Spatial and temporal coherence of paleomonsoon records ......520 MANISH TIWARI et al qualitative climate proxies omitted here, reference is made to Jagadheesha et al (1999a), Korisettar

Spatial and temporal coherence of paleomonsoonrecords from marine and land proxies in the Indian region

during the past 30 ka

MANISH TIWARI1, S MANAGAVE2, M G YADAVA2 and R RAMESH2,∗

1National Centre for Antarctic and Ocean Research, Vasco de Gama, Goa 403 804, India.2Physical Research Laboratory, Navrangpura, Ahmedabad 380 009, India.

∗e-mail: [email protected]

Recent paleomonsoon data derived from various natural archives, such as tree-rings, speleothemsand deep sea sediments, buttressed by reliable chronology, with high temporal resolution aids toassess monsoon variability, both spatial and temporal, during the past ∼30 ka. While 1000 yearlong coniferous chronologies from the Himalaya help to decipher past temperature changes (e.g.,Little Ice Age), it is the tropical teak trees from peninsular India that hold promise for sub-seasonal monsoon reconstruction, as shown by oxygen isotopic studies. Likewise oxygen isotopes ofspeleothems from central India have provided proxy rainfall data for the last 10,000 years, albeitwith some gaps that need filling. Good coherence is observed from distant speleothem records.Oxygen isotopes in foraminifers have helped to understand ocean-atmosphere interactions for thepast ∼35 ka, and allowed comparison of monsoon record based on wind proxies and run-off (rain)proxies.

1. Introduction

During the past century precipitation has beenestimated to have increased by 0.5–1% per decadeover most of the mid- and high latitudes of thenorthern hemisphere continents and 0.2–0.3% overthe tropical (10◦S to 10◦N) land areas; in partsof Asia and Africa, the frequency and intensityof droughts seem to have increased (IPCC AR42007; Houghton et al 2001) during 1900–1995 AD.Future projections using climate models pointto an increase in the monsoon rainfall in mostparts of India with increasing greenhouse gasesand sulphate aerosols (Rupa Kumar et al 2002).Many north Indian rivers such as the Ganga,Yamuna, etc., have shown a sharp decline in thesummer discharge in the recent past, possiblydue the shrinking of the Himalayan glaciers thatfeed them (Gosain and Rao 2003). These obser-vations lead us to an important question whether

these are consequences of global warming or onlya part of the low frequency climate variabilityinherent in the system. To answer this wouldrequire reliably dated, high-resolution records ofthe past climate/monsoonal precipitation for aduration longer than the spatially and temporallylimited instrumental weather records. Monsoonis known to exhibit variance over a range ofperiods such as annual, decadal and centennialto millennial timescales (Webster 1987; Gadgil2003). For studying the monsoon variability in thelatter two timescales, recourse is made to variousocean- and land-based natural archives suchas marine/lacustrine/riverine/aeolian sediments,Tibetan/Himalayan glaciers/ice, speleothems, tree-rings, corals, etc. Here we present our currentunderstanding of past variations of south Asianmonsoon based on various marine and terrestrialmonsoon proxies, which are well constrained, reli-able, chronologies. For a more detailed account of

Keywords. Monsoon; Arabian Sea; tree-rings; speleothems; marine sediments; stable isotopes.

517

518 MANISH TIWARI et al

Table

1.

Table

show

ing

abrief

acc

ount

ofden

dro

clim

ato

logic

alin

ves

tigati

ons

carr

ied

out

inth

eIn

dia

nre

gio

n.

Abbre

via

tions:

Tem

p–

Tem

per

atu

re,ppt

–pre

cipitation,E

–ea

st,W

–w

est,

C–

centr

al,

S–

south

.

Clim

ate

/E

nvironm

enta

lR

efer

ence

Spec

ies

Loca

tion

Alt

itude

Tim

esp

an

Met

hod

signal

1Pant

and

Borg

aonka

r(1

984)

Pin

us

roxb

urg

hii

WH

imala

ya

1525–1850

1805–1980

Rin

g-w

idth

Tem

pra

nge

and

min

imum

tem

p

2R

am

esh

etal(1

985)

Abi

espin

dro

wW

Him

ala

ya

2650

1903–1932

Cel

lulo

seδD

,δ1

8O

,δ1

3C

Isoto

pic

coher

ence

bet

wee

ndiff

eren

ttr

ees

and

diff

eren

tra

dia

ldir

ecti

ons

wit

hin

tree

s

3H

ughes

and

Dav

ies

(1986)

Abi

espin

dro

w,Pic

easm

ithia

na

Kash

mir

2380–3400

1620–1982

Rin

g-w

idth

,ri

ng

den

sity

Spri

ng,ea

rly

sum

mer

,w

hole

sum

mer

tem

p

4R

am

esh

etal(1

989)

Tec

tona

grandis

WIn

dia

1917–1980

δDM

onso

onalra

infa

ll

5H

ughes

(1992)

Abi

espin

dro

wW

Him

ala

ya

2620–3400

1620–1982

Rin

g-w

idth

,ri

ng

den

sity

Apri

l–M

ay,A

ug–Sep

tte

mp

and

Apri

l–Sep

tppt

6B

hatt

ach

ary

ya

etal(1

992)

Ced

rus

deo

dara

,P.ro

xburg

hii,

P.wallic

hia

na,Pic

easm

ithia

na,A

.pin

dro

,A

.sp

ecta

bilis,

J.re

curv

a,

Tsu

gadum

osa

,L.po

tanin

i

Nep

al

1320–3720

1569–1979

Rin

g-w

idth

Moistu

reco

ndit

ions,

spring

and

sum

mer

tem

p

7B

org

aonka

ret

al(1

994)

Abi

espin

dro

w,Pic

easm

ithia

na

Kash

mir

2600–2900

1612–1982

1775–1982

Rin

g-w

idth

Sum

mer

tem

pand

ppt

8B

org

aonka

ret

al(1

996)

Ced

rus

deo

dara

WH

imala

ya

>2000

1676–1988

Rin

g-w

idth

Sum

mer

tem

pand

sum

mer

ppt

9B

org

aonka

rand

Pant

(1997)

Ced

rus

deo

dara

,A

.pin

dro

wP.ro

xburg

hii,Pic

easm

ithia

na

WH

imala

ya

1590–1990

Rin

g-w

idth

Sum

mer

,pre

-monso

on

tem

pand

ppt

10

Yadav

etal(1

997)

Ced

rus

deo

dara

,P.wallic

hia

na

Pic

easm

ithia

na

WH

imala

ya

2700–3000

1698–1987

Rin

g-w

idth

April–

May

tem

p

11

Pant

etal(1

998)

Pic

easm

ithia

na

WH

imala

ya

1673–1990

Rin

g-w

idth

April–

May

tem

pand

ppt

12

Chaudhary

etal(1

999)

Abi

esden

sa,Lari

xgr

iffith

iana

EH

imala

ya

1503–1994

Rin

g-w

idth

Tem

p

13

Yadav

etal(1

999)

Ced

rus

deo

dara

EH

imala

ya

>2730

1390–1987

Rin

g-w

idth

Marc

h–M

ayte

mp

South

ern

Osc

illa

tion

14

Pant

etal(2

000)

Ced

rus

deo

dara

WH

imala

ya

1845–1986

Rin

g-w

idth

,ri

ng

den

sity

Pre

-monso

on

tem

pand

ppt

15

Yadav

and

Park

(2000)

Ced

rus

deo

dara

WH

imala

ya

>2730

1171–1988

Rin

g-w

idth

Tem

pand

ppt

16

Sin

gh

and

Yadav

(2000)

Pin

us

wallic

hia

na

WH

imala

ya

3400

1590–1999

Rin

g-w

idth

Win

ter

(Dec

–Feb

)te

mp,gla

cier

retr

eat,

warm

ing

tren

d

17

Chaudhary

and

Bhatt

ach

ary

ya

(2000)

Lari

xgr

iffith

iana

EH

imala

ya

3320

1891–1996

Rin

g-w

idth

May

tem

p

18

Borg

aonka

ret

al(2

001)

Abi

espin

dro

w,Pic

easm

ithia

na,Ced

rus

deo

dara

,P.ro

xburg

hii

WH

imala

ya

2100–3000

1849–1990

Rin

gden

sity

,R

ing-w

ith

Pre

-monso

on

(Marc

h–M

ay)

tem

pand

ppt

19

Bhatt

ach

ary

ya

etal(2

001)

Abi

espin

dro

wW

Him

ala

ya

1625–1995

Rin

g-w

idth

Tem

pand

ppt,

gla

cialm

ovem

ent

SPATIAL AND TEMPORAL COHERENCE OF PALEOMONSOON RECORDS 519Table

1.

(Conti

nued

).

Clim

ate

/E

nvironm

enta

lR

efer

ence

Spec

ies

Loca

tion

Alt

itude

Tim

esp

an

Met

hod

signal

20

Chaudhary

and

Bhatt

ach

ary

ya

(2002)

Pin

us

kesi

aShillo

ng,

Meg

hala

ya

950–1560

1859–2000

Rin

g-w

idth

Rain

fall

Dec

ember

(pre

vio

us

yea

r)and

Marc

h(c

urr

ent

yea

r)

21

Yadav

and

Sin

gh

(2002a)

Ced

rus

deo

dara

WH

imala

ya

2720–3260

1480–1987

Rin

g-w

idth

Mea

nsp

ring

(Marc

h–M

ay)

tem

p

22

Yadav

and

Sin

gh

(2002b)

Taxu

sbu

ccata

WH

imala

ya

2910

1656–2000

Rin

g-w

idth

Pre

-monso

on

(Marc

h–June)

tem

p

23

Cook

etal(2

003)

Abi

essp

ecta

bilu

s,T

suga

dum

osa

,Pin

us

wallic

hia

na,

J.re

curv

a,Pic

easm

ithia

na,

Ulm

us

wallic

hin

a

Nep

al

1830–3630

1796–1992

(com

mon

inte

rval)

Old

-es

t856–1996

Rin

g-w

idth

Tem

p

24

Bhatt

ach

ary

ya

and

Chaudhary

(2003)

Abi

esden

saE

Him

ala

ya

3300–3900

1507–1987

Rin

g-w

idth

Late

sum

mer

(July

–Sep

t)te

mp

25

Sin

gh

etal(2

004)

Ced

rus

deo

dara

WH

imala

ya

805–2002

Rin

g-w

idth

Tem

pfr

om

Feb

–M

ay

26

Yadav

etal(2

004)

Ced

rus

deo

dara

WH

imala

ya

1226–2000

Rin

g-w

idth

Pre

-monso

on

tem

pLit

tle

Ice

Age

27

Sin

gh

and

Yadav

(2005)

Ced

rus

deo

dara

WH

imala

ya

1556–1987

Rin

g-w

idth

Spring

ppt

28

Sano

etal(2

005)

Abi

essp

ecta

bilu

sW

Nep

al

3850

1717–2000

Rin

g-w

idth

,ri

ng

den

sity

Marc

h–M

ayte

mp

and

ppt

29

Bhatt

ach

ary

ya

etal(2

006)

Bet

ula

utilis

WH

imala

ya

3700

1571–2002

Rin

g-w

idth

Ppt

ofM

arc

h–June,

tem

pof

Marc

h–A

pri

lG

laci

alm

ovem

ent

30

Yadav

etal(2

006)

Junip

erus

macr

opo

da

WH

imala

ya

2600–3300

420–2003

Rin

g-w

idth

Ppt

31

Tre

ydte

etal(2

006)

Junip

erus

exce

lsa,

J.tu

rkes

tanic

aN

Pakista

n2900,>

3700

950–1990

Cel

lulo

seδ1

8O

Ppt

32

Sin

gh

and

Yadav

(2007)

Pin

us

gera

rdia

na

WH

imala

ya

919–2005

Rin

g-w

idth

Tem

pand

ppt

33

Borg

aonka

ret

al(2

007)

Him

ala

yan

conifer

sW

Him

ala

ya

2900–3450

1747

onw

ard

sR

ing-w

idth

,ri

ng

den

sity

Pre

-monso

on

sum

mer

tem

pand

ppt,

glo

balw

arm

ing

34

Borg

aonka

ret

al(2

007)

Tec

tona

grandis

CIn

dia

SIn

dia

1654–2001

(CIn

dia

)1481–2003

(SIn

dia

)

Rin

g-w

idth

Annualra

infa

ll(C

India

)

35

Shah

etal(2

007)

Tec

tona

grandis

CIn

dia

1835–1997

Rin

g-w

idth

June–

Sep

tra

infa

ll

36

Bhatt

ach

ary

ya

etal(2

007)

Tec

tona

grandis

SIn

dia

1743–1986

Early

wood

ves

selare

aO

ct–N

ovra

infa

llofpre

vio

us

yea

rand

Aprilra

infa

llof

curr

ent

yea

r

37

Ram

etal(2

008)

Tec

tona

grandis

CIn

dia

1827–2001

Rin

g-w

idth

Rain

fall

and

moistu

rein

dex

38

Managav

eet

al(2

009a,b

)Tec

tona

grandis

Cand

SIn

dia

Cel

lulo

seδ1

8O

Under

standin

gsu

b-a

nnual

variati

ons

inδ1

8O


qualitative climate proxies omitted here, referenceis made to Jagadheesha et al (1999a), Korisettarand Ramesh (2002) and Singh et al (2007).

2. Tree-ring based climatic records

Instrumental rainfall data collected by IndianMeteorological Department (IMD) dates back toAD 1813 when the first rain gauge station wasestablished at Chennai. By 1871, a fairly good net-work of rain gauge 312 stations was established. Tostudy the variations in rainfall prior to 1871, it isnecessary to use climate proxies such as tree-rings.Dendroclimatological investigations carried out atthe Birbal Sahni Institute of Palaeobotany (BSIP),Lucknow, Indian Institute of Tropical Meteorology(IITM), Pune and Physical Research Laboratory(PRL), Ahmedabad have demonstrated the poten-tial of high altitude as well as tropical trees inreconstructing past climate. Table 1 summarizesmajor dendroclimatological investigations in theIndian region.

Several aspects of the global climate changehave been revealed by dendroclimatic investiga-tions over Indian region. The prominent feature ofdendroclimatic temperature reconstruction in theHimalaya is the lack of a pronounced increase intemperature during the early 20th century, a globalwarming trend observed in most places in theNorthern Hemisphere (Mann et al 1999; Crowleyand Lowery 2000; Esper et al 2002). Further, theHimalayan region, especially Tibet and centralAsia, shows a decreasing trend in temperature inthe late 20th century. Another contrasting featureobserved in western Himalayan temperature recordas compared to the Northern Hemisphere recordis lack of evidence for the Little Ice Age (LIA),a feature it shares with the record observed atTibet and other central Asian regions. These evi-dences could point to the presence of regional andtemporal differences in the past climate. However,hydrogen isotopic measurements clearly show theevidence for LIA (Ramesh 2000). Also Nijampurkaret al (2002) found evidence for the LIA in acentral Himalayan glacier. Thus it appears thatthe detrending procedure employed for construct-ing ring-width indices removes the climatic signalto some extent. More recently, Cook et al (2003)have reported presence of Little Ice Age and warm-ing trend of the 20th century based on their investi-gations in Nepal. Yadava et al (2004) demonstratedpresence of the Little Ice Age in western Himalaya.

Although mid-altitude chronologies do notshow any long term trend, some high altitudeand periglacial tree-ring chronologies appear tooffer evidence of global warming. Chronologiesbuilt by Singh and Yadav (2000), Bhattacharyya

et al (2006) and Borgaonkar et al (2007) showincreased tree growth in recent times possibly indi-cating warming trend over the region. The stableisotopic ratios are measured using a mass spec-trometer and expressed as δ18O and δ13C where,δ18O = [((18O/16O)sample/(18O/16O)reference) − 1]×1000� (per mil) and δ13C = [((13C/12C)sample/(13C/12C)reference) − 1] × 1000�. The referenceis either SMOW or V-PDB supplied by IAEA(International Atomic Energy Agency, Vienna).Treydte et al (2006) based on oxygen isotope ratios(δ18O) of trees from northern Pakistan and Singhand Yadav (2005) using ring-width variations oftrees from western Himalaya have reported increas-ing precipitation trend over the area. Decrease inthe pre-monsoon temperature was observed in theinstrumental and tree ring reconstructed recordsduring the late 20th century, in contrast to highlatitude regions, due likely to increased anthro-pogenic influence (Yadav et al 2004).

Dendroclimatological potential of trees in centraland southern India has not been exploited fullyas only a few tropical trees offer reliable growthrings and are long lived. Teak (Tectona gran-dis) and toona (Cedrela toona) are two specieswith good potential for the reconstruction of pastclimate (Pant and Borgaonkar 1983; Ramesh et al1985) and a few chronologies have been successfullybuilt for teak. Recent studies (Borgaonkar et al2007; Shah et al 2007; Ram et al 2008) show thatvariations in ring-widths of teak can be used toreconstruct past monsoon rainfall. Bhattacharyyaet al (2007) have shown mean vessel area of theearly wood in teak from southern India is corre-lated with the north–east (NE) monsoon rainfallof the previous year and based on this relationreconstructed past NE monsoon rainfall. Basedon high resolution δ18O composition of teak treesManagave et al (2009a,b) has demonstrated thepossibility of reconstructing sub-annual monsoonrainfall. In future, more chronologies need to bedeveloped for peninsular India using isotope andring-width studies. Future efforts may concentrateon comparing the tree-ring δ18O record with thatof speleothems to understand and document spa-tial heterogeneity of rainfall.

3. Monsoon record from speleothems

Speleothem refers to minerals precipitated in acave environment. It can be used for paleoclimaticreconstruction through the analysis of the varia-tions in stable oxygen isotopic composition (δ18O)of the carbonate (Yadava et al 2004; Yadava andRamesh 2005a,b). In the last two decades therehas been growing interest among paleoclimatolo-gists to explore potential of speleothem deposits

SPATIAL AND TEMPORAL COHERENCE OF PALEOMONSOON RECORDS 521

because: (i) these deposits are well protected fromphysical damages (erosion) by rain or wind action(ii) U-Th dating techniques using mass spectro-metry were refined and applied to speleothemdating; it has become possible to date small quan-tities (∼100mg) of even young speleothems withreasonable precision.

All speleothems growing in a cave are notsuitable for paleoclimatic reconstruction. Whena cave has several openings through which out-side dry air enters and makes rapid evapora-tion of seepage water possible, the dissolved ionicspecies (mainly calcium and bicarbonate) recom-bine quickly and precipitate carbonate. There isnot enough time left for a complete ion exchangeamong different species, a condition necessary toattain isotopic equilibrium. Therefore it becomesdifficult to interpret isotopic data from suchspeleothems. Samples with ‘promise for climateinterpretation’ do occur mostly in the interiorof caves where air ventilation is poor, ensuringboth slow carbon dioxide emanation and littleevaporation of water satisfying isotopic equili-brium. The cave environment affects the stable iso-tope ratios of oxygen (δ18O) and carbon (δ13C)while CaCO3 is precipitated, and we discuss theseseparately.

The isotopic composition of seepage water incaves reflects the average isotopic composition ofthe meteoric water falling atop the cave (e.g.,Schwarcz 1986; Ivanovich and Harmon 1995).Air temperature in the cave is equal to the averageannual surface air temperature outside, because ofthe large thermal inertia of land mass (soil andbedrock).

The fractionation factor αcw for the calcite-water system, known experimentally (Friedmanand O’Neil 1977) is given below:

αcw = (18O/16O)calcite/(18O/16O)water

Δcw = (2.78 × 106/T 2) − 2.89 (1)

where, T is the temperature during precipitationof calcite, in K and Δcw = 103 ln αcw. Any shiftin temperature in a karst region affects thecave temperature, and therefore calcite δ18Oc.Higher δ18Oc values imply lower tempera-tures (dΔcw/dT = −0.21�◦C−1 at 25◦C, fromequation 1). In addition, there are severalregional factors which influence the δ18Oc values.For example, at high latitudes rain water isrelated to the surface air temperature, d(δ18O)/dT = 0.69� ◦C−1, while at tropical locationsδ18Ow is dominantly controlled by the amount ofrain and surface temperature dependence is eithervery weak or insignificant (Dansgaard 1964).

Figure 1. Locations of different caves studied so far inIndia.

The isotopic sensitivity of any speleothem to theambient climate depends upon several factors butprimarily on cave location and local meteorology.Therefore, interpretation of speleothem δ18Oc isnot straightforward and often ambiguous (Yadavaand Ramesh 2006). This calls for the monitoring ofisotopic signatures of rain and seepage water nearthe cave.

Speleothem locations studied so far fromIndia are shown in figure 1. Speleothems fromlimestone caves in Koraput (Orissa), Jagdalpur(Chhattisgarh), Chitrakoot (Uttar Pradesh),Uttar Kannada District (Karnataka) and westernHimalaya have been studied until now. Caves whichare least disturbed by human activity are the mostsuitable for palaeoclimatic studies as the samplesare in pristine condition. Following are the cavesexploited so far for monsoon reconstruction inIndia (figure 1):1. Gupteswar cave,Koraput district, Orissa.

An actively growing stalactite was collected in1996.

2. Dandak cave, Jagdalpur district, Chhat-tisgarh. This cave lies within ∼30 km fromGupteswar. An active stalagmite was collectedin 1996.

3. Akalagavi cave, Uttar Kannada District,Karnataka. (AKG) This small cave is locatedin a tropical dense forest (C3 type). A stalag-mite was collected in 1997 and it was dated bycounting annual laminations and radiocarbonmethod.


Figure 2. Time span covered by speleothems. Radio-carbon dated speleothems are shown by single crossed boxesand U-Th dated speleothems are shown by double crossedboxes.

4. Sota cave, Chitrakoot district, UttarPradesh. This small cave is located in theKarwi taluk, a tourist place; however, humansrarely visit this cave. A stalactite from a nar-row chamber in the interior part was collectedin 1997.

5. Timta cave, Pithoragarh district, westernHimalya. A stalagmite collected from this cavewas dated by U-Th mass spectrometric method(Sinha et al 2007).

Further, samples have been collected fromAndaman and Andhra, which are being processed.Large variations in the δ18Oc are observed, indicat-ing that the temperature dependence of δ18Oc isnegligible (Yadava and Ramesh 1999a,b, 2005a,b,2007; Yadava et al 2004; Yadava et al 2007a,b).It was found that due to the dominant control ofrainfall it is difficult to recover the small-magnitudepast temperature variations from δ18Oc, or evenfrom trace elemental data, for example, Mg/Ca,Sr/Ca (Yadava and Ramesh 2001).

Although the fractionation of various carbonisotopic species is mildly temperature dependent(Hendy 1971), changes in the equilibrium frac-tionation factors with temperature are compen-sated by parallel changes in the molar ratiosof carbon species in the solution (Hendy 1971;Dulinski and Rozanski 1990). Hence, use ofδ13Cc for climate reconstruction is not successful.Presently, speleothem δ18O is being used to addressvariations in the past monsoon.

For dating speleothems both radiocarbon andU-Th methods have been applied. As shown infigure 2, time duration covered by these have somegaps. Figure 3 shows a comprehensive monsoonscenario now available from Indian speleothems.

Some important observations in the reconstruc-tions are:

(a) Teak tree ring widths have been reported(Bhattacharyya and Yadav 1999) fromsouthern India. Around 1660 AD, they arebroader, coinciding with the highly depletedδ18O signal (i.e., higher rain) around 1666 ADin AKG.

(b) Two other high resolution (comparable withthat of speleothems, i.e., ∼1 yr to ∼15 yr)paleomonsoon records from southern Asia,spanning the last 3400 yrs are available: (i) ahigh resolution (∼7 yr) record has been recon-structed by von Rad et al (1999), using thethickness variations in the varved sedimentscollected from off Karachi, Pakistan (north-eastern Arabian Sea): precipitation and hence

Figure 3. Comparison of δ18O of the Indian speleothems.Akalagavi cave has higher resolution (∼1yr). Dandak cavehas two reconstructions (B & C). Before Holocene a recon-struction is available from Timta cave (A).


the river runoff are assumed to control thevarve thickness. The precipitation at the samp-ling site (von Rad et al 1999) occurs bothduring summer (Jun.–Sep.) and winter mon-soons (November–March). These alternativelydark and light colored sediment sequences forman annual couplet. The precipitation may havefluctuated (Luckge et al 2001) due to variationsin the extreme positions of the ITCZ (inter-tropical convergence zone) and hence, thevariability in varve thickness is interpreted as aproxy for past rainfall variations. (ii) Anotherhigh resolution (∼decadal) record is basedon the stable oxygen isotope variations fromstalagmites in southern Oman (Fleitmann et al2003).

(c) Gupteswar and Oman speleothem records arewell matched. The monsoon was strongeraround 3000 BP as indicated by more depletedδ18O values and also by the increased growthrate (higher sampling density). The increas-ing trend between 1200 yr BP and 400 yr BPis seen in both the records. Also during theextremely low rainfall epochs of 1700 and2000 yr BP shown by Gupteswar, the Omanstalagmites did not grow, probably due to thecomplete lack of rain. It must be noted thatOman is more like a desert relative to easternIndia and therefore growth of stalagmites thereis more sensitive to rainfall fluctuations.

(d) The decreasing trend of rainfall from 3400to 1900 yr BP is reflected both in the varveand Gupteswar record (Luckge et al 2001,have reported lowest Ti/Al ratio (terrigenousorigin) in the same core around 2000 yr BP).But the two records differ significantly duringthe last 1500 yrs. As the speleothem recordsseparated by a larger distance agree very well,it is likely that the varve thickness response tothe monsoon is nonlinear.

4. Monsoon and associated oceanographiceffects from marine proxies

During the summer and winter monsoons the sur-face oceanic circulation in the northern IndianOcean (Arabian Sea and Bay of Bengal) experi-ences changes in direction in consonance with thechanging wind patterns (Wyrtki 1973; Schott andMcCreary Jr 2001). Intense upwelling occurs alongthe Somalian and Omanian coasts with a transportof 1.5–2 Sv in the upper 50 m (Smith and Bottero1977). The typical temperature of the upwelledwater is 19◦–24◦C (Schott and McCreary Jr 2001).The reasons attributed for such intense coastalupwelling is the Ekman divergence due to the flowof strong winds parallel to the coast. The central

Arabian Sea exhibits a bowl shaped mixed layerdeepening under the effect of Findlater jet wind-stress forcing and Ekman pumping (McCreary andKundu 1989; Rao and Sivakumar 2000). The coldand dry northeast monsoon winds accompanied bythe Ekman pumping cause subduction of the highsalinity surface waters in the northern Arabian Sea(Morrison 1997; Schott and Fischer 2000).

The upwelling zones along the Somalian andOman coasts cause intense biological and geo-chemical changes in this region with sea surfacetemperature (SST) falling by ∼4◦C as nutrient richdeeper water surfaces that enhance the sea sur-face biological productivity considerably (Wyrtki1973; Nair et al 1989; Haake et al 1993). Weakupwelling also occurs along coastal southwest India(Wyrtki 1973; Shetye 1984). During the Northeastmonsoon, minor upwelling is observed in thenortheastern Arabian Sea (Wyrtki 1973). The coldand dry NE monsoon winds causes the deepening ofthe mixed layer to a depth of 100–125 m due to con-vective mixing in the northern Arabian Sea, whichleads to nutrient injection and hence high produ-ctivity during winter monsoon in this region (Banseand McClain 1986; Madhupratap et al 1996).The typical productivity values for the westernArabian Sea are 2.0, 1.0 and 0.5 g C/m2

/day forthe SW monsoon, NE monsoon and the intermon-soon periods, respectively (Codispoti 1991; Barberet al 2001). Similarly for the eastern Arabian Sea,the typical productivity values are 0.6, 0.3 and0.2 g C/m2

/day for the SW monsoon, NE mon-soon and the intermonsoon periods, respectively(Bhattathiri et al 1996). As the moisture laden SWmonsoon winds approach the Western Ghats theyare forced to ascend resulting in copious precipita-tion and runoff into the coastal Arabian Sea, reduc-ing the sea surface salinity considerably (Sarkaret al 2000). Denitrification takes place due to thevery low concentration of oxygen in the entireArabian Sea from 250 m to 1250 m water depths(Naqvi 1987; Olson et al 1993). This oxygen mini-mum zone (OMZ) is due to the high oxygen con-sumption below the thermocline for the oxidationof organic matter supplied by the high overheadsurface productivity. Furthermore the sluggish flowof the oxygen poor intermediate water (Olsonet al 1993; You and Tomczak 1993) along witha strong tropical thermocline (due to relativelyhigh SST that prevents mixing of the oxygenrich surface waters with the deeper waters) main-tains the OMZ (Spencer et al 1982; Qasim 1982).Thus OMZ and denitrification are the interplayof monsoon winds and the ensuing productivityalong with other climatically controlled factorssuch as ocean ventilation rate (Reichart et al1997, 1998, 2002a; Schulz et al 1998; Altabet et al2002).


Figure 4. Core locations discussed in the text.

Such pronounced changes in the seawater charac-teristics make the Arabian Sea ideal for decipheringthe past changes in monsoon intensity. The surfaceproductivity that manifests itself in many formssuch as organic, calcareous and siliceous produc-tivity, also affects the carbon isotopic compositionof the seawater, which is preserved in the calciticshells of various foraminifera. Similarly the SSTand sea surface salinity alter the oxygen isotopiccomposition of these shells and they get recordedin the sea sediments. The nitrogen isotopic compo-sition of sedimentary organic matter can indicatethe denitrification intensity relatable to produc-tivity variations. Thus the downcore variations ofsuch proxies could help document the past varia-tions in monsoon intensity and the related climaticchanges.

Locations of some recent marine paleoclimatestudies are shown in figure 4 and some results aresummarized in figure 5.

4.1 Western/northern Arabian Sea

The western Arabian Sea has received the maxi-mum scientific attention for deciphering the pastmonsoon fluctuations as it experiences the mostintense biogeochemical changes during the mon-soon season. The earliest and very comprehen-sive studies were carried out by Prell et al (1980),Prell (1984), Prell and Van Campo (1986) in whichthey found that SW monsoon was weaker dur-ing the glacial periods and stronger during inter-glacials. They showed that much of the 103 to105 year variability in the monsoon is linked tosolar radiative forcing and the associated feedbackeffects.

Prell and Kutzbach (1987) proposed that glacialboundary conditions such as SST, earth’s albedo,sea level, extent and elevation of large ice massesplay equally important roles in modifying mon-soon patterns. Later Clemens et al (1991), Clemens

and Prell (1991) argued that monsoon is mainlygoverned by the precession induced insolationchanges and not by the changing glacial boundaryconditions. Several authors such as Anderson andPrell (1992), Sirocko et al (1993), Overpeck et al(1996) refuted this hypothesis and maintainedthat glacial boundary conditions indeed are instru-mental in modifying monsoon intensities. Sirockoet al (1993) measured dolomite content in thecore 74 KL off the Oman margin (figure 4), anindicator of aridity (when SW monsoon windsweaken, the dolomite content of the core increases,as north-westerlies carrying dolomite from theArabian peninsula reach the core site). Basedon this high resolution, centennial scale study,they reported that monsoonal climate changedin abrupt steps and not in a gradual mannerwith increase in its intensity observed at ∼15.5 kaand a maximum at ∼8.5 ka (figure 5C), whichthey attributed to albedo changes during the peri-ods of deglaciation. Sirocko et al (1996) pro-posed that SW monsoon intensified at 11.4 ka,coinciding with the climate transition observedin polar ice cores; thus monsoon exhibits corre-lation with the high latitude climatic changes.Naidu and Malmgren (1996) analyzed the coresfrom the western Arabian Sea and concluded thatSW monsoon was relatively stronger during 22–18 ka than ∼18–13.8 ka with a major intensifica-tion at 13 ka BP and a maximum between 10 and5 ka BP, after which it declined with the weak-est phase at 3.5 ka BP. They also observed a sub-Milankovitch periodicity of 2,200 years exhibitedby the SW monsoon induced upwelling indices fromwhich they inferred that SW monsoon is influ-enced by oceanic circulation changes that con-trols the ∼2,300 year periodicity also observedin atmospheric 14C. They concluded that lowerCaCO3 in the western Arabian Sea during inter-glacials along with higher δ13C is due to highernon-carbonate productivity and higher dissolutionof CaCO3 due to enhanced Antarctic bottom waterventilation in the equatorial Indian Ocean (Naiduet al 1993; Naidu and Malmgren 1999). Altabetet al (1995) and Ganeshram et al (2000) haveshown that denitrification intensity is controlled bySW monsoon induced productivity changes, whichwas weaker during the glacials and stronger duringinterglacial periods. Reichart and coworkers havecarried out extensive work regarding the monsoonand oxygen minimum zone (OMZ) variability forthe late Quarternary (covering the past ∼225 ka)in the northern Arabian Sea. Reichart et al (1997)studied productivity and dust input records in acore from Murray Ridge (Northern Arabian Sea)and concluded that productivity in this region ismainly controlled by the SW monsoon and theintensity of OMZ is governed by the sea surface


Figure 5. Records of SW monsoon variations for the past 35 ka from the Arabian Sea (Panel A-J) & Bay of Bengal(Panel K); the last panel represents temperature record from the Greenland Ice Sheet Project (GISP2) with Arabic numerals(1–6) indicating Dansgaard/Oeschger interstadials (LGM & YD refer to the Last Glacial Maximum & Younger Dryas,respectively); Shaded areas highlight the climatically important periods.

productivity, which was lower during the weakSW monsoon during glacial periods. Reichart et al(1998, 2002a) showed that OMZ and SW mon-soon strength varied synchronously with polar icerecords. The stadials as deciphered by the polar

records are characterized by light colored, bio-turbated sediments with low Corg implying weakOMZ due to reduced productivity that in turn isbecause of weaker SW monsoon. Reichart et al(1998, 2002b) showed that during stadials the OMZ


was destroyed because of deep convective over-turning due to intensified cool and dry wintermonsoon winds that led to surface water cool-ing and enhanced the salinity due to evapora-tion. Reichart et al (2004) further showed thatjust after strong stadials at stadial-interstadialboundaries, a brief episode of hyperstratificationtakes place due to weakened winter monsoon(and hence less cooling due to evaporation) thatfacilitates the formation of strong OMZ duringinterstadials.

Schulz et al (1998) studied the monsoon inducedOMZ variability by analyzing the TOC (totalorganic carbon) content in cores from the northernArabian Sea (136 KL, figure 4). They found astrong correlation between laminated, organic-carbon-rich bands, reflecting strong monsoon-induced biological productivity, and warminterstadials (Dansgaard–Oeschger events) fromGreenland ice core records (GISP2, figure 5G).The fact that SW monsoon and high-latitudetemperature records exhibit similar variability ledthem to propose common forcing agents such asatmospheric moisture and other greenhouse gases.Similarly, Altabet et al (2002) studied denitrifi-cation intensity in cores from the Oman margin(RC7-23, figure 4), which is strongly coupled withSW monsoon wind strength. They showed thatmonsoon intensity is closely related with the polar(Greenland as well as Antarctic) ice core recordseven on short timescales with enhanced/reducedmonsoon during the warmer/cooler periods(figure 5F) and concluded that high latitude andtropical climates are most probably linked viarapid atmospheric forcing. Leuschner and Sirocko(2003) studied three cores from the northernArabian Sea analyzing the aeolian dust contentthat reflects arid continental climate and foundthat it exhibits good correlation with the GISP2and Vostok ice records with humid periods coincid-ing with the temperature maxima. Zonneveld et al(1997) obtained a core from the Somalian upwellingregion and studied the relative dominance ofdinoflagellates cysts of the (SW) monsoon-inducedupwelling and non-upwelling species. They foundthat broadly, monsoon follows the insolation forc-ing, which is nonlinear due to the effect of snowcover over the central Asia and Tibetan plateau.The other forcing factors they identified includeglacial–interglacial boundary conditions (varyingthermohaline circulation) and tropical land coverforcing (that influences albedo). Von Rad et al(1999) and Luckge et al (2001) analyzed varvesequences in cores from the northern ArabianSea raised from OMZ and studied monsoon vari-ations for the past 5000 years. They deducedthat precipitation decreased after ∼4 ka BP with

minima centered at ∼2 ka BP and ∼500 ka BPwith higher precipitation during the interveningperiods. Almogi-Labin et al (2000) studied a corefrom the Gulf of Aden, spanning last 530 ky,and found that the productivity is controlled byNortheast (NE) monsoon. They reported two longperiods of NE monsoon intensification lasting from460–430 ky and ∼60–13 ky and concluded thatglacial boundary conditions appear to control themajor portion of the NE monsoon variability. Junget al (2002) have interpreted the high frequencydecadal to centennial scale oxygen isotopic varia-tions in a core off Somalia as due to solar insolationinduced SST variations. Staubwasser et al (2002,2003) reconstructed, using δ18O of G. ruber, theIndus river discharge that depends on SW mon-soon precipitation from a core (63 ka, figure 4)off the Pakistan coast (figure 5H). They reportedreduced SW monsoon during the Younger Dryasand a major drought event at ∼4.2 ka, which theyproposed as the cause of the termination of theurban Harappan civilization. Ivanova et al (2003)estimated primary productivity from three dif-ferent regions of the Arabian Sea for the past135 ky and concluded that variation in paleopro-ductivity is most pronounced in the northern andeastern Arabian Sea, and comparatively weak inthe western Arabian Sea. Ivanochko et al (2005)studied productivity proxies in a core from theSomali margin (figure 4) and found excellentcorrelation with the Greenland ice core record(figure 5I) with higher productivity/wind strengthduring interstadials (DO events). This led them tohypothesize that global scale millennial climaticvariability is in part driven by modulations inthe tropical hydrological cycle and tropical emis-sions of the greenhouse gases such as CH4 andN2O. Anderson et al (2002), based on percentageof G. bulloides, proposed that monsoon strengthhas been increasing for the past 400 years due tonorthern hemisphere warming and will continueto do so as the greenhouse gases concentrationincreases. G. bulloides is a temperate foraminiferalspecies and occurs in tropics where water is coolerdue to upwelling. Thus their abundance, i.e.,percentage of total planktic foraminifera, is anexcellent indicator of upwelling, controlled by thewind strength. Gupta et al (2003, 2005) comparedthe abundance of G. bulloides (figure 5J) in coresfrom off-Oman (figure 4) with North Atlanticsediment records and concluded that SW mon-soon exhibits excellent correlation with the highlatitude climate on centennial timescales (weakermonsoon during colder periods). Based on thissingle proxy, they proposed that monsoon broadlyfollows the insolation curve at 65◦N with a mon-soon maximum at ∼8.5 ka and declining since


then to ∼1.5 ka. A recent, multi-proxy study byTiwari et al (2009) in a well-dated core has shownthat the SW monsoon did not decline until laterthan 5.5 ka, consistent with phase-lag analysisof longer time series data of monsoon proxies.Thus considering the overall trend during theHolocene, SW monsoon did not follow insolation,which corroborates that monsoon lagged sum-mer insolation maxima by several thousand years(Clemens et al 1991, 1996; Clemens and Prell2003) and highlights the importance of internalfeedbacks.

Recently, several attempts have been made toreconstruct the SST from this region that yieldedhigher Last Glacial Maximum (LGM) values (e.g.,Dahl and Oppo 2006) than originally reconstructedby the CLIMAP study (CLIMAP Project Members1981). Naidu and Malmgren (2005) have recon-structed annual, summer, and winter SST throughthe last 22 ky using artificial neural networks(ANNs) based on quantitative analyses of plank-tic foraminifera. They reported that annual, sum-mer, and winter SST were 2, 1.2, and 2.6◦Ccooler, respectively, during the last glacial periodthan in the Holocene. Huguet et al (2006) deter-mined SST variations over the last 23 ky using twoorganic molecular proxies, viz., alkenone unsatura-tion index and a newly proposed TEX86 derivedfrom the membrane lipids of crenarchaeota. Thealkenone SST record shows a ∼2◦C increase sincethe LGM. They observe a cold phase between 14.5and 12 ka that may correspond to the Antarcticcold reversal, which implies a Southern Hemi-sphere control on tropical SST reconstructed bythe TEX86. Saher et al (2007) have reconstructedthe integrated SST of the SW and NE monsoonseasons based on Mg/Ca ratio in G. ruber, whichshow a glacial-interglacial SST difference of ∼2◦C,matching with the earlier studies. Recently, Anandet al (2008) attempted to reconstruct seasonal SSTfor the past 35 ky from western and eastern Ara-bian Sea using Mg/Ca ratio in G. ruber and G.bulloides and reported an annual SST difference of3–4◦C between the LGM and the Holocene in bothregions. In the eastern Arabian Sea, increased sea-sonal contrast was found during LGM (∼4◦C) thanthe present (∼1◦C), indicating enhanced deep con-vective mixing (causing lowered SST) during theNE monsoon. In the western Arabian Sea, warmingoccurred, that is, upwelling decreased, during thestadial periods (Northern Hemisphere cold events).Gupta et al (2008) studied faunal and TOC datathat indicated an early Holocene characterized bya relatively well oxygenated OMZ where the influ-ence of intense monsoon-related production wasreduced by increased Circumpolar Deep Waterincursion.

4.2 Eastern Arabian Sea

Duplessy (1982) reconstructed the Holocene andLGM sea surface conditions such as salinity andtemperature based on δ18O of G. ruber fromthe cores spread over the Bay of Bengal andthe Arabian Sea. He concluded that SW mon-soon was weaker during the LGM and NE monsoonwas stronger than present with more precipitationsouth of 10◦N. Sarkar et al (1990) analyzed a corefrom the eastern Arabian Sea and proposed thatwinter monsoon was stronger during LGM as evi-dent by enhanced NE monsoon current then. Theybased their conclusion on the negative excursionshown by oxygen isotope in four different speciesof foraminifera, which they attributed to influx ofenhanced low salinity water via the NE monsooncurrent and SST increase due to vanishing of SWmonsoon induced mixed layer deepening. Sarkaret al (2000) carried out oxygen and carbon isotopicanalysis on the planktic foraminifera, G. sacculiferand G. menardii, in cores off the west coast of India(3268G5, figure 4). They inferred that excess ofevaporation over precipitation (E–P) has decreasedsteadily from the 10 ka to ∼2 ka (figure 5E) imply-ing a steady increase in the SW monsoon pre-cipitation during the Holocene. Thamban et al(2001) measured the oxygen and carbon isotopesin planktic and benthic foraminifera along withCaCO3 and Corg content from near Cochin (GC-5,figure 4). They reported an glacial-interglacialδ18O difference of ∼2.1� (figure 5D) and signifi-cant fluctuation in δ18O during Holocene due tochanges in sea surface characteristics relatable tovariations in SW monsoon over land. They con-cluded that major increase in SW monsoon pre-cipitation occurred after ∼9 ka and in contrast towestern Arabian Sea records the productivity inthis region was lower during 13–6 ka and was maxi-mum between ∼18–15 ka, attributable to increasedwinter monsoon that led to greater nutrient injec-tion by enhanced convective mixing. Bhushan et al(2001) calculated the paleoproductivity using theburial flux of CaCO3 and Corg in 11 cores fromthe continental margin of eastern Arabian Sea andinferred that SW monsoon intensity increased from∼10 ka to ∼2 ka. Agnihotri et al (2003a) studiedvarious sedimentary proxies regarding productivitysuch as CaCO3, Corg, nitrogen, Sr and Ba, etc., inthe cores from eastern Arabian Sea and concludedthat surface productivity was lower during the lastglacial-interglacial transition and higher during theHolocene. Agnihotri et al (2003b) found increas-ing denitrification intensity in a core from theeastern Arabian Sea from ∼10–2 ka that impliesincreasing SW monsoon intensity during Holocene.Tiwari et al (2005a, 2006a) studied a core off


the Mangalore coast that yielded sub-centennialscale resolution, which helped them reconstructSW monsoon precipitation proxies, in contrast tomany studies from western Arabian Sea that reportSW monsoon wind strength. Periods of aridity wereobserved at ∼2 ka, 1.5 ka, 1.1 ka, 0.85 ka, and 0.5 ka;these appear to be widespread and are reflectedin diverse proxies from different geographical loca-tions surrounding it. Also, precipitation signalsfrom the eastern Arabian Sea exhibit a good coher-ence with the wind speed indicators from thewestern Arabian Sea on centennial timescale, dur-ing the past to ∼3 ka. High resolution data for moreremote past are yet to be obtained. The precipita-tion record also exhibit strong correlation with theTSI (total solar irradiance) variability and spec-tral analysis of precipitation proxies yield signifi-cant period of ∼200 yr, which matches with the∼200 yr Suess solar cycle indicating that SW mon-soon is governed by solar variability on centennialtimescales (Tiwari and Ramesh 2007).

4.3 Equatorial Arabian Sea

Studies from the equatorial Arabian Sea are verysparse compared to other parts of the Arabian Sea.A giant piston core MD900963 has been obtainedeast of the Maldives covering the past 910 ka.Rostek et al (1993) measured the oxygen isotopesin the G. ruber in this core and inferred that glacialstages were characterized by increased evapora-tion and/or decreased precipitation, which theyattributed to enhanced dry NE monsoon and/orreduced SW monsoon. Various other studies suchas Beaufort et al (1997), Rostek et al (1997),Schulte et al (1999), Pailler et al (2002) were car-ried out on the core MD900963. They inferred thatprimary productivity in this region was enhancedduring the glacial periods and was lower dur-ing the interglacial periods in contrast to otherproductivity records from the Arabian Sea. Theyattributed it to increased convective overturningdue to stronger NE monsoon winds that led tonutrient injection in the surface layer and henceincreasing the productivity. In all the studies, theproductivity proxies exhibit a ∼23 ky precessioncyclicity induced by insolation variations. Further-more, they maintain that deep water at this siteremained oxygenated for the past 350 ka. Saraswatet al (2005) reconstructed the SST for the past137 ky from this region using Mg/Ca ratio ofG. ruber. They observed that, during LGM, equa-torial Indian Ocean SST was lower than present by∼2.1◦C comparable to cooling observed in otherparts of the Arabian Sea and during last inter-glacial (isotopic stage 5e), SST were higher by∼1.5◦C than present. Tiwari et al (2005b, 2006b,c)studied a core, SS3827G (figure 4) from this region

for various isotopic (δ18O, δ13C, δ15N) and chemi-cal proxies (CaCO3 and Organic carbon content,C/N ratio) spanning the past ∼35 ka. Based onimproved chronology (AMS radiocarbon datingon selected planktic foraminiferal species), Tiwariet al (2005b) noted that during the early deglacialperiod (∼19 to ∼17 ka BP), δ18O values decreasedby about 1% in the surface dwelling G. ruber andG. sacculifer (figure 5A,B) due to influx of isotopi-cally lighter, low salinity water from the south-western Bay of Bengal via the strengthened NorthEast Monsoon Current. This indicates that thenortheast monsoon intensification occurred duringthe early deglacial period (∼19 to ∼17 ka BP)and not during LGM as proposed earlier based onbulk 14C dates (Sarkar et al 1990). Results furthershow that at the core site, minimum SW monsoonprecipitation occurred at the Last Glacial Maxi-mum, with a subsequent increase at TerminationIA. During the Holocene, SWM precipitation eitherintensified or stayed uniform up to the core top(∼2.2 ka), as revealed by generally decreasing δ18Ovalues (figure 5A,B). Variations in precipitation areconsistent with climate changes recorded in polarice sheets. Although the different resolutions of thetwo records preclude a rigorous comparison, abruptcooling/warming events appear to be accompa-nied by sudden reduction/enhancement in (SWM)rainfall. Thus mechanisms other than changesin the thermohaline circulation, with timescalemuch shorter than a millennium, for example,natural greenhouse warming (e.g., CH4 concentra-tion), modulated by emissions from the tropics,could have played a major role in high-latitudeclimate change (Tiwari et al 2006b). The pro-ductivity in this region is governed by the vari-ations in the Indian Ocean Equatorial Westerlies(IEW). The IEW in turn is positively correlatedto the southern oscillation index (SOI), related toEl Nino, SW monsoon, and east African rainfall(EAR). The productivity data show that Indianand east African rainfalls declined from ∼35 ka upto the LGM, with the maximum El Nino frequencyduring the last glacial period. From ∼14.5 ka to∼2 ka (i.e., core top), strengthening SW monsoonand EAR is observed as is also declining El Ninofrequency (Tiwari et al 2006c).

4.4 Bay of Bengal

This region records a strong signal of Indian mon-soon in the form of huge river discharge throughthe Ganges–Brahmaputra rivers (the peak summerdischarge is 50,000m3/s) with an average δ18Ocontent of −8� (Feng et al 1999) that reducesthe mean salinity by about 7% in its northern-most region (Laviolette 1967). But very few studieshave been carried out in this region as compared


to Arabian Sea, possibly because of the fact thathuge sediment load (2000 million ton per year) isbrought along with the river discharge with greatforce (forms the Bengal Fan, the world’s largestone, spreading for 3000 km) that intermingles thesediments of different ages and makes it very dif-ficult to obtain a undisturbed, turbidity-free sedi-ment core. Nevertheless a few, rare successfulattempts have been made starting with Duplessy(1982) and Fontugne and Duplessy (1986) whostudied cores from the northern Indian Oceanfor different time slabs. Recently, Chauhan andSuneethi (2001), Chauhan (2003) and Chauhanand Vogelsang (2006) have carried out studiesfrom different locations of the Bay of Bengalcovering eastern, central and western regions. Theyreport arid climate at 18–15, 12.5, 11.5, ∼4.8and 2.2 14C ka based on δ18O of G. sacculiferand clay mineral assemblage. Kudrass et al (2001)studied a core (126 KL) from the northernmostpart of the Bay of Bengal (figure 4) and recon-structed SW monsoon history for the last 80 kabased on salinity fluctuations derived from δ18Oof G. ruber and alkenone SST data. As evidentfrom figure 5K, they observed a strong correlationbetween the Indian summer monsoon variabi-lity and Greenland temperature record; intersta-dials (warmer, D-O events) and stadials (colderepisodes) correspond respectively to stronger andweaker monsoon. This observed correlation alongwith other such studies, led them to hypothesizethat the feedback processes involving snow anddust of the Tibetan plateau vary the summer mon-soon capacity to transport moisture into centralSouth Asia and into the atmosphere, which initi-ates, amplifies, and terminates climatic cycles inthe Northern Hemisphere. Ahmad et al (2008) ana-lyzed a core from the southern Bay of Bengal forδ18O and δ13C of G. ruber and benthic foraminifera(C. wuellerstorfi) spanning the past ∼60 ky. Theynoted two prominent negative δ18O excursions at∼8–7 and ∼18–20 ka and attributed to the suddeninflux of freshwater as a result of intensified mon-soonal precipitation, while large fluctuations inδ18O during the Holocene suggest variability inriverine input. They attributed δ13C variations tochanges in deep ocean circulation and postulatedthat during the glacial period, this region receivedlesser of the North Atlantic deep water (NADW)while contribution from the southern ocean deepwater may have increased.

5. Conclusions

Several studies discussed above have shown strongcorrelation between the North Atlantic climateand the SW monsoon on centennial to millennial

timescales. Abrupt climate changes (Dansgaard–Oeschger (D/O) oscillations and the Heinrichevents) documented in Greenland ice core records(Dansgaard et al 1993; Grootes et al 1993) areclearly reflected in the tropical SW monsoonrecords. But the exact mechanism linking the cli-mate of these far off regions is still not clear: severalmediums, such as thermohaline circulation changes(e.g., Broecker 1994) point to the North Atlanticas the driver of the global climate. The naturalgreenhouse gas concentrations modulated by cli-mate conditions in the tropics (Manabe et al 1991;Schulz et al 1998; Tiwari et al 2006b), and theatmospheric water vapor governed by SW mon-soon strength (Kudrass et al 2001) are suggested tobe the principal factors governing the world’s cli-mate. Recently, Goswami et al (2006) have identi-fied a coupled ocean-atmosphere phenomenon thatprovides a new teleconnection between the NorthAtlantic climate and the SW monsoon on inter-decadal timescale and is suggested to be effectiveon longer timescales as well. It involves the AtlanticMultidecadal Oscillations (AMO), a basin widephenomenon of oscillating sea surface temperatures(SST) with period of 65–80 years (Kerr 2000), andthe North Atlantic Oscillation (NAO), a sea levelpressure fluctuation between the Icelandic Lowand the Azores High (essentially wind strength)with strong annual and weak inter-decadal perio-dicities (Wallace 2000). AMO modulates the NAOin such a way that the strong negative NAOevents influence the winds and storm tracksthat lead to tropospheric temperature anomaliesover Eurasia. These anomalies decrease (increase)the meridional gradient of tropospheric tempera-ture resulting in below (above) normal monsoonrainfall.

On a glacial-interglacial timescale, the broadpicture regarding the paleomonsoon variation thatemerges out of the above discussion is that SWmonsoon was stronger during interglacials andweaker during glacial periods with the NE mon-soon exhibiting the opposite behavior. A first-orderunderstanding is that the snow cover increased overthe Tibetan plateau/central Asia during LGM thatreduced the land-ocean temperature contrast dur-ing summer and enhanced it during winter that ledto weakened SW monsoon and strengthened wintermonsoon. On exploring further, during LGM,we find that NE monsoon strengthened duringthe early deglacial/late glacial period (∼19–17 ka;Tiwari et al 2005b). During winter, the sea inthe southern tropical Indian Ocean is warmer(low-pressure region) than the cooler Asian land-mass (high-pressure region) resulting in NE mon-soon. Earlier studies proposed that during glacialperiods, such as LGM, land was cooler thanpresent (due to permanent ice cover) in the vicinity


of a warmer (than land) ocean resulting in therelatively strong NE monsoon. But during LGM,the SSTs also fell by ∼2◦C in the tropical regions(Rostek et al 1993; Saraswat et al 2005), that is,oceans were not much warmer than the land. Theoutcome was that NE monsoon was not as strongas it could be. But during the early deglacial period(∼19–17 ka), the SSTs in the equatorial IndianOcean rose by 1◦C (Rostek et al 1993; Saraswatet al 2005; Anand et al 2008), which accountsfor about 50% of the total LGM to Holocenetemperature change, resulting in warmer oceansand strengthening of the NE monsoon. On longertimescales (orbital), the intensity of monsoon isbasically controlled by the sun-earth geometry;mainly by precessional cycle of the earth. It hasbeen verified using the paleoclimate models thatthe tropical climate is affected more by the preces-sion of the earth’s perihelion, while high latitudesare more affected by the changes in earth’s obliq-uity (Kutzbach 1981; Jagadheesha et al 1999b).

During the early Holocene, the SW monsoonintensified with a maximum at ∼8–9 ka and hasexhibited centennial/millennial scale fluctuations.On centennial timescale, a study has shown thatpast fluctuations in SW monsoon precipitationover the Indian subcontinent followed the windintensity records from the western Arabian Seaon (Tiwari et al 2006a). But if we considerthe long-term trend throughout the Holocene,as inferred from the upwelling proxies from thewestern Arabian Sea, which are manifestationof SW monsoon wind alone, the SW monsoonappear to decline in accordance with the decreas-ing summer insolation. On the contrary, precipi-tation proxies from the eastern Arabian Sea andother studies from Arabian Sea/Bay of Bengal indi-cate that SW monsoon has consistently increasedfrom 10 ka to ∼2 ka or has shown decline onlyafter ∼5.5 ka; in the Holocene SWM did notfollow insolation, highlighting the importance ofinternal feedbacks. Thus more well dated, high-resolution cores are needed to answer the ques-tion that how good is the correlation between thewind and rain proxies from the western and easternArabian Sea on different timescales and whetherincreasing wind intensity in the west (sea) alwaysfavours enhanced precipitation in the east (land)or not.

Acknowledgements

We thank ISRO-GBP for funding, Dr M Sudhakarand Director, NCAOR for encouragement. We alsothank the editors of this volume for the invitationto write this review. This is NCAOR contributionno. R-48.

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Spatial and temporal coherence of paleomonsoon records ......520 MANISH TIWARI et al qualitative climate proxies omitted here, reference is made to Jagadheesha et al (1999a), Korisettar