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CHAPTER 5

DISCUSSION AND CONCLUSION

5.1 PROVENANCE OF THE MID-HOLOCENE MARINE

SEDIMENTS

Marine sediments are mostly derived from the recycling of

continental materials (Windom 1976), thus changes of weathering intensity in

the source areas are recorded in the marine sediments. Climate, weathering

processes, erosion and tectonics are the main factors for controlling the degree

of chemical weathering and generating detritus at source areas. The main

parameters that control chemical weathering on continents consist mainly of

precipitation and temperature and tectonic factors (Milliman & Syvitski1992,

Raymo et al 1988, Raymo & Ruddiman 1992, White & Blum 1995, Berner &

Berner 1997). Physical denudation of the continents and the products of

continental weathering such as the detritus silicates are major components of

marine sediments (Fu et al 2011).

Chemical weathering is principally driven by rainfall in tropical

regions, with temperature being almost constant (Warrier &Shankar 2009).

Weathering in the Andaman and Nicobar ridge terrain is enhanced due to the

high monsoonal activity and the deposition of detritus sediments in the ocean

reveals physical and chemical erosion of the terrain over which they flowed.

This is strongly dependent on the monsoonal intensity, especially the SWM

(Bhushan et al 2007, Sarin et al 1979) and also by sea level change.

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The intense summer monsoon rain over the southwest Asian

continent drains into the Bay of Bengal and Andaman Sea through the major

rivers such as Ganges, Brahmaputra, Irrawaddy and the Salween rivers

(Duplessy 1982). But the Andaman fore and back arc basins near the

Andaman ridge accumulates a little amount of detritus sediments from the

local streams due to their input since these are accretionary prism and are the

result of highly oblique subduction at the western Sunda Trench (Cochran

2010).

The important aspect of provenance study is to infer the sediments

derived from such as source rock, source area, and tectonic setting etc.

(Schieber 1992).Geochemistry of sedimentary rocks is widely used to

discriminate the tectonic setting of the sedimentary basin (Bhatia 1985; Roser

&Korsch 1986; Floyd &Leveridge 1987). The immobile chemical elements

such as Al, Ti, and Zr are extremely helpful for estimating the nature of

source rock (Taylor &McLennan 1985). In the present study geochemical

variations in the sediments have been applied to infer their source.The

bivariate plot of TiO2 Vs Zr as illustrated by Pearce (1982) (Figure 5.1a) to

differentiate the mid oceanic ridge basalt (MORB), within plate basalt (WPB)

and volcanic arc lavas, was plotted to understand the origin of the sediments.

The plot points to volcanic arc lava source. This inference is supported by the

several ternary plots, such as the ternary plot of MnO *10 - TiO2 - P2O5 * 10

proposed by Mullen (1983) that suggests that the sediments are derived from

the Ocean Island Arc (OIA) (Figure5.1b). This inference is also supported

by a comparison of the present study trace element data and their ratio with

the trace elements values of mudrocks collected from various tectonic settings

(Bhatia 1985) (Table 5.1).

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The ternary plot of MgO - FeO (total) - Al2O3 percentage (Pearce

et al 1977) and Ti/100 - Zr - Y*3 (Pearce & Cann 1973) exhibits the source

from the ocean floor basalts (OFB) (Figure 5. 2 a, b).

Figure 5.1 A) Bivariate plot of TiO2 vs Zr (ppm) followed after Pearce(1982), B) Ternary plot of MnO*10 - TiO2 - P2O5*10 drawnafter Mullen (1983) genesis of sediemtns from the Arc lavasand from the ocean Island arc.

Table 5.1 Comparison of trace elements and their ratio with the traceelement characters of mudrocks from various tectonicsettings (after Bhatia 1985)

Elements BengalBasin OIA CIA ACM PM

Present work(Average of24 samples)

Nb(ppm)

16.00 3.70 9.00 16.50 15.80 1.96

Th (ppm) 18.00 5.50 16.20 28.00 22.00 3.00Cr (ppm) 99.00 39.00 55.00 58.00 100.00 101.1Ni (ppm) 42.00 15.00 18.00 26.00 36.00 23.1Zr/Th 13.30 28.00 12.00 7.00 7.00 20.72Zr/Nb 14.70 38.00 21.00 11.00 10.00 29.98Nb/Y 0.50 0.17 0.35 0.50 0.54 0.12

[ OIA-Ocean Island Arc, CIS - Continental Island Arc, ACM - Active

Continental Margin, and PM- Passive Margin]

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Figure 5.2 A) Ternary plot of MgO - Feo - Al2O3 proposed by Pearceet al (1977) and Panel B) ternary plot of Ti/100 - Zr - Y*3following Pearce &Cann (1973) points to a source from theocean floor basalt (OFB).

These inferences are further corroborated by the Rosser & Korsch

(1986) bivariate plot of Log K2O/Na2O against SiO2 percentage (Figure 5.3)

that points to the sediments derived predominantly from the oceanic island-

arc setting. The bivariate plot of SiO2/Al2O3- Log K2O/Na2O (Figure 5.4) also

supports that the sediments are a consequence of tectonic settings of A1 and

A2, i.e. from the basalt and andesitic detritus.

A binary plot of SiO2 contentto the ratio of Zr/TiO2 was put forward

by Winchester & Floyd (1977) to classify the type of source rock, and this

reveals that the sediment is largely of andesite composition (Figure 5.5).

Further, the ratio of Zr/Ti and Nb/Y shows that the sediments are composed

basaltic andesite (Winchester& Floyd 1977) (Figure 5.6).

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Figure 5.3 Bivariate plot of log K2O/Na2O against SiO2 % points to asource from the Island arc settings (after Rosser &Korsch1986).

Figure 5.4 Bivariate plot of SiO2/Al2O3- Log K2O/Na2O exhibiting asource from A1 and A2 settings (Rosser &Korsch 1986)

All the geochemical parameters points to an basalt source (depleted

in free silica content) and this is in agreement with Srivastava et al (2004)

who suggested that the Andaman ophiolites are mainly composed of basalt

with some andesite and trachy basalt.

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Figure 5.5 Bivariate plot of SiO2 with ration of Zr/TiO2 afterWinchester &Floyd (1977).

Figure 5.6 Bivariate plot of the ratio of Zr/Ti versus Nb/Y indicatingan andesite/basalt supply after Winchester and Floyd1977)

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The CIA and ICV weathering indicators were used to deduce the

weathering trend of the sediments. The Chemical index of alteration (CIA)

value calculated following Fedo et al (1995) varies between 50 and 60,

indicating an early stage of weathering ie. low chemical alteration.. The Index

of Compositional variability (ICV) was calculated as proposed by Cox et al

(1995). The ICV values range from 1.50 to 1.83. The bivariate plot of CIA

and ICV (Potter et al 2005) indicates that the sediments follow a basaltic

weathering trend (Figure 5.7)

Figure 5.7 A plot of Chemical Index of Alteration (CIA) and Index ofChemicalVariation (ICV)show a trend of low chemicalweathering of basalt

Recent studies carried out by Ramaswamy et al (2004) andRao

et al (2005) indicate that the sediment influx from the Irrawaddy and Salween

rivers (the sediments that are discharged from Myanmar and drain the

metamorphic and intrusive rocks catchment in the NE Himalayas (Chauhan

et al 1993) is transported along the shelf by monsoonal currents and carried to

the deep sea floor through submarine canyons. Apart from fluvial sources

(Roonwal et al 1997), the Andaman Forearc Basin (AFB) also receives

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detritus from pelagic, eolian, weathering of sea-floor rocks, and possible

hydrothermal sources (Rao et al 1996, Chernova et al 2001, Venkatesan et al

2003, Siby 2004).

In the present study smectite and illite are the two principal clay

minerals present in the sediment core other than kaolinite and chlorite. The

high smectite content in the sediment core suggests that the sediments are

derived mainly from nearby basic rock source (Griffin et al1998, Siby et al

2008). Illite and chlorite are often derived mainly by physical erosion or

degradation of metamorphic and granitic rocks or from the erosion of

sedimentary rocks (Liu et al 2005). Illite is a common clay mineral, which is

found in almost every river suspension (Potter et al 1975). The catchment area

of the rivers Irrawaddy and Salween (which discharge from Myanmar and

have catchment in the NE Himalayas) comprise metamorphic and intrusive

rocks (Chauhan et al 1993). Thus it can be inferred that considerable amount

of fine grained illite and chlorite sediment deposition near the Landfall Island

is controlled by the Irrawaddy and Salween Rivers.

SWM controls precipitation in BOB (Ding 1994) and the

strengthened winter monsoon causes deep chemical weathering of the

bedrocks (Wang 1999) resulting in deep weathering of the nearby continental

margins of the Andaman Islands. Thus, the terrestrial detritus in the sediments

in northern Andaman Islands exhibit stronger chemical weathering signals

(Wei et al 2003).

A comparison of major elements data from the present study

(Table 5.2) with the other sources contributing sediment load to the Bay of

Bengal basin reveals that the Irrawaddy delta to be one of the significant

sources. This is based on comparing the data on Al and Fe content from the

Irrawady and the present study.

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Table 5.2 A comparison of major and trace elements with the othermajor source of deposition to Bay of Bengal Basin

Metals Gangesa

Brahmaputraa Coastal

BOBb Irrawaddy c

Presentwork

(n = 24 )Al (%) 4.66 5.6 2.07 8.68 7.70Fe (%) 2.16 2.9 1.5 5.95 3.68Mg (%) 1.32 1.66 1.72 - 6.48Ca (%) 2.34 1.93 0.91 - 1.28Na (%) - - 1.18 - 3.75K (%) 1.33 1.24 0.49 - 1.43Ti (%) 0.3 0.31 - - 0.55V (ppm) 86 137 - - 40.2Cr (ppm) 52 100 57 156 101.1Ni (ppm) 20 47 30 140 23.1Cu ppm) 21 17 20 31 16.0Ba (ppm) 348 347 58.6 - 57.8

a. Biksham and Subramanian (1988), b. Selvaraj et al(2004), c. Siby et al (2008).

In order to understand the relationship within the major oxides of

the sediment and their controlling factors, Pearson correlation coefficient was

carried out using the statistical software (SPSS, V-21)(Table 5.3).Positive

correlation was observed between Al2O3 and Sr, CaCO3, K2O, CaO, P2O5 and

MgO, while SiO2 shows negative relationships with all the major oxides

barring Sr.SiO2 shows a strong positive correlation with Zr (0.82) and a

significant negative correlation with Sr indicating its concentration in

minerals other than quartz (Das & Haake 2003). The significant negative

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Table 5.3 Correlation matrix of major oxides (wt %), CaCO3 (%), Sr and Zr (ppm)

MajorOxides

SiO2 Al2O3 K2O CaO TiO2 Na2O MgO P2O5 MnO Fe2O3 CaCO3 Sr Zr

SiO2 1.00Al2O3 -0.61 1.00K2O -0.58 0.72 1.00CaO -0.82 0.71 0.76 1.00TiO2 0.28 -0.31 -0.10 -0.14 1.00Na2O -0.34 -0.36 -0.34 -0.18 -0.14 1.00MgO -0.86 0.56 0.54 0.89 -0.32 0.02 1.00P2O5 -0.82 0.82 0.71 0.89 -0.21 -0.18 0.84 1.00MnO 0.15 -0.16 0.14 0.01 0.42 -0.12 -0.26 -0.14 1.00Fe2O3 -0.10 0.30 0.63 0.34 0.43 -0.38 -0.03 0.22 0.66 1.00CaCO3 -0.71 0.68 0.45 0.80 -0.30 -0.20 0.90 0.86 -0.43 -0.13 1.00Sr -0.75 0.66 0.80 0.93 0.06 -0.23 0.77 0.87 0.13 0.54 0.68 1.00Zr 0.82 -0.71 -0.63 -0.90 0.21 0.11 -0.88 -0.85 0.16 -0.15 -0.85 -0.82 1.00

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correlation of between SiO2 and Al2O3 indicates that the majority of the SiO2

is not associated with Al2O3. This is probably due to presence of SiO2 as free

quartz, and anticorrelation between SiO2 and Fe2O3 suggests that the

occurrence of these oxides have a contribution from different sources and

rock types (Das &Haake 2003, Ogala et al 2009).

Al2O3 is correlated well with K2O (0.72), CaO (0.71), MgO (0.56),

and P2O5 (0.82). CaO shows a strong correlation with Sr (0.93), P2O5 (0.89)

and MgO (0.89) suggesting that Sr is associated with calcium carbonate and

P2O5 as apatite (Nath et al 1989). Negative correlation between Zr and CaO (-

0.90) suggests that the detritus component comprises a mixture between arc-

derived and continental detritus (Gamble 1996).

The bivariate plot of all the major oxides against Al2O3

(Figure 5.8) exhibits a positive trend with CaO, K2O and P2O5, whereas SiO2

evidently shows an reverse trend, that points two different sources (Ogala

et al 2009).

5.1.1 Neodynimum

According to Amakawa et al (2000) the concentration of Nd in the

BOB are much higher near the coast. Dissolved Nd from rivers and release of

Nd during reaction of river sediments with seawater in estuarine settings, is

the main source of Nd to surface waters. Stoll et al (2007) found that in the

entire Indian Ocean highest dissolved Nd concentrations are in the BOB and

Andaman Sea, and these basins are characteristic of very high input of

riverine sediments.

It is also noted by Stoll et al (2007) that the BOB has significant

isotopic heterogeneity among the principal river tributaries such as the

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Ganges, Brahmaputra drains from Himalayan crystalline series in the north,

the Irrawaddy drains the eastern and northern Indo-Burman Ranges and

eastern Himalaya and Indo-Burman ranges (Figure 5.9).

Figure 5.8 Major oxides of the marine sediment core plotted againstAl2O3 percentage.

In order to delineate and characterise the source of the Nd data of

the present study is compared with the other major sources contributing the

radiogenic isotopes into the basin (Figure 5.9). It evidently shows that

Irrawaddy has considerable contribution of sediments in this region as

suspended particles, this is also supported by the eastern BOB and

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Upper - Middle fan Nd (0) values (Colin et al 1999). The Nd (0) values also

suggeststhat the sediments are of Holocene to late Holocene in age, Since

these Nd(0)values are similar and equivalent to the Nd (0) values of Stoll

et al (2007) as he differentiated the Nd values for Late Glacial Maxium

(LGM), Holocene (H) and Late Holocene (LH) (Figure 5.10).

Figure 5.9 A comparison of Nd (0)values of the present study with theother source of Nd into the basin. [1) Colin et al (1999), 2)Galy &France-Lanord (2001) 3) Singh &France-Lanord(2002) 4) Szulc et al (2006) 5) Robinson et al (2001) 6) DeCelles et al (2004) 7) Najman et al (2008) 8) Allen et al (2008)9)Pierson-Wickmann et al (2001) 10) Kessarkar et al (2005)11) Fagel et al (1994, 1997)].This shows that the sedimentsare from Irrawaddy river, it is supported by the easternBOB and upper and middle distal fan Nd data

The chemical, textural and isotopic signatures observed at the AFB

sediment core near the Landfall Island reflects a mixture of sediments

distinctly from a nearby source i.e from the Andaman terrain of ophiolite

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basalt suites and sediments from Irrawaddy, Salween and Sittang rivers that

transport weathering products from the land masses to the AFB. The

sediments derived into the basin are contributed not only by the local

ophiolite provenance but it also from the quartz-rich sediments derived

dominantly from the north Irrawaddy delta (Pal et al 2010). This inference

also supports the argument suggested by Moore et al (1982) that the sediment

are transported from Irrawady delta to the south of Sumatra along the

Andaman Fore arc basin. However the sediment fluxes vary during the dry-

warm/wet phases.

Figure 5.10 A comparison of the epsilon Nd (0) with the values of lastglacial maximum (LGM), Late Holocene and the Holocene(after Stoll et al 2007)

5.2 MIDDLE HOLOCENE PALAEOCLIMATE

From the review of literature it has been observed that Late

Quaternary palaeoclimate record has been reconstructed using marine

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sediment cores with the application of various proxies such as geochemistry,

oxygen isotope, and clay mineralogy etc. Palaeoclimate reconstructions are

predominantly from the Arabian Sea marine sediment cores as compared to

the BOB (Chauhan et al 1993, 2004). Based on the marine and terrestrial

records, Patnaik et al (2012) advocated that the Holocene period is

characterized by a strong monsoonal phase known as “Holocene Climatic

Optimum (HCO)” i.e. the middle Holocene period, that subsequentlywas

followed by weak monsoonal phases around ~2500-1500, 1000, 650-450

years BP and the Little Ice Age (LIA) (1450-1850 AD). Chauhan et al 2004

and Chauhan & Vogelsang (2006) found an intense weakening of SWM

~ 4300 and 2200 – 1800 years BP. Sarkar et al (1990, 2000) and Yadava &

Ramesh (2005) found an increase in the intensity of SWM around ~ 3200

years BP.

Hence in the present study a detailed analysis of sediment texture,

clay mineral composition (clay mineralogical aspects such as occurrence of

kaolinite, smectite, chlorite, illite, K/C, C/I)major oxides, Nd(0) and its

distribution along the marine sediment core collected near the Landfall Island,

Bay of Bengal, supported by stable oxygen isotope ( 18O) data on G. ruber

and Phytolithhas beenused to infer climate regime responsible for their

deposition.

5.3 CLAY MINERALOGY

The production, release and deposition of clay minerals are largely

related to the geology of the region, its drainage and climate (Weaver 1989).

The sediment load of the Himalayan and the Indian peninsular rivers are

markedly different (Konta 1985, Subramanian 1987), and these have been

distinguished and identified as the clay loads of the Himalayan and the

peninsular rivers respectively. Studies carried out on the clay minerals

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(Rao 1983) near the study area reveal that the clays and the very fine

sediments are rich in kaolinite, chlorite and illite whereas montmorillonite

(smectite group) is less. Rao (1983) found two distinct clay mineral

assemblages: i. Fe-rich chlorite in the eastern side, ii. Fe-poor chlorite in the

western side of the Andaman Islands, however this study was not quantified.

The fluctuations in the magnitude of the monsoon influence the

formation and release of clays and detritus, temporally and spatially

(Thamban et al 2001). Clay minerals are brought by the river systems

surrounding the BOB (with a major contribution from Irrawaddy, Salween

and Sittang and, more importantly, the Ganga-Brahmaputra and the Andaman

Islands). Clay mineralogy analyses signify alteration/erosion in the catchment

area. Release of sediments from the parent material and formation of clays

depend upon the climate, source rock characteristics and relief of the area

(Birkland 1974). It is well accepted that the formation and occurrence of

kaolinite and gibbsite indicates a very humid climate, illite and smectite

points to a moderate-humid climate and chlorite by physical weathering under

arid cold climate (Biscaye 1965, Griffin 1968, Chauhan et al 1993).

The variations in the fine grain detritus particularly in clay mineral

assemblages can be applied to delineate the intensity of the SWM. The clay

mineral assemblage and coarser sediment layers as observed in the sediment

core points towards two major incursions of intense SWM during ~ 5400 and

~ 3400 years BP with higher sand and silt components. A dominant

occurrence of sand at the depth 120 - 110 cm (~ 6300 to 6000 years BP),

90 - 85 cm (5400-5300 years BP) and 45-35 cm (3500 -3200 years BP)

followed by an predominant occurrence of silt and sand at the depth

45- 35 cm (3500 - 3300 years BP) is observed.

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High content of silt occurs at the depth 60-55 cm (4000 - 3900years BP), while laminar clays and fine silt are deposited at the depth 35 cm(3200 years BP). Sediment texture and clay mineral composition reveal lowclay content (16 %), with higher coarser components at the 40 cm depth andthis indicates an high energy event, unaffected chemical weathering of theLandfall Island bed rocks (conglomerates, sandstone, shale and the ophiolitesuite of rocks). This clearly shows that the rate of sedimentations has beennon-linear. It is also noted that the decrease in the sedimentation rate(Figure 5.11) post 4300 years BP that points to probably a thicker vegetationcanopy in the hinterland and with low carbonate content indicating strongchemical weathering of the Island bed rocks. The low values of calciumcarbonate and organic matter also points to low productivity due to reducedSWM or dilution by terrigeneous material.

Clay mineral and 18O (Globigerinoides sacculifer) record from themarine sediment core BOB reveal an intense arid climate from 18, 000 to 15,000 years BP with fluctuations between 15, 000 and 13, 000 years BP(Chauhan & Suneethi 2001). A humid phase occurred around 12, 000 yearsBP and an arid phase from 10, 300 years BP to 11,500 years BP with twoextreme arid events at ~4800 years BP and 2300-2200 years BP (Chauhan &Suneethi 2001, Chauhan et al 2004) (Table 1.1).

Based on 13Corg, major and trace elements compositions and pollendata, the monsoon variability deduced from the Ganga plain lake-fill depositsindicate a dry period was evidenced from 15,000 to 13,000 years BPsubsequently followed by a humid/ wet period from 13,000 - 5800 years BP.This wet period eventually passed into a dry event from 5000 to 2000 yearsBP. Palaeoclimate and archaeological records of eastern Mediterranean, WestAsia, and the Indian subcontinent show abrupt climate changes in theHolocene period with widespread drought around 8200, 5200 and 4200 yearsBP (Staubwasser & Weiss 2006) (Table 1.1). These authors reported that

Figure 5.11 Down core variation of sand, silt, clay, Organic matter (OM), Calcium carbonate (CaCO3), Oxygenisotope ( 18O ‰) and Carbon isotope ( 13C ‰) with rate of sedimentation

105

during the Holoceneperiod, the Indian monsoon weakened over the northern

(Ganges and Indus catchments, Western Arabian Sea) region, whereas the

southern (Peninsular India) region experienced an increase in summer

monsoon rains.It has been postulated that occurrence of kaolinite (abundant in

the tropics), chlorite and illite (prevailing at high latitudes) are directly

inherited from the continents, and particularly from the tropical soils (Thiry

2000, Chauhan & Suneethi 2001).

The BOB (area 2.2 million km2) (Landfond 1966, Stow &Cohran

1989), receives several-fold higher fluvial influx particularly from the

Himalayan, Irrawaddy and the Salween and the Peninsular rivers as compared

to the Arabian Sea. This affects the sea surface salinity (SSS) and temperature

of the BOB sea surface water (Levitus et al1994). Planktonic foraminifer are

sensitive to sea surface salinity (SSS) and sea surface temperature (SST)

variations, the 18O of the foraminifera test composed of CaCO3 which is a

reliable proxy for understanding palaeoclimate changes (Chauhan 2003 and

references therein). The palaeoclimatic variations inferred from 18O in the

planktonic foraminifers species G. sacculifer (without sac), and

Globigerinoides ruber (G. ruber) mostly sustain in the upper mixed layer, and

their shell secretion is influenced by SSS and SST changes (Duplessy 1982).

Griffin et al (1968), Liu et al (2004), Weaver (1989), Chauhan &

Suneethi (2001), and Chauhan et al (2004) suggested that the clay minerals

such as chlorite and illite are often produced under the arid/cold climate while

the smectite and kaolinite are formed under humid/warm conditions causing

strong weathering coupled with intense SWM.

Occurrence of varying amounts of kaolinite, smectite, chlorite,

illite, kaolinite/chlorite (K/C) and chlorite/illite (C/I) ratio suggest periods of

intense wetter conditions (intense SWM) with incursions of shorter weaker

monsoonal conditions from mid to late Holocene. High kaolinite, K/C ratio

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with reduced chlorite and low C/I ratio around 6500 to 6000 years BP

(124-111 cm depth) points to intense SWM causing strong weathering of the

bedrocks in the hinterland. The reduced CaCO3 content (72.32- 70.63 %) and

occurrence of kaolinite, around mid Holocene 3300 years BP and 6000 years

BP (depth 38-40 cm and 108-110 cm) respectively indicates strong

weathering of bed rocks, intensified SWM that led to an upwelling resulting

in higher surface biogenic productivity. This observation is also supported by

the large coarser component influx at these depths.

In the studied sediment core an overall fluctuating occurrence in the

abundance of smectite, kaolinite, illite and chlorite followed by K/C and C/I

ratio is observed from 6000 to 2000 years BP (111 to 20 cm depth) indicating

an overall weakening of SWM. However, within this 4000 years period a

short incursion of extreme weakening of SWM occurred from 4400 years BP

to 4200 years BP i.e. from 70-65 cm depth (Figure 5.12) as revealed by high

smectite content, reduced kaolinite and illite clay minerals with high C/I ratio.

The characteristic clay mineral association of smectite and illite

also implies dry to semi-drier conditions around 4400 to 4200 years BP.

Subsequently an amelioration of SWM is observed with high percentage of

kaolinite and chlorite with higher C/I ratio from 3900 years BP to 3600 years

BP (depth 55 - 48 cm). An occurrence of arid phase is noticed around ~2000

to 1700 years BP (20-15 cm) due to the presence of higher illite content and

reduced C/I ratio. Similar observations were made by Chauhan &Vogelsang

(2006) using clay mineral composition of the sediments from the western

BOB. They reported (Table 1.1) four arid phases during the Holocene period

(10, 000- 9600, 7300, 5800 – 4300 and 2200-1800 years BP) and our

inferences gains support from the observations made by Chauhan &

Vogelsang (2006) and Staubwasser & Weiss (2006).

Figure 5.12 Down core variation of clay minerals , Sand, Silt, Clay, CaCO3, Organic matter (OM), Smectite (S), Illite(I), Kaolinite (K), Chlorite (C), ratio of K/C and C/I with calibrated age.

108

5.4 PALAEOCLIMATE IMPLICATION FROM

GEOCHEMICAL DATA

Marine sediments are made up of the weathering products from

continents (Windom 1976), so changes, shifts and weathering intensity in

source areas are archived in marine sediments. Geochemical data indicate that

Al is resistant to leaching during chemical weathering and is enriched in

weathering products (Nesbitt & Young 1982). Oxides such as CaO, MgO, and

Na2O are considered more soluble and mobile, while Al2O3, SiO2, and TiO2

are considered more insoluble and resistant oxides (Mackereth, 1966

Engstrom & Wright 1984). High molar ratios of (CaO + MgO + Na2O)/Al2O3

calculated for chemical weathering intensity (CWI) of sediments indicate

stronger weathering and presumably wetter conditions and a strong monsoon,

while low CWI reflects a less weathering and drier conditions (Sun et al

2009).

Moreover, normalization of major elements such as Mg, Ca, Na by

Al for the bulk sediments can be related to wet and dry cycles(Sun et al 2010)

Al is insoluble under both toxic and anoxic conditions and is a common

terrestrially derived element (Brown et al 2000). Al is generally used to

estimate the percentage of the terrestrial materials in marine sediments, and to

subtract the contribution of the detritus materials when calculating the

authigenic components of marine sediments for both pelagic and terrigenous

materials-dominated sediments (Murray & Leinen 1996, Klump et al 2000,

Wei et al 2003). The vertical profiles of the Na/Al, Ca/Al, and Mg/Al ratios

reflect the abundance of detritus components, which can be considered as

proxies for climate change in the provenance region (Wei et al 2003).

From the chemical data generated it is observed that significant

shifts in Ca/Al, Mg/Al, Na/Al ratios and CWI values generally occurred

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during major palaeoclimate changes. The increasing trend of CWI values,

followed by high Ca/Al, Na/Al, Mg/Al ratios indicate wetter conditions

~5486 years BP (98-90 cm) and from ~3425years BP to ~3214 years BP (42-

36 cm) (Figure 5.13) and intense ISM. Low ratios of Na/Al, Mg/Al, Ca/Al

and CWI from ~4270 years BP to ~4059 years BP (66-58 cm) indicate dry

and warm conditions and weak ISM. The magnitude of change in these ratios

over wet/dry cycles is nearly 20-22% indicating that such variation in patterns

is significant for these sediments and therefore, the element ratios and CWI

values of the studied sediment core are climatically responsive (Sun et al

2010.).According to Xiao et al (2006) CaCO3 is a good indicator of

temperature variation. Higher content CaCO3 (75.56 %) percentage indicates

warmer and dry conditions in the hinter land (around 66-64 cm, 4270 years

BP) (Figure 5.13).

In the present study, high sand content in the sediment core

occurring at depths 114-112 cm (~6191 years), 90-88 cm (~5486 years) and

42-40 cm (~3425 years BP) with less CaCO3 content indicates terrigenous

influx of sediments and wet conditions (intense monsoon). The increasing

OM content also point towards wetter conditions. Moreover, a virtual increase

in sand component (from 8.28% to 20.02%) and the decrease in clay content

(from 23.67% to 16.00%) between ~ 4059 to 3425 years BP suggests high

terrigenous input from the surrounding islands and an unaffected chemical

weathering of the Landfall Island bed rocks during the mid and late Holocene

period. The sand layers with less silt component during this period also points

to the river influx such as Irrawaddy, Sittang and the Salween rivers. On the

contrary, from the depths such as 96 - 94 cm (~ 5655 years), 60 - 58 cm (~

4059 years) and 24 - 16 cm (~2038 years) the sediments are dominantly fine

silt with decreasing trend of sand content and elevated CaCO3percentage

indicating warm and dry conditions (Figure 5.14).

Figure 5.13 Down sediment core variation of Na/Al, Mg/Al, Ca/Al, CaCO3 (%), CWI, stable isotope of 18O and 13Cof G. ruber species. red lines represent the four significant steps of palaeoclimatic changes ages when theseprincipal changes occurred.

111

A clear positive correlation between sand content and Nd and a

negative correlation between silt content and Nd implies that during warm

and dry climate. Nd seawater values tend to be less radiogenic than during the

wetter phases. It is observed that when the sand content is high, silt content is

low, Nd values high (more radiogenic) and CWI values is also high indicating

intense SWM. While in turn, when the sand and clay contents are low, silt

percentage is high, Nd values low (less radiogenic) the CWI values are low

pointing to weaker SWM. Considering that Nd values can be explained by a

mixing between two end members (one less radiogenic and the other more

radiogenic), a change in the source input related to the climate (warm and dry

to wet conditions or the reverse situation) has directly affected the Nd

seawater values, as well as the other geochemical data, at the AFB core.

These changes on the source input however must be interpreted within the

regional geological and hydrological contexts.

Figure 5.14 Down sediment core variations in sand, silt, claypercentages, Nd, and CWI ratios, red lines are presentedthe age values of these principal shifts

112

Integration of all these results indicate a SWM palaeoclimate

variability since mid Holocene period: wet conditions prevailed from 6500

years BP to 5486 years BP, then dry and warm conditions with a peak at 4059

years BP, followed by a major wet phase reaching its maximum at 3425-3214

years BP, further followed by a dry phase since that time to the present. This

inference drawn based on multiproxy approach is in agreement withChauhan

& Suneethi (2001) who suggested extreme arid events at ~4800 years BP and

~ 2300 to 2200 years BP and Sharma et al (2006) who showed an

amelioration in climate since ~1700 years BP. Chauhan &Vogelsang (2006)

also revealed an arid phase between 5800 - 4300years BP and 2200-1800

years BP from a clay mineralogical study.

5.5 STABLE ISOTOPE INFERENCES

The 13C of G. ruber varies from 0.70 ‰ (16-18 cm; ~1700 years

BP) to 1.33 ‰ (112-114 cm; ~6100) years BP. These variations in the 13C

are not coherent with the Holocene variations however we attribute the shifts

in 13C to the variations in carbon isotope composition of surface waters due

to the strengthening or weakening of SWM and wind induced variability.

During enhanced high wind velocity period, sub surface water with low 13C

might have been brought from to the surface, resulting in depletion of 13C

values and vice versa (Ahmad et al 2012). Our 13C values conform well with

the values of planktonic foraminifera 0.74 ‰ to 1.25 ‰ from the sediment

core SK157-14 (Ahmad et al 2012) data.

Duplessy (1982) analysed the variation in the stable isotope G. ruber18O) from the marine sediment cores collected from the BOB, Andaman Sea

and Arabian Sea. The data ranges between -3.70 ‰ and -1.58 ‰ for the

Holocene times and for the glacial times it varies from +0.55 to -1.17 ‰.

Chauhan (2003) also observed the heavy 18O values of G. ruber during

20,000-15, 000 years BP (- 0.9‰) and lighter 18O value of G. ruber

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in the BOB is around - 3.38‰ in the Holocene period, which is much lighter

compared to the Last Glacial Maxima (LGM). Thus a comparison of stable

isotope data reveals that the beginning of the Holocene is marked with much

lighter 18O values which persisted for the entire Holocene, with two pauses

at ~ 5000-4300 years BP and ~ 2000 yearsBP compared to the last glacial

maximum (LGM) (magnitude 2.44‰ for G. ruber).

The sediment influx data also corroborate well with the stable

isotope data on G. ruber. 18O isotope data on G. ruber reveals fluctuations in

the monsoonal regime from 6000 to 4300 years BP (-4.03 to -3.39 ‰).

Negative 18O value occurs at ~5500 years BP and the more negative 18O

around 4200 years BP (Figure 5.11 and 5.13) indicating weakened SWM

during this period. However around 4400 to 4200 years BP a signature of

weakened monsoon with cooler conditions is discerned by the values of 18O

moving from -3.39 to -2.33‰ (values becoming more positive). The lighter

values of 18O in the species G. ruber could be due to local factors such as

changes in sea surface salinity (SSS) by higher fluvial influx from the

Ganga-Brahmaputra system, and/or higher sea surface temperature (SST) in

the BOB. Further, there is a slight amelioration in the monsoons as revealed

by the 18O values ranging from -2.76 to -3.43‰ (~3425 years BP). Higher18O values of G. ruber and an increase in its abundance (Figure 5.11 and

5.13) can be strongly linked to an increase in SSS caused by a decrease in

riverine runoffs strongly associated with reduced SWM.

Thus an integration of the down core sediment texture, clay

mineralogy and stable isotope data on G. ruber points that the decrease in

sediment influx into the Bay of Bengal near the Landfall island reflects

weakened SWM and higher salinity conditions of the sea water during

(~ 6000 to 4400 years BP). An increasing occurrence of kaolinite and illite

indicate intense chemical alteration of island bed rocks from ~ 4400 to 1700

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years BP. The arid/dry episodes of 5000-4300 and ~ 2000 years BP largely

coincide with similar events reported from the northern Indian Ocean

(Von et al 1999, Chauhan et al 2000, Chauhan &Suneethi 2001).

Based on G. bulloidesforaminifer time series data from the Arabian

Sea, Gupta et al (2005) compared it with the sun spot data inferred from 14C

records as a proxy for solar activity including smoothed time series of 14C

production rates (Stuiver et al 1998) and showed that the summer monsoon

gradually weakened over the past 8000 years BP with a more or less stable

dry phase beginning since 5000 years BP. This coincided with the onset of an

arid phase in India (Sharma et al 2004) and termination of the Indus Valley

civilization (Staubwasser et al 2003, Gupta 2004). A high-resolution study of

mineralogy and major element geochemistry combined with Sr, Nd and

Oxygen isotopes ( 18O) was conducted by Colin et al (2006) and Colin et

al(1999) in two marine sediment cores collected off the Irrawaddy river

mouth (MD77-180) in the Bay of Bengal and in the Andaman Sea

(MD77-169). Pedogenic clays (smectite and kaolinite) to primary mineral

(feldspar, quartz, illite and chlorite) ratios showed strong precessional cycles,

suggesting a control due to the past changes in the summer monsoon intensity

(Colin et al 1999, Colin et al 2006). They also demonstrated wet periods of

summer monsoon reinforcement corresponding to an increase in weathering

of the Irrawaddy plain soils with a decrease in 87Sr/86Sr ratio.

Based on clay composition from the Arabian sea sediment cores,

Thamban et al (2002) suggested that the Early Holocene precipitation maxima

occurred after 9000 years BP and the late Holocene period witnessed reduced

rainfall activity and the resultant decrease in hydrolysis began around 5600

years BP. The clays were derived from the hinterland rocks and soils

(Thamban et al 2002). Stable oxygen and carbon analyses of marine sediment

cores from the eastern Arabian Sea, Sarkar et al (2000) found one arid period

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phase around 4000 years BP and two wet phase that occurred around 10, 000,

6000, 3500 and 2000 years BP. This enhanced precipitation was due to

intense SWM (Table5.4). Marine records also point two phases of aridity and

extremely weakened SWM around 4500 years BP (Sarkar et al 2000) and

reduced discharge of Indus during 2200-1800 years BP (Von Rad et al 1999).

However, the arid event with upwelling record during 3500 years BP of the

western Arabian Sea as reported by Naidu (1996) is not prominent in the Bay

of Bengal

5.6 PHYTOLITHS AND ITS INFERENCES

Phytoliths are siliceous remains, formed due to the absorption of

silica in the form of monosilic acid by the plants, which is precipitated within

the cells or intercellular regions of living plant tissues. These remains are a

redeposit in the form of phytolith (from Greek, phyto- plant, lith-stone i.e.

plant stone). Phytoliths occur in various sizes, shape and structure as it takes

the shapes of cellular bodies of the plant. The production and abundance of

phytolith is influenced by the concentration of monosilicic acid in the soils,

temperature, pH, water content, climate and environment (Jones &Handrek

1965, Piperno1988, Madella et al2002).

Apart from their morphological variability, being siliceous and

inorganic in nature the other strength of the phytoliths lies in their durability

as they are not susceptible to bacteria corrosion as are many other

microfossils. They are stable across a wide pH range (3-9) and are well

preserved in damp and desiccated as well as alternating wet and dry

conditions (Piperno 2006). Hence, unlike organic plant remains, they do not

depend on exceptional conditions for survival and are therefore widespread in

sediments and soils.

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As a consequence of their abundance, durability and diagnostic

morphologies, phytoliths have been increasingly used to reconstruct aspects

of Late Quaternary paleoenvironments in numerous sediment types, including

loess (Lü & Wang 1991, L et al 1996, Madella 1997, Blinnkov et al 2002),

lake muds (Carter 2002, Thorne 2004), sand dunes (Horrocks et al 2000,

Boyd 2005), tephra sequences (Sase et al 1987, Parr 1999, Lentfer et al 2001)

and coastal plain sequences in addition to other sediment types (Fredlund &

Tieszen1997, Carter & Lian 2000, Prebble & Shulmeister 2002, Lü et al 2002,

Abrantes 2003, Piperno & Jones 2003, Charles & Isabel, 2005).

According to Twiss et al (1969), Tieszen et al (1979) and Twiss

et al (2001) the diversity in the morphology of phytoliths indicate

palaeoclimatic conditions, as they are of C3 or C4 photosynthetic pattern.

Thus the phytoliths have proved to be important reliable indicators of

palaeoenviornment (Pearsal l2000, Barboni et al 1999, Ghosh et al 2008).

Application of phytolith study has been widely used in identifying the type of

vegetation specifically in elucidating the relative amount of grass cover as

opposed to forest cover, past vegetation, palaeoclimate allowing speculation

on past climate in the hinterland. Thus, phytolith analysis in dated

geo-archives are often used as an reliable tool for examining both changes in

paleoenvironmental/palaeoecological and also cultural records

(archaeological), including evidence of early agriculture, plant domestication,

early irrigation techniques, prehistoric stone tool function, pottery vessel

function, historic/prehistoric diet and food processing (Kajale & Eksambekar

2001, Madella 2003).

Phytolith assemblages have been used for reconstructing coastal

environmental changes and for validating the overwash sand layers in

palaeotempestology studies (Hou-Yuan Lu & Kam-biu Liu 2005). Phytolith

assemblages recovered from marine sediments off west equatorial and

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tropical Africa have been used as indexes of the continental aridity through

the climatic cycles (Diester-Haas et al 1973, Parmenter & Folger1974, Jansen

et al 1989). Although Palmer (1976) and Runge (1995) have considered the

potential of phytoliths for African vegetation reconstruction, however, no

paleoenvironmental study from phytolith assemblages have been made in

inter-tropical Africa. Opal phytoliths that are silicified casts of epidermal cells

of terrestrial grasses have been described and illustrated from Atlantic Ocean

sediment cores by Kolbe (1955) and Bukry (1979 a, b) and from

Mediterranean Sea cores by Dumitricà (1973) and Hajós (1973). No data

however, till date has been presented on the phytolith assemblages from the

Bay of Bengal. This is the first report of high resolution phytolith study from

a marine sediment core retrieved near the Landfall Island, Bay of Bengal

analyzed to understand the Middle Holocene SW monsoonal shifts.

Phytoliths are mainly transported by wind and the long shorecurrent system to the marine sedimentary environment (Wang et al1997,Wang 1997). Grasses are known to be the major accumulators of opalsilica/phytolith. In this study detailed phytolith assemblages down thesediment core indicate that the vegetation was stable around 4400 years BP (-24.3‰ 13C).The abundance festucoid indicate cool and dry climate whichwas suitable for the growth of small shrubs and trees of dicotyledonous typeas indicated by siliceous woody elements. This theory may be supported bythe lowering frequency of Diatoms and sponge spiclues.

5.6.1 Phytolith Assemblage

The various grasses of different subfamilies produces uniqueshapes, such as Panicoidea produces dumbbell (bilobate) and cross shapephytoliths which represents a warm moist climate. Similarly cold dry climatecan be revealed from the trapezoide and rondel phytoliths produced fromFestucoidea. While warm dry habitat was represented by the short-saddle

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shape formed from Chloridoideae (Twiss et al 1969, Piperno & Pearsall1998).

5.6.1.1 Climate index

Twiss (1987, 1992) has estimated the relative proportion of C3

grasses of the American Great plains. Climate index (Ic) has been extensively

utilized in the tropical countries to evaluate climatic changes (Barboni et al

1999, Bremond et al 2005, Masud Alam et al 2009). Ic is a ratio calculated

using the abundance of pooid (festucoid) versus the sum of pooid, chloridoid

and panicoid. High values of Ic point to a cool climate conditions which in

turn is considered to be in reduced precipitation or arid conditions.

Climate index (Ic) = (pooid / (pooid + chloridoid + panicoid)) × 100 (5.1)

5.6.1.2 Phytolith index

Calculation of phytolith Index (Iph) was proposed by Diester-Hass

et al (1973). The phytolith index (Iph) has been adopted by several workers

(Alexandre et al 1997, Wang et al 1999, Ghosh et al 2008, Prasad et al 2007

and Li et al 2010). A high Iph value indicates aridity or a reduced monsoon

precipitation. The phytolith index is expressed as a ratio of saddle phytolith

versus sum of saddle, dumbbell and cross phytoliths and it is represented as

follows:

Phytolith Index (Iph) = [Saddle / (Saddle + dumble +cross)] × 100 (5.2)

5.6.1.3 Cold / warm ratio

As the various types of phytoliths are produced under different

climatic condition, the ratio of cold - dry (trapezoid and roundel) to warm -

moist (bilobate, cross, and saddle) phytoliths indictes terrestrial climatic

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conditions and vegetation of the source area (Wang & Lu 1993, Lu et al

2002). Warm and humid conditions can be inferred from the lower values and

vice versa for the cold and dry conditions. In the present study warm and dry

condition is due to intense SWM and cold-dry condition is due to weakened

SW rains. The ratio can be calculated using the formula given below:

Cold / Warm (C/W) = Elongate/ (panicoid + Chloridoid) (5.3)

With reference to the variations in the phytolith counts, Ic, IpH and C/W

ratios were calculated down the sediment core profile and based on the data

generated four zones have been identified.

Zone IV (~6500 - 5000 years BP; Table 5.4): In this zone the

panicoid was abundant with slight increase in bulliform, woody elements and

spongy spicules than the other three zones whereas the Chloridoid, Festucoid

and Elongate presence are low. The calculated climate index (Ic), Iph and

C/W ratio are also low, which suggest that warm and humid climate prevailed

during this period indicating intense SWM rains. This corroborated with high

coarser components with mild increase in CaCO3 content, low organic matter

and clay percentage (Figure 5.15).

Zone III (5000 - 3600 years BP): This zone is in contrast to the

zone IV, zone III and is represented by low panicoid with higher chloridoid,

festucoid and trichome. It was also noted that the bulliform, spongy spicules

are low. The higher value of Ic, Iph, and cold/warm ratios point to a cold and

dry conditions that prevailed from 5000-3600 years BP. This observation is

corroborated with a rise in the sedimentation of finer components and organic

matter which is followed with reduced sand and high CaCO3 percentage

(Figure 5.15 and 5.16) indicating a calm and quieter environment.

Figure 5.15 Down sediment core variation of the sand, silt, clay, organic matter (OM), calcium carbonate (CaCO3)with the climate index (Ic), phytolith index (Iph) and cold/warm (C/W) ratio

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Table 5.4 Four important zones of climate variations classified from thephytolith assemblage

Zone Climate Remarks

Zone IV

(6500 - 5000 years BP)

(Intense SWM at 5600 years

BP)

Intense SWM

High panicoid, low

chloridoid, fetstucoid,

elongate, trichome, slight

increasing bulliform woody

elements and Spongy Spicules

with low Ic, Iph, Cold/Warm

ratio

Zone III

(5000 - 3600 years BP)

(Extreme weak SWM ~4400

years BP)

Weak SWM

Lowering panicoid,increasing

chloridoid, festucoid,

trichome, decreasing

bulliform, S.Spicules and

increasing Ic, Iph, Cold/Warm

ratio

Zone II

(3600 - 2400 years BP)

(strengthened SWM ~3400

years BP)

Fluctuation of

SWM

Lowering chloridoid,

festucoid, elongate, trichome,

increasing bulliform, woody

elements, and S. Spicules with

low Ic, Iph, Cold/Warm ratio

Zone I

(2400 - 800 years BP and

above )

(Weakening of SWM begins

from ~ 2400 years BP)

Weak SWM

Slight increase in chloridoid,

festucoid, elongate, lowering

trichome, bulliform, woody

elements, with increasing Ic,

Iph, Cold/Warm ratio

Zone II (3600 - 2400 years BP): In this zone II, panicoid

percentage was low with a slight decrease in chloridoid. Fluctuating high to

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low percentage of festucoid, elongate, and trichome was found along with

increasing bulliform, woody elements, and S. spicules. The low value of Ic,

fluctuating trend of Iph, Cold/Warm ratio reflects fluctuations in the intensity

of SWM. An increase both in sand and fine silt with less clay detritus in the

early part of the zone II followed by a decrease in sand and silt content

indicates fluctuating SWM.

Zone I (2400 - ~800 years BP): In zone I the occurrence of

chloridoid and panicoid is almost stable since 2400-~800 years BP with a

slight increase in festucoid phytoliths and diatoms. The woody elements and

bulliform phytoliths were found decreasing with increasing percentage of

elongate and spongy spicules. A slight increase in Ic, Iph values and

cold/warm ratio point to a cold and dry period due to weakened SWM. A

slight increase in silt and CaCO3 content indicates a dry environment due to

weakened monsoons (Table5.5).

5.6.1.4 Inference drawn from the results (sediment texture, CaCO3,

OM and Phytolith data)

Integration of the down core variations in sediment texture, CaCO3

percentage, and OM and phytolith data reveals an overall weakening of SWM

since mid Holocene period ~5000 years BP. Extremely weakened SWM

occurred around ~ 4400 years BP (-24.3‰ 13C)indicated by fairly stable

vegetation coverage due to the occurrence of high frequency of panicoid,

chloridoid and festucoid phytolith morph types. It is important to note that the

occurrence of other macrofossils such as diatoms and spongy spicules are

present in the entire profile in varying percentages.

Planktonic diatoms generally decrease with climate warming

because of reduced nutrient redistribution and increasing sinking velocities.

Climate warming exhibits a selection pressure on diatom cell size, this favors

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small-sized diatoms that are able to out compete larger sized cells and expand

under intensified stratification. This observation was made in this sediment

core recovered from the Bay of Bengal High frequency of diatoms is found in

lower zones as compared to the upper levels. This points also to a

hydrodynamic conditions that created a calmer environment, with strong

sedimentation of clay particles with abundant plant remains such as branches,

roots and leaves as indicated by phytolith morph types and silicified woody

elements, forming an organic clay sedimentary deposit, colonized by small

shrubs and vegetation. These environmental conditions probably contributed

to the retention of organic muddy sediments that favored the dominance of

diatom species. The presence of freshwater sponge spicules in sediments

obviously indicates an aquatic environment, but their gemmoscleres/

megasclare besides allowing for species identification, point to seasonal

changes in temperature and sea water column levels.

Siliceous sponge spicules provide a basis for palaeoecological

interpretation. Sponges require relatively clear, non-turbid water for living

conditions because muddy waters clog their pores (Eksambaker 2002). Thus it

appears that only some spicules were from in-situ sponges while the rest were

probably re-deposited. The low spicule abundance in the sediment core

resulted from local current conditions and the dilution effect through input of

terrestrial sediment. Other possible explanations for the varying spicule

abundance in the sediment core are due to the difference in local fauna, such

as coral reefs which usually have high diversity and abundance of sponges.

An integration of all the results show that the variations in the

sand/clay ratio influenced the ratio of the benthic/planktonic diatom species,

sponge spicules and phytolith and this probably was caused by local

hydrodynamic changes in SWM and also due to the fluctuations in the marine

water column, temperature and salinity. A comparison of paleoenvironmental

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inferences drawn based on phytolith study on the marine sediment core and

with the sediments from Itola site, Dhadhar river basin, Singh et al(2007)

suggested an occurrence of intense winter precipitation during early - mid

Holocene period (~3960years BP) and SWM from mid - late Holocene

period. Phytolith records from the Ganga plain revealed four dry phases 10,

300-9200, 5300-4100 years BP (most prominent), 1650-1200 years BP, and

950-700 years BP (Saxena et al 2013). Phytolith records from archeological

soils from Bangladesh (Masud Alam et al 2009) also point to a dry phase

from 730 to 1080 AD (i.e. ~1000 - 1300 years BP).

The present study using marine sediments and phytolith distribution

supplemented by radiocarbon ages and sediment texture corroborates well

with all the above mentioned climatic interpretations (Figure 5.15 and 5.16).

The data points to an overall weakening of SWM from 5000 years BP to 1200

AD. The Early middle Holocene period (~6500 years BP to 5000 years BP)

experienced strong SWM that eventually tapered. The period from ~5000-

3600 years BP witnessed weakening of SWM rains, followed by fluctuations

in its intensity from 3600-2400 years BP. Further the SWM conditions did not

improve but continued to weaken from 2400 years BP to 800 years BP.

However, much variability exists in the continental records about

the prevalent climate of the Himalayas during Mid Holocene period. In the

Gangotri area, the Shivling Stage of glaciations has been reported around

5100 years BP (Sharma &Owen 1996) and the pollen records from the Dokini

Glacier- periglacial peat indicate prevalence of arid climate during 4000-3500

years BP (Phadtare 2000). In our records this phase terminates at around 4300

years BP. Phytolith records from Deoria Tal (2770 masl; Uttarkashi),Sharma

et al (1995) reported an occurrence of humid perid around 4000 years BP and

a persistence of drier climate between 3200 and 1700 years BP.

Figure 5.16 The down sediment core variation of the phytolith assemblage, diatoms, spongy spicules with the climateindex (Ic), phytolith index (Iph) and cold/warm (C/W) ratio.

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5.7 NANOPLANKTON

Planktonic community was one of the wide varieties of organisms

that contribute the base of the marine food web and drive the biogeochemical

cycles of carbon and nutrients (Smith et al 2009). Ammonia is considered to

be the main nitrogen source for the primary production of plankton

community (Jochem 1989).

The planktons are capable to adapt in the physical environment

(salinity, temperature and nutrient concentration) and morphology to live in

the water of ocean (Jochem 1989). The planktons are classified mostly

according to their size of the organisms, such as picooplankton, nannoplanton

(2 - 20 m), microplankton (20-200 m) etc. (Dussart 1965, Sieburth et al

1978). Nanno and picoplankton are collectively termed as ultraplankton

(Murphy and Haugen 1985).

Picoplanktons (< 2 m) and nannoplankton (2 m to 20 m) can

also contribute downward carbon flux (Gorsky et al 1999), other than diatoms

and the dinoflagellates (two chief component of phytoplankton) (Severdrup et

al 1942), Hence the calcareous nannoplankton is also used as one of the most

important climate indicator among other proxy such as foraminifera,

alkenones, dinoflagellate cysts, especially in Mediterranean sea (Antonioli et

al 2001, Triantaphyllou 2009).

The study of nannofossil distribution and its application in

reconstructing palaeoclimate is rising in Indian Ocean also (Agnihotri &

Kurian 2008). The nannoplankton assemblage in the studied marine sediment

core (0-124 cm) reveals the abundance of Emiliania huxleyi and

Gephyrocapsa Oceanica rare of Pleistocene to recent age (Bukry 1971). It is

also noted that the presence of some important species of nannoplanktons of

Pliocene (Discoaster brouweri) and Miocene (Coccolithus pelagicus) suggests

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that the sediments are derived from the reworked sediment of Mio-Pliocene

age (Table 5.5).

High content of sand with reduced silt, OM, and nannoplankton

productivity ~ 6300 - 6000 years BP, 5600 - 5200 years BP suggest intense

SWM. It is also observed that ~ 3600 - 3300 years BP the sand and silt

content was higher with reduced nannoplankton and clay reveals high river

influx due to the SWM. In contrast two major phase of weakening SWM was

found around 4200-3800 years BP and 2000-1800 years BP inferred from the

reduced sand with increasing percentage of silt, OM and nannoplankton

productivity (Figure 5.17). It was in agreement with Jochem (1989), that the

primary production of nannoplankton was lower in winter monsoon whereas

it is higher during the summer season as a dominating size fraction which can,

contribute 70 to 100% of primary production compared to the other plankton

(Pico and microplankton)

The nannoplankton data from the BOB sediment core revealed

fluctuations in its abundance for the past 5680 years BP. This variation is

associated with the fluctuations in the nutrient enriched /depleted water and

changes in the thermocline/nutricline depth. A significant increase in the

surface water productivity during the deglaciation period is evident from the

phytolith proxy data.

The overall correlation between ascidian spicule abundance and the

relative abundances of upper photic zone calcareous nannoplankton reflect a

close connection between the enrichment in nutrients in the upper photic zone

and an increment in bottom productivity. This supports the idea of benthic

communities in this area strongly benefiting from surface water productivity

(bentho-pelagic coupling).

128

Table 5.5 Abundance and the age determination of the species after(Bukry 1971, Gartner 1971).

Age Species Abundance

Pleistocene to recent

Emiliania huxleyi Fai and CommonGephyrocapsa Oceanica AbundantGephyrocapsa Caribbeanica Fair and RarePseudoemiliania lacunosa Rare

Pliocene

Discoaster brouweri(Middle Pliocene) Rare

Reticulofenestra pseudoumbilica RareSphenolithus abies( Early Pliocene) Rare

Miocene

Discoaster quinqueramus RareCoccolithus pelagicus(Late Miocene) Fair and Rare

Discoaster calcaris (Late Middle Miocene) Rare

5.8 RECONSTRUCTION OF PAST RAINFALL

The moisture during summer mainly originates from the equatorial

Indian Ocean (Laskar et al 2013b). There is no evidence of any significant

change in the SWM moisture source during the mid to late Holocene (Laskar

et al 2013b). However, a latitudinal shift of the Inter tropical Convergence

Zone could be responsible for changing the strength of the SWM (Fleitmann

et al 2007, Laskar et al 2013a, b and references therein). However, in a small

island station, change in monsoon seasonality for the last 6000 years BP is not

expected. Therefore, the main governing factor is plausibly the amount effect

during the monsoon season.

The basis of palaeomonsoon reconstruction in the tropics is the amount

effect (Yadava & Ramesh 2005). They observed significant amount effect in

the plains of central India. However, exceptions were also observed in the

tropical to subtropical SWM rainfalls in hilly terrains of NE India

Figure 5.17 Downcore variations of Sand, Silt, Clay Organic matter (OM) percentage and nannoplankton countsreveals phases intense and reduced SWM.

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(Breitenbach et al 2010, Laskar et al 2013a) and some parts of Southern India(Yadava et al 2007). Variations in the amount of past rainfall have been due

to factors, such as changes in the 18O of the ocean waters due to evaporation(‘‘ice volume’’effect), shifts in the source of moisture (Bar-Matthews et al

1999, Fleitmann et al 2003, 2007), and changes in the seasonality ofprecipitation (e.g., change in the proportion of winter and summer

precipitation). It has been noted that in the last 4000 years BP, there is nosignificant change in the ocean 18O, that can cause change in rainwater and

hence sedimentation.

In the present study, based on the data collected a strong reduction

in the monsoon during 4400- 4000 and 2200 - 1800 years BP is observed. Theperiod 2200 – 1800 years BP coincides with the RWP, previously thought to

be restricted to higher latitude regions (Vollweiler et al 2006, Desperat et al2003, Martinez-Cortizas et al 1999, Laskar et al 2013a, b). The BOB sediment

core data indicate that the effect of RWP is evident in the tropics also.Weakening of the monsoon has also been observed around 1500 and 400-800

years BP, the latter period is the transition from MWP to LIA. A significantreduction in the SWM during this period was also reported by several others

(Fleitmann et al 2004, Sinha et al 2007). The SWM was significantly strongerduring 800-1200 years BP, which is the MWP. Enhanced monsoon during

MWP was also observed in Oman (Fleitmann et al 2004). The discrepanciesin the timing of these events recorded at the different locations are probably

the result of age uncertainty of 100-200 years BP in the present chronology.

5.9 COMPARISON WITH OTHER PALAEOMONSOONRECORDS

Influences of the global climate instabilities, such as Roman Warm

Period (RWP), Medieval Warm Period (MWP) and Little Ice Age (LIA) areevident in the SWM precipitation reconstructed here. To check whether our

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marine sediment core based chronology reflects a local or broader regionalrainfall variability, we compared the 18O record (G. ruber) with other marine

and continental climate proxy records from some of the areas receiving SWM(Chauhan et al 2004, Chauhan & Vogelsang 2006, Chauhan & Sunethi 2001,

Mathien & Bassinot 2008, Rashid et al 2007, Chauhan et al 2000, Sharma etal 2006, Saxena et al 2013, Masud Alam et al 2009, Singh et al 2007, Enzel

et al 1999, LuÈckge et al 2001, Sarkar et al 2000, Thamban et al 2002,Patnaik et al 2012, Kupusamy & Ghosh 2012, Thamban et al 2007,

Laskar et al 2013a,b). A comparison with some of the well studied sites ispresented in Figure 5.18. By and large, variations observed in the marine

sediment record are similar to the other proxy records, except with somemismatches and sharp changes that may be attributed to uncertainty in the

chronologies and relative differences in sampling resolution. The mostcommon feature in all these records is the reduction in the SWM intensity

during the RWP and MWP.

Climate fluctuation during the RWP affected the SWM

significantly. Reduction in the SWM during the RWP is evident in the

abundances of Globigerina bulloides (Gupta et al 2003) and 18O of G. ruber

and G. Sacculifer (Tiwari et al 2006) in sediment cores from Arabian Sea. A

significant decrease in monsoon strength is visible in all the studies the period

between 800-400 years BP, i.e. the transition period from the MWP to the

LIA. A relatively dry episode at the beginning and a wet phase at the end of

the LIA are also reported from the Himalayan region (Kotlia et al 2012). The

monsoon was significantly stronger as compared to the present during

1200-800 cal years BP, the MWP. Evidence for a stronger monsoon during

this interval is also found in records of varve thickness (Von Rad et al 1999)

and stalagmite 18O from the Indian subcontinent (Sinha et al 2007). The

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SWM was similar to the present without any major change in the source,

extent, and seasonality before 6000 years BP except two short-term arid

phases around 4400 - 4000 years BP and 2000 - 1700 years BP (Sinha et al

2007).

There has been much discussion about the magnitude and

geographical extent of the Medieval Warm Period and Little Ice Age. Soon et

al (2003) reviewed a wealth of palaeoclimatic data from around the world.

Their study indicate a relatively warmer and cooler epochs that were of global

extent and that the MWP was at least as warm as and in many places and

times, even warmer than - the current MWP.

Andersson et al (2003)reconstructed surface temperatures of the

eastern Norwegian Sea based on the planktic stable isotopes and foraminiferal

assemblages. The climate history derived from their study is remarkably

similar to that derived by McDermott et al (2001) from a high-resolution 18O

record obtained from a stalagmite discovered in a cave in southwestern

Ireland. McDermott et al (2001) data shows that at the beginning of the 3,

000 year long Norwegian Sea record, both regions were in the end stage of

the long cold period that preceded the RWP. Hence, both the records depict

warming from that point in time to the peak of the RWP, which occurred

about 2000 years BP.

Berglund (2003) identified several periods of expansion and decline

of human cultures in Northwest Europe and compared them with a history of

reconstructed climate based on insolation, glacier activity, lake and sea levels,

bog growth, tree line, and tree growth. Starting from the climatic warmth and

elevated human productivity of the RWP, there was a retreat of agriculture

about 1500 years BP and this was coincident with a period of rapid cooling -

which ushered in to the Dark Ages (DA) Cold Period. This was also indicated

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in the tree-ring study data (Eronen et al 1999) as well as sea surface

temperatures based on diatom stratigraphy in the Norwegian Sea (Jansen and

Koc 2000). This period correlated well with Bond's event 1 in the North

Atlantic sediments (Bond et al 1997) (Figure 5.18).

Several centuries from 700 to 1100 AD interval of time proved to

be a favourable period for agriculture in marginal areas of Northwest Europe,

leading into the so-called Medieval Warm Epoch when the climate was warm

and dry, with high tree lines, glacier retreat, and reduced lake catchment

erosion (Berglund 2003). This lasted until around 1200 AD, when there was a

gradual change to cool/moist climate, the beginning of the LIA with severe

consequences for the agrarian society.Castagnoli et al (2002) developed a

1400-year record of 13C values from remains of the foraminifera G.ruber,

which were obtained from a sediment core retrieved from the Mediterranean

Sea. Their data revealed an initial increase in 13C values that coincided with

the climatic transition from the DA to MWP, over which marine productivity

in the Mediterranean rose dramatically.

Ma et al (2003) assessed the climatic history around the Northern

Hemisphere of the past 3000 years at 100-year intervals on the basis of 18O

data, the Mg/Sr ratio, and the solid-liquid distribution coefficient of Mg.They

derived the data by careful analysis of a stalagmite from Jingdong Cave about

90 km northeast of Beijing, China. Ma et al (2003) reported that the air

temperature between 200 and 500 years BP was about 1.2°C lower than that

of the present, corresponding to the LIA in Europe. Between 1000 and 1300

years BP, there was an equally aberrant but warm period that peaked at about

1100 years BP, and this they corresponded to the MWP (~900-1300 AD) in

Europe.

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Xu et al (2002) studied plant cellulose 18O variations in sediment

cores retrieved from peat deposits at the northeastern edge of the Qinghai-

Tibetan Plateau in China. Following the decline of the RWP, their data

revealed the existence of three particularly cold intervals centered at

approximately 500, 700 and 900 AD during the Dark Ages (cold period).

Then, from1100-1300 AD, the 18O of Hongyuan peat cellulose increased,

consistent with that of Jinchuan peat cellulose and corresponding to the MWP

(Xu et al 2002). Finally, Xu et al (2002)observed that the periods: 1370-1400

AD, 1550-1610 AD, and 1780-1880 AD recorded three cold events,

corresponding to the LIA. Xu et al (2002) detected the relative warmth and

coolness of the RWP, DA, MWP and LIA.

Paulsen et al (2003) used high-resolution records of 13C and 18O

analysed from a stalagmite taken from Buddha Cave to infer changes in

climate in central China for the last 1270 years.They categorised the climate

in terms of warmer, colder, wetter and drier conditions.Among the climatic

episodes revealed in their data, they specifically identified those

corresponding to the MWP, LIA and 20th-century warming, lending support

to the global extent of these events. In addition, their record begins in the

depths of the DA, which ends about 965 AD with the commencement of the

MWP, that continues to approximately 1475 AD, whereupon the LIA sets in

and holds sway until about 1825 AD, after which the warming responsible for

the modern warm period begins.

Yang et al (2002) used nine separate proxy climate records, derived

from peat, lake sediment, ice core, tree ring and other proxy sources, to

compile a single weighted temperature history for China spanning the past

two millennia. Their composite record revealed five distinct climate epochs:

a warm stage from 0 to 240 AD (the tail-end of the RWP), a cold interval

between 240 and 800 AD (the DA), a return to warm conditions from 800-

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1400 AD (which included the MWP between 800 and 1100 AD), a cool

interval between 1400 and 1820 (LIA), and the current warm regime (MWP)

that followed the increase in temperature that began in the early 1800s.

Another important finding of this study was that the warmest temperatures of

the past two millennia were observed during the second and third centuries

AD near the end of the RWP.

Kar et al (2002) explored the nature of palaeoclimate change over

the past 2000 years using pollen data analysed from a 1.25-meter sediment

profile. The sediement core was collected from an outwash plain located

about 2.5-3 km from the Gangotri glacier's snout. Their data revealed the

existence of a cooler climate than the one prevailing at present between 2000

and 1700 years BP. Comparing this result with the findings of McDermott et

al (2001) in Ireland, this period of time is seen to be part of the DA. Between

1700 and 850 years BP, there then occurs what Kar et al (2002) call an

amelioration of climate, which represents the transition from the depth of the

DA to the midst of the MWP. Subsequent to 850 years BP, the climate

became much cooler, indicative of its transition to LIA conditions. Between

300 and 200 years ago, Kar et al (2002) observed the ceasation of the long-

term retreat of the Gangotri Glacier, possibly with some minor advancement.

During the last 200 years, however, the study of Esper et al (2002) indicates

there has been a rather steady warming of the Northern Hemisphere - the

Gangotri glacier's snout has retreated by about 2 km.

The results of the present study based on the sediment core

collected from the BOB near the Landfall Island clearly demonstrates that

there exists harmony of climate change in the region as observed in North

Atlantic Ocean, distant Himalayas and East Asia providing evidence for

RWP, MWP and LIA.There are a number of conclusions that may be drawn

from the present study.

Figure 5.18 A comparison of monsoonal shifts of the present study with global climate events, Bond events (shaded) and otherpalaeoclimate records, reviews from Arabian sea, Bay of Bengal, Andaman sea and Indian sub-continent.[Reference. 1)Chauhan et al (2004), 2) Chauhan and Vogelsang (2006), 3) Chauhan and Sunethi (2001), 4) Mathienand Bassinot (2008), 5)Chauhan et al (2000), 6) Rashid et al (2007), 7) Laskar et al (2013), 8) LuÈckge et al (2001),9) Sarkar et al (2000), 10) Thamban et al (2002), 11) Sharma et al (2006), 12) Saxena et al (2013),13) Masud Alam etal (2009), 14) Singh et al (2007), 15), Enzel et al (1999), 16) Patnaik et al (2012), 17) Kupusamy and Ghosh (2012),18) Thamban et al (2007) and 19,20,21 our present data of phytolith, clay mineralogy and Geochemistry].

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5.10 CONCLUSIONS

A marine sediment core of 124 cm retrieved at a depth of 250 m

water column, near Landfall Island, North Andaman, Bay of Bengal,was

studied using various proxies such as clay mineralogy,oxygen ( 18O) and

carbon ( 13C) stable isotope, neodymium ( Nd), geochemistry, phytolith and

nannoplankton assemblage. Based on the data collected and integration of

results following conclusions are drawn:

i) The sediment texture is predominantly represented by clayey

silt. The sediment core reveals layers of coarser sand flux

since ~ 6500 to 6000 and ~ 3300 years BP, that reflects a

strengthened SWM in an overall weak SWM of the middle to

late Holocene.

ii) The study of geochemical, grain size and isotopic data of these

sediments show alternate dry/warm and wet (humid)

conditions prevailing in the southern tropical areas since

~6500 years BP with weaker southwest monsoon during 4300

to 4000 years BP, strengthened SWM conditions ~3300 years

BP and an amelioration in climate since ~ 2000 years BP. Due

to the intense SWM, the short wet conditions resulted in

intense chemical weathering of the Island rocks. The arid

event with upwelling record during 3500 years BP of the

western Arabian Sea as reported by Naidu (1996) is not

prominent in the Bay of Bengal

iii) The geochemistry, neodymium and clay mineral assemblage

suggests that the sediments deposited in this region have been

contributed by two major provenances. A nearby source i.e.

from Andaman Sediments predominately derives from the

Andaman terrain of ophiolite suites, followed by sediments

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from Irrawaddy, Salween and Sittang rivers from Myanmar.

The transport of weathering products from these land masses

to the AFB site allows us to explain the chemical, isotopic and

grain size signals observed in the AFB core from Mid-

Holocene to late Holocene.

iv) Smectite and illite are the two dominant clay minerals in this

sediment core. High content of smectite reflects the

weathering product of mafic rock contributed from the nearby

island whereas illite is predominantly coming from the rivers

such as Irrawaddy and Salween.

v) An overall weakening of SWM is observed from ~6000 to

2000 years BP, and this t is inferred from the high smectite,

high C/I ratio and low K/C ratio. Within this period; an

incursion of a more intense weakening of SWM is noted from

the elevated C/I ratio and lesser K/C ratio from 4400 to 4200

years BP, which is also supported by the 18O data of G. ruber

(-3.39 to -2.33‰). Occurrence of high kaolinite percentage

and K/C ratio points to strong weathering during 3900 to 3600

years BP and subsequently an arid phase is observed between

20-15cm (~2000-1700 years BP). A major wet phase also

found reaching its maximum at 3400-3200 years BP and

amelioration in climate since ~ 2000 years BP to present. The

intensity of SWM tends to repeat itself.

vi) The phytolith analysis shows a significant variation in the

SWM from ~ 6500 to present, with weakening of the SWM

from 5000 years BP. The reduced Ic, Iph and C/W ratio with

increased panicoid, sand content during the earlier mid -

Holocene period (~ 6500 to 5000 years BP) suggest an intense

SWM. In contrast an elevated value of Ic, Iph and C/W ratio

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with high clay component around 4400 years BP (-24.3‰13C) points to a weakened SWM regime. A fluctuation in the

SWM was also observed from 3600 to 2400 years BP. This is

inferred from low Ic value with fluctuating Iph and C/W ratio.

Amelioration in SWM is noticed around 3400 years BP

pointed out by higher coarser component with low clay

content followed by a low Ic value. A slight increasing trend

of Ic, Iph and C/W ratio suggest a weakened SWM that

persisted since 2400 to 800 years BP.

vii) An integration of all the results show that the variations in the

sand/clay ratio influenced the ratio of the benthic/planktonic

diatom species, sponge spicules and phytolith and this

probably was caused by local hydrodynamic changes in SWM

and also due to the fluctuations in the marine water column,

temperature and salinity.

viii) An integeration of various palaeoclimate proxies of the marine

sediment core of the present study with review and global

climate events reflects the significant signatures of intense and

weakeaned SWM. Such as wet phase during 6500 to 5600

years BP (HCO), 3400 - 3200 years BP (SBO) and warm dry

phase ~ 2200 to 1800 years BP (RWP) and 1000 to 800 years

BP (MWP) with an extreme weakened monsoon ~ 4400 -

4000 years BP (Sub - Boreal Optimum).

5.11 SCOPE FOR FUTURE STUDY

The present study deals with a proper understanding and

reconstruction of palaeo monsoonal (SWM) shifts over the Bay of Bengal and

Indian subcontinent during the middle to late Holocene period. This was

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carried out using a marine sediment core off Landfall Island, North Andaman

Bay of Bengal. A multiproxystudy using clay mineralogy, oxygen ( 18O) and

carbon ( 13C)stable isotope, neodymium ( Nd), geochemistry and phytolith

assemblage was attempted.This study presents an preliminary endeavour

using multi and interdisciplinary approach to reconstruct palaeo southwest

monsoon The importance of this study lies in the fact that in the overall

weakening of SWM fgrom 6000 years BP to presentthe major dry event

occurred around 4200-4000, 2200 - 1800 years Bp (RWP), 1000 - 800 years

BP (MWP) and a small wet phase ~ 3300 years BP events are revealed in the

sediments. However, there are still number of the significant gaps that need to

be filled,for instance a detailed record of C/N ratio of biogenic sediments,

geomagnetic, palynology,radiogenic isotope, Cd/Ca ratio in foraminiferal

tests, 15N, Radiolaria and Dinoflagellate upwelling index.

Reconstruction of the the Indian summer mosoon over the past

several thousands of years is still in an infancy stage. Hence high resolution

studies need to be carried out using marine sediment archives along the

eastern and western margin of India, Gulf of Bengal and deep sea (BOB).

Serious attention needs to be paid for building to be paid towards

accuratechronostratigraphy (age control)from a long (> 40 to 50 m),

undisturbed, continuous, and large diameter sediment cores. Application of

Ce, Sr, Pb and S geochemical data would help in understanding source of

dusts and aerosols other than the bed rock weathering and suspended

particles. Numerical data generated from marine sediment cores would help

in the preparation of climate models with the help of REGCM3, GCM and

AMMA for the prediction of future climate.