-times the Holocene amplitude) fluctuations in

the residual-y18OG. sacculifer contrast indicate consid-

erable variability in the past summer monsoons.

4. Conclusion

The time-series residual-y18OG. sacculifer contrast

between the northern- and southern-region of the low-

salinity tongue in the Eastern Arabian Sea records

significant variations in the past salinity-gradient of

the region, in turn summer monsoon-driven changes

in the strength of low-salinity tongue.

The highest residual-y18OG. sacculifer contrast indi-

cating largest salinity-gradient during the LGM

appears to be the result of greatly reduced freshness

in the coastal current system responsible for freshening

the Eastern Arabian Sea due to reduced freshening of

the Bay of Bengal, suggesting weakest summer

monsoons. On the other hand, the lowest contrast

indicating smallest salinity-gradient during the last

warm period suggests increased freshness in the

eastern Arabian Sea due to northward extension of

the low-salinity tongue and weakened formation of the

high-salinity water in the northern Arabian Sea region,

together resulting from intense summer monsoons.

Thus, it is possible to understand past variations in the

Indian monsoon system by monitoring the time-series

changes in the salinity-gradient of the eastern Arabian

Sea characteristic low-salinity tongue.

Acknowledgements

We have greatly benefited from discussions with S.

R. Shetye, D. Shankar, and suggestions of Keith

Alverson and two anonymous reviewers. VKB thanks

the CSIR for the award of Raman Fellowship and

many colleagues of the Geosphere Laboratory, Hok-

kaido University for help during isotope measure-

ments. ARC thanks NIO for providing her an

opportunity to work for a doctoral thesis, EMR, CSIR

for awarding SRF, and START-TWAS for an oppor-

tunity to present this work during the Young Scientists

Global Change Conference-2003, Italy. The ship-time

for sampling was provided by DOD (this is NIO

contribution No. 3943).

References

Banakar, V.K., Oba, T., Chodankar, A.R., Kuramoto, T., Yamamoto,

M., Minagawa, M., in press. Monsoon related changes in sea

surface productivity and water column denitrification in the

Eastern Arabian Sea during the last glacial cycle. Mar. Geol.

Bassinot, F.C., Labeyrie, L.D., Vincent, E., Quidelleur, X.,

Shackleton, N.J., Lancelot, Y., 1994. The astronomical theory

of climate and the age of the Brunhes–Matuyama magnetic

reversal. Earth Planet. Sci. Lett. 126, 91–108.

A.R. Chodankar et al. / Global and Planetary Change 47 (2005) 135–142142

Cayre, O., Bard, E., 1999. Planktonic foraminiferal and alkenone

records of the last deglaciation from the eastern Arabian Sea.

Quat. Res. 52, 337–342.

Conte, M.H., Weber, J.C., King, L.L., Wakeham, S.G., 2001. The

alkenone temperature signal in Western North Atlantic surface

waters. Geochim. Cosmochim. Acta 65, 4275–4287.

Cullen, J.L., 1981. Microfossil evidence for changing salinity

pattern in the Bay of Bengal over the last 20,000 years.

Palaeogeogr. Palaeoclimatol. Palaeoecol. 35, 315–356.

Delaygue, G., Bard, E., Rollion, C., Jouzel, J., Stievenard, M.,

Duplessy, J.C., 2001. Oxygen isotope/salinity relationship in the

northern Indian Ocean. J. Geophys. Res. 106, 4565–4574.

Duplessy, J.C., 1982. Glacial to interglacial contrast in the northern

Indian Ocean. Nature 295, 494–498.

Duplessy, J.C., Be, A.W.H., Blank, P.L., 1981. Oxygen and carbon

isotopic composition and biogeographic distribution of plank-

tonic foraminiferain the Indian Ocean. Palaeogeogr. Palae-

oclimatol. Palaeoecol. 33, 9–46.

Emiliani, C., 1955. Pleistocene temperatures. J. Geol. 63, 538–578.

Hembleben, C., Spindler, M., Anderson, R.O., 1989. Modern

Planktonic Foraminifera. Springer-Verlag, NY, pp. 363.

Imbrie, J., Hays, J.D., Martinson, D.G., McIntyre, A.C., Mix, A.C.,

Morley, J.J., Pisias, N.G., Prell, W.L., Shackleton, N.J., 1984.

The orbital theory of Pleistocene climate: support from revised

chronology of the marine y18O record. In: Berger, A., Imbrie, J.,

Hays, J., Kukla, G., Saltzman, B. (Eds.), Milankovitch and

Climate Part-1. Plenum Reidel, pp. 269-305.

Kudrass, H.R., Hofmann, A., Doose, H., Emeis, K., Erlenkeuser, H.,

2001. Modulation and amplification of climatic changes in the

northern hemisphere by the Indian summer monsoon during the

past 80 k.y. Geology 29, 63–66.

Levitus, S., Boyer, T.P., 1994. World Ocean Atlas, vol. 94. NOAA/

NESDIS/OCL, Washington, DC, pp. 1–4.

Niitsuma, N., Oba, T., Okada, M., 1991. Oxygen and carbon isotope

stratigraphy at site 729, Oman margin. Proc. ODP, Sci. Results

117, 321–341.

Ninkovich, D., Shackleton, N.J., Abdel-Monem, A.A., Obradovich,

J.D., Izett, G., 1978. K–Ar age of thelate Pleistocene eruption of

Toba, north Sumatra. Nature 276, 574–577.

Pattan, J.N., Shane, P., Banakar, V.K., 1999. New occurrence of

youngest Toba Tuff in abyssal sediments of the central Indian

basin. Mar. Geol. 155, 243–248.

Prasannakumar, S., Prasad, T.G., 1999. Formation and spreading of

Arabian Sea high-salinity water mass. J. Geophys. Res. 104,

1455–1464.

Prell, W.L., Kutzbach, J.E., 1992. Sensitivity of the Indian monsoon

to forcing parameters and implications for its evolution. Nature

360, 647–652.

Rao, K.L., 1979. India’s Water Wealth: Its Assessment, Uses, and

Projections. Orient Longman, p. 261.

Rostek, F., Ruhland, G., Bassinot, F.C., Muller, P.J., Labeyrie, L.D.,

Lancelot, Y., Bard, E., 1993. Reconstructing sea surface

temperature and salinity using y18O and alkenone records.

Nature 364, 319–321.

Rostek, F., Bard, E., Beaufort, L., Sonzogni, C., Ganssen, G., 1997.

Sea surface temperature and productivity records for the past

240 kyr in the Arabian Sea. Deep-Sea Res. II 44, 1461–1480.

Sarkar, A., Ramesh, R., Bhattacharya, S.K., Rajagopalan, G., 1990.

Oxygen isotope evidence for stronger winter monsoon current

during the last glaciation. Nature 343, 549–551.

Sarkar, A., Ramesh, R., Somayajulu, B.L.K., Agnihotri, R., Jull,

A.J.T., Burr, G.S., 2000. High resolution Holocene monsoon

record from the eastern Arabian Sea. Earth Planet. Sci. Lett.

177, 209–218.

Shackleton, N.J., 1967. Oxygen isotope analyses and Pleistocene

temperatures re-assessed. Nature 215, 15–17.

Shackleton, N.J., 2000. A 100,000 year ice-age cycle identified and

found to lag temperature, carbon-dioxide, and orbital eccen-

tricity. Science 289, 1897–1902.

Shackleton, N.J., Opdyke, N.D., 1973. Oxygen isotopes and

paleomagnetic stratigraphy of equatorial Pacific core V28-238:

oxygen isotope temperatures and ice-volume on a 105-year and

106-year scale. Quat. Res. 3, 39–55.

Shankar, D., Vinayachandran, P.N., Unnikrishnan, A.S., 2002. The

monsoon currents in the north Indian Ocean. Prog. Oceanogr.

52, 63–120.

Shenoi, S.S.C., Shetye, S.R., Gouveia, A.D., Michael, G.S., 1993.

Salinity extrema in the Arabian Sea. In: Ittekkot, V., Nair, R.R.

(Eds.), Monsoon Biogeochemistry, vol. 76. SCOPE/UNEP

Sonderband, pp. 37–50.

Shetye, S.R., Gouveia, A.D., Shenoi, S.S.C., Michael, G.S., Sundar,

D., Almeida, A.M., Santanam, K., 1991. The coastal current off

western India during the northeast monsoon. Deep-Sea Res. 38,

1517–1529.

Sirocko, F., Sarnthein, M., Lange, H., Erlenkeuser, H., 1991.

Atmospheric summer circulation and coastal upwelling in the

Arabian Sea during the Holocene and the last glaciation. Quat.

Res. 36, 72–93.

Sonzogni, C., Bard, E., Rostek, F., 1998. Tropical sea-surface

temperatures during the last glacial period: a view based on

alkenones in Indian Ocean sediments. Quat. Sci. Rev. 17,

1185–1201.

Thamban, M., Rao, V.P., Schneider, R.R., Grootes, P.M., 2001.

Glacial to Holocene fluctuations in hydrography and produc-

tivity along the southwestern continental margin of India.

Palaeogeogr. Palaeoclimatol. Palaeoecol. 165, 113–127.

Van Campo, E., Duplessy, J.C., Rossignol-Strick, M., 1982.

Climatic conditions deduced from a 100 kyr oxygen isotope-

pollen record from Arabian Sea. Nature 296, 56–59.

Wyrtki, K., 1973. Physical oceanography of the Indian Ocean. In:

Zeitzschel, B. (Ed.), The Biology of the Indian Ocean. Springer,

Berlin, pp. 18–36.

Yang, S., 1996. Snow–monsoon associations and seasonal–inter-

annual prediction. Int. J. Climatol. 16, 125–134.