indonesian throughflow and monsoon activity records in the timor sea since the last glacial maximum
TRANSCRIPT
Marine Micropaleontology 101 (2013) 115–126
Contents lists available at SciVerse ScienceDirect
Marine Micropaleontology
j ourna l homepage: www.e lsev ie r .com/ locate /marmic ro
Indonesian Throughflow and monsoon activity records in the Timor Sea since the lastglacial maximum
X. Ding a,⁎, F. Bassinot b, F. Guichard b, N.Q. Fang a
a China University of Geosciences (Beijing), Beijing 100083, Chinab Laboratoire des Sciences du Climat et de l'Environnement Domaine du CNRS, Gif-sur-Yvette 91198, France
⁎ Corresponding author. Tel. +861082334643.E-mail address: [email protected] (X. Ding).
0377-8398/$ – see front matter © 2013 Elsevier B.V. Allhttp://dx.doi.org/10.1016/j.marmicro.2013.02.003
a b s t r a c t
a r t i c l e i n f oArticle history:Received 26 February 2012Received in revised form 7 February 2013Accepted 10 February 2013
Keywords:Timor SeaLGMHolocenePlanktonic foraminiferaIndonesian ThroughflowMonsoonSea level changes
Indonesian Throughflow (ITF) is known to play an important role in the heat exchange between the Pacificand the Indian Oceans. However, our understanding of the long-term evolution of the ITF and, in particular,the mechanism of heat transport is limited. Here, we present a high-resolution foraminifera-basedmulti-proxy study in the main ITF outflow area of the Timor Sea, to reconstruct the ITF variability and to un-derstand the relationship between the ITF changes and monsoon activity from the last glacial maximum(LGM) to the Holocene. Our results show that when the strong surface water ITF occurs, high productivityis related to the mixing of the upper water column owing to the wind-driven upwelling rather than theshoaling of the depth of thermocline (DOT). By contrast, the DOT is affected more strongly by the ITF thanby the monsoonal wind-driven upwelling in the Indonesian Seas. During the LGM (23–19 ka) and middle Ho-locene (8–6 ka), warm surface water ITF was dominated owing to the lowered sea level and (or) the highersteric height difference between the western Pacific and eastern Indian Oceans as a result of the strong south-east monsoon. During the early Holocene (11–8 ka) and late Holocene (last ~6 ka), because of the postglacialhigh sea level, the strong northwest monsoon and heavy rains, large amounts of freshwater flowed into theJava Sea from the South China Sea (SCS). The freshwater plug at the southern tip of the Makassar Straitblocked the warm surface flow, thus initiating the enhanced thermocline ITF. In the Timor Sea, the changesin the vertical profile of the ITF were influenced by the glacio-eustatic sea-level changes that have modifiedthe geometry of the pathways within the Indonesian Seas, as well as by the monsoon activity which wasmodulated by the changes in the insolation with a precessional cyclicity.
© 2013 Elsevier B.V. All rights reserved.
1. Introduction
Today, the wind stress between the Pacific Ocean and the IndianOcean maintains a sea-level height difference between these twooceans (Bray et al., 1996), as seen in the Indonesian Throughflow(ITF), a network of currents, surface and thermocline waters thatare transported from the western equatorial Pacific Ocean into the In-dian Ocean (Hirst and Godfrey, 1993; Linsley et al., 2010). The ITF isthe only low-latitude connection along the return branch of theGreat Conveyor Belt, which ultimately brings upper thermocline andsurface waters from the Pacific to the North Atlantic (Gordon, 1986;Hirst and Godfrey, 1993; Bray et al., 1996; Gordon and Fine, 1996;Müller and Opdyke, 2000). The annual mean heat transport throughthe Indonesian Throughflow region (about 1.4×1015 W) representsa heat sink for the upper Pacific Ocean and is an important heatsource for the Indian Ocean (Schiller et al., 1998; Ganachaud andWunsch, 2000). However, modeling experiments and recent oceano-graphic measurements indicate that the modern ITF transport occurs
rights reserved.
mainly within the thermocline rather than at the sea surface (Gordonet al., 2003; Potemra et al., 2003; Song and Gordon, 2004; Gordon,2005), and thus the net effect in terms of heat transport to the IndianOcean could be negative instead of being positive.
The major component of the ITF is the Mindanao Current thatoriginates from the upper thermocline of the North Pacific and istransported into the Indonesian Seas through the Makassar Strait(Gordon, 1986; Murray and Arief, 1988; Gordon and Fine, 1996)(Fig. 1). Within the Makassar Strait, the 680-m-deep Dewakang sillpermits only the upper thermocline waters to enter the Flores Seaand flow eastward to the Banda Sea, or to directly exit into the IndianOcean via the shallow Lombok Strait (Sprintall et al., 2009). Today,the ITF transport of warm, low-salinity water into the Indian Oceanaverages ~16 Sv (1 Sv=106 m3 s−1) per year (Gordon and Fine,1996; Schiller et al., 1998; Gordon et al., 2003). Mooring measure-ments show that only a small portion (~1.7 Sv) (Murray and Arief,1988) of the waters flowing through the Makassar Strait across theIndonesian Seas directly enters the Indian Ocean via the LombokStrait (with a sill depth of 350 m) between the islands of Bali andLombok (Fig. 1). The largest part of these waters turns eastwardinto the Banda Sea and Flores Sea before spreading into the Indian
Fig. 1. Locations of core sites MD98-2172, SHI9034 and SHI9022 and those referred to in the text, with main oceanographic surface current (black line arrows), thermocline current(black dashed arrows) and main geographic locations mentioned in the text. Core SHI9016 was analyzed by Spooner et al. (2005) and core MD01-2378 by Xu et al. (2006).
116 X. Ding et al. / Marine Micropaleontology 101 (2013) 115–126
Ocean through two main pathways: the Ombai Strait (with a sill depthof 3250 m and a yearly average flow of ~5.0±1 Sv (Molcard et al.,2001)) and the Timor Sea (with a sill depth of 1890 m and a yearly av-erage flow of 7.0 Sv (Creswell et al., 1993)) (Fig. 1). These ITF watersflow into the Indian Ocean as parts of the west-flowing South Java Cur-rent, the South Equatorial Current, and the south-flowing Leeuwin Cur-rent that runs along the western Australian margin (Gordon and Fine,1996; Siedler et al., 2001) (Fig. 1).
Themodern climate of the Indonesian Seas is dominated by bi-annualmonsoonal shifts. Heavy rain accompanies northwesterly winds betweenNovember and March (Austral summer), during the Northwest (NW)Monsoon. The dry season corresponds to the Southeast (SE)Monsoon pe-riod fromMay to September (Austral winter) (Spooner et al., 2005).
The intertropical convergence zone (ITCZ), a pressure troughwhere the southeast and northeast trade winds meet, usually liesabout 10°–15° north of the equator in the Austral winter and migratessouthward, close to or over northern Australia in the Austral summer(Spooner et al., 2005). During the Austral summer, the NW Monsoongathers large amounts of moisture while crossing the sea from theAsian high-pressure belt on itsway to the ITCZ,which has shifted south-ward. At the ITCZ, the moisture-laden air rises, resulting in heavy rains(van der Kaars et al., 2000). During the Austral winter, the SE Monsoonoriginates from the Southern Hemisphere high-pressure belt and is rel-atively dry and cool (van der Kaars et al., 2000; Spooner et al., 2005).
The modern ITF is also closely related to the Asian monsoon dy-namics. The main flow of the ITF in the key passages of the Makassarand Timor Straits shows a strong seasonal variability (Gordon et al.,1999; Potemra et al., 2003). During the NW Monsoon (Austral sum-mer), a thermocline flow of relatively cool water dominates, as thewarm surface flow becomes blocked by the development of a fresh-water plug at the southern tip of the Makassar Strait, driven by mon-soonal winds from the Java Sea (Gordon et al., 2003; Gordon, 2005).As a result, the tropical Indian Ocean is cooled rather than warmedby the ITF (Song and Gordon, 2004).
On the Milankovitch time scale, two main mechanisms may affectthe ITF: (1) orbitally driven, low-latitude changes in insolation thataffect monsoon dynamics and (2) glacio-eustatic sea-level changesthat modify the geometry of the pathways within the IndonesianSeas. We have very limited understanding of the ITF evolution atthis time scale. Using the records of the δ18O and Mg/Ca of planktonic
foraminifera, Linsley et al. (2010) suggested that the freshening of thesurface ocean in the southern Makassar Strait 9.5 ka ago increasedthe northward pressure gradient and inhibited the flow of warmersurface-layer water into the Indian Ocean. Thus, 9.5 ka may havemarked the initiation of the thermocline-enhanced cool ITF transportthat is observed today. Xu et al. (2006, 2008) used a multi-proxy (plank-tonic foraminiferal census data and oxygen isotopic and Mg/Ca records)approach to reconstruct changes in the vertical profile of the ITF (depthof the thermocline and changes in sea surface and upper thermoclinetemperature), as well as monsoonal wind and precipitation variations inthe Timor Sea during Terminations I and II. These studies indicated thatthe vertical structure of the ITF probably varied considerably over preces-sional and glacial–interglacial time scales, with the thermocline flowdominating during warm periods, i.e. the thermocline shoaled from thelast glacial maximum (LGM) to the Holocene. However, through recon-structions of the vertical structure of the water column in the Banda Seaover the last ~80 ka using the abundance ratio of the planktonic forami-nifera thermocline and mixed-layer dwellers, Spooner et al. (2005)showed that the mixed layer was thinner during the LGM, but thickenedat the beginning of the Holocene. The latter may have been related to thestrengthened surfacewater ITF in the early Holocene. Owing to the differ-ent climate proxies used in these studies, the existing data appear to beinconsistent.
A multi-proxy study in the main ITF outflow area and Java upwell-ing area is presented here, with the goal of comparing the resultsobtained using different climate proxies and reconstructing the ITFvariations, hence improving our understanding on the relationshipbetween the ITF changes and monsoon activity from the LGM to theHolocene. Our main objectives are to track changes of the ITF outflowduring a period of extreme climate change and sea-level variationsand to assess links between high- and low-latitude climate changes.
2. Materials and methods
Core MD98-2172 (8°31′S, 128°09′E) was obtained from the TimorSea at a water depth of 1768 m during the International Marine Glob-al Change Study (IMAGES) cruise IV of the R/V Marion Dufresne(Fig. 1). This Calypso giant piston core is 54 m long, but only workcarried out on the upper 7.5 m of the core is discussed in this paper.Core MD98-2172 was sampled at 2 to 10-cm intervals for stable
Table 1Radiocarbon dates from cores MD-982172, SHI9022 and SHI9034. After adjusted bysubtracting 400 years for the marine reservoir effect, 14C dates were calibrated usingthe radiocarbon calibration program Fairbanks0701.
Lab code Depth 14C dates 14C cal(cm) (year BP) (year BP)
MD98-2172GifA 101513 0–2 1330±80 845±87GifA 101512 350–351 9120±100 9702±164GifA 101511 418–419 11,310±130 12,803±109GifA 101510 465–466 13,270±290 14,973±382GifA 101509 511–512 15,210±180 17,726±375GifA 101508 581–582 18,250±190 21,159±380
SHI9022GifA 95334 84–85 3800±70 3644±89GifA 95335 154–155 5580±70 5930±74GifA 95336 204–205.5 6880±80 7397±69GifA 95339 385–386 9800±90 10,630±131GifA 95340 474.5–475.5 14,850±130 17,053±255GifA 95341 504.5–505.5 15,110±130 17,529±288GifA 95342 594.2–595.8 18,620±160 21,713±267GifA 95343 694.2–695.8 26,290±300 31,112±352
SHI9034GifA 9504 34–36 3110±60 2802±53GifA 9505 44–46 3930±60 3813±82GifA 9506 105–106 5960±70 6340±62GifA 9507 125–126 6610±60 7122±87GifA 9508 360–362 12,360±90 13,780±83GifA 9510 390–391 12,760±90 14,209±78GifA 9511 430–431 14,350±100 16,267±177
117X. Ding et al. / Marine Micropaleontology 101 (2013) 115–126
isotope analysis of foraminifera and at 10-cm intervals for foraminifercensus, carbonate content measurements, and grain size analysis.
After being dried and weighed, samples were washed through a150-μm sieve. The coarse fraction was dried at 40 °C overnight, andthen split with a micro-splitter to provide a subsample with at least300whole tests of foraminiferawhichwere identified and counted. Fora-miniferal fragments were also counted to provide a dissolution index(Thunell, 1976) and assess any possible bias resulting from the preferen-tial dissolution of themore fragile planktonic species. Planktonic forami-niferal census data were used to estimate the depth of the thermocline(DOT), based on the transfer function developed by Andreasen andRavelo (1997) for the tropical Pacific Ocean, where the 18 °C isothermalis arbitrarily defined as the DOT. Today, the position of the 18 °C isother-mal in the Timor Sea is located at 170 m, corresponding to the depthrange at which water temperature changes most rapidly (Xu et al.,2006). The transfer function equation has a standard error of ±22 m,and an additional ±5 m error is introduced by low species counts inthe core-top database (Andreasen and Ravelo, 1997).
A total of 10–15 specimens of planktonic foraminifer Globigerinoidesruber in the size range of 250–315 μm and 4–6 specimens (>250 μm)of the epifaunal benthic foraminifer Cibicides wuellerstorfi werehandpicked and then ultrasonically cleaned in methanol for less than10 s. Carbon and oxygen stable isotopes were measured using aFinnigan MAT-251 mass spectrometer at the Laboratoire des Sciencesdu Climat et de l'Environnement, Gif-sur-Yvette, France. The externalreproducibility is ~0.06‰ for δ18O and 0.03‰ for δ13C. Isotopic dataare reported relative to the PDB standard through calibration to theNBS19 standard.
AMS 14C ages were obtained on monospecific samples of G. ruber(>250 μm) picked from six intervals. The specimens were ultrasoni-cally cleaned in distilled water and then analyzed with the TandetronAccelerator in Gif-sur-Yvette. Foraminifera 14C ages were adjusted forthe apparent reservoir effect on the ages of surface seawater bysubtracting 400 years; then, dates were calibrated using the radiocar-bon calibration program Fairbanks0701 (Fairbanks et al., 2005). Theages are given in Table 1.
Carbonate weight content (in percent) was measured on dry bulksamples using a standard titration method at the School of MarineSciences, China University of Geosciences (Beijing), and the analyticalerror was ±0.2%.
For grain size analysis, 0.5–1 g samples were placed in beakers with~25 ml water. 2–3 drops of pure hydrogen peroxide and 1–2 ml 10%hydrochloric acid solutionswere added in the beakers. After the carbon-ate and organic carbon fractionswere completely removed, the grain-size(>65 μm, 65–3 μmand b3 μm) distribution of the non-carbonate, terrig-enous material was analyzed with aMastersizer 2000 laser grain size an-alyzer at the School of Marine Sciences, China University of Geosciences(Beijing).
Two other piston cores, SHI9022 and SHI9034, were also studied.The first one was obtained from the Timor Sea (11°35′S, 122°03′E,water depth 2313 m, length 7.11 m), and the second from the Java up-welling area (9°09′S, 111°01′E, water depth 3330 m, length 8.84 m)during a joint French–Indonesian marine geological research programin February 1990 (Fig. 1). The paleoceanographic proxy records ofcores SHI9022 and SHI9034 (the G. ruber δ18O and the planktonic fora-minifera census data) had previously been used to analyze the changesin the Indonesian Seas heat transport pathways (Ding et al., 2002,2006). The AMS 14C ages of the two cores were recalibrated using Fair-banks0701 (Fairbanks et al., 2005).
3. Results
3.1. Age models
The age models for cores MD98-2172, SHI9022, and SHI9034 werebased on the calibrated radiocarbon dates (Table 1) and planktonic
foraminiferal δ18O records. Depths were converted to calendar agesby linear interpolation between stratigraphic control points. Agesfor the older samples in these cores were extrapolated by assuminga constant sedimentation rate throughout the LGM. Clark et al.(2009) recently proposed to extend the duration of the LGM basedon the careful study of maximum extension of ice sheets, andsuggested that it lasted from 26.5 to 19 ka (cal yr BP). However, inthe present study, we decided to stick to the time interval 23–19 kadefined by the EPILOG working group (Mix et al., 2001). This wouldallow us to remain in phase with numerous studies over the last de-cades in which the definition had been used.
The temporal resolution between samples is ~200 years for δ18Orecords, ~400 years for the planktonic foraminifera census in coreMD98-2172 and ~400 years in core SHI9034, and the average sedi-mentation rate is ~26 cm/kyr in core MD98-2172 and ~22 cm/kyrin core SHI9034. The resolution in the upper 500 cm (after 17.5 ka)of core SHI9022 is also ~200 years, but lower downcore, and the av-erage sedimentation rate is ~22 cm/kyr.
3.2. Stable isotope record
Over the time interval covered, theplanktonic δ18O values varied be-tween −0.89‰ and −2.93‰ in core MD98-2172, between −0.83‰and −2.97‰ in core SHI9022, and between −0.74‰ and −3.35‰ incore SHI9034 (Fig. 2). The planktonic and benthic δ18O values wereheavier during the LGM (Fig. 2). Between 19 and 11 ka, the planktonicand benthic δ18O curves displayed sudden distinct declines. Duringthe interval of global warming from the end of the LGM approximately19 ka to the early Holocene 11 ka, virtually every component of the cli-mate system underwent large-scale changes (Clark et al., 2012). Thisdramatic time of global change was triggered by changes in insolation(Clark et al., 2012). The maximum of summer solstice insolation at65°N is at 11 ka (Laskar et al., 2004).Here,wediscuss this 19–11 ka inter-val as the last deglaciation, and the interval after 11 ka as the Holocene.
For core MD98-2172, the planktonic and benthic δ13C values var-ied between 0.64‰ and 1.84‰ and between −0.29‰ and 0.48‰,
Fig. 2. Plots of planktonic and benthic foraminifera δ18O and δ13C and the carbon isotope difference (Δδ13CPF−BF) between planktonic and benthic foraminifera of cores MD-982172,and the δ18O contrast of three cores MD98-2172, SHI9022, and SHI9034. Gray shadings mark the Last Glacial Maximum (LGM, 23–19 ka), the late deglaciation after the meltwaterpulse (MWP)-1A (14.5–11 ka) and strong Southeast (SE) Monsoon periods (~8–6 ka) respectively.
118 X. Ding et al. / Marine Micropaleontology 101 (2013) 115–126
respectively (Fig. 2). The difference in δ13C values between planktonicand benthic foraminifera (Δδ13CPF−BF) was high during the LGM andduring 7–5 ka in the Holocene (Fig. 2).
3.3. Planktonic foraminifera abundance
Variations of planktonic foraminiferal abundances in MD98-2172,SHI9022 and SHI9034 are plotted in Figs. 3–5. The abundance ofGlobigerina bulloides in core MD98-2172 is high during the LGM andat about 8–6 ka in the Holocene, but low during the late deglaciation(Fig. 3). The G. bulloides peak abundance at about 8–6 ka is even moreobvious in core SHI9022 (Fig. 4). For core SHI9034, collected in theJava upwelling area, the abundance of G. bulloides is higher, and thepeak abundance of this species in the Holocene appeared earlier, withhigher values and longer duration (about 10.5–6 ka) than in coresMD98-2172 and SHI9022 (Fig. 5). However, the variation in G. bulloidesabundance during the LGM in core SHI9022 does not resemble that ob-served in core MD98-2172 (Figs. 3 and 4).
Variations in the abundance of Globigerinita glutinata from coreSHI9022 is similar to what was observed for the G. bulloides recordfrom core MD98-2172 (high abundance in the LGM, followed by a
Fig. 3. The abundance of the important planktonic foraminifera species of core MD98-2172after the meltwater pulse (MWP)-1A (14.5–11 ka) and strong Southeast (SE) Monsoon per
decrease during deglaciation and the early Holocene, and an increaseagain about 8–6 ka; Figs. 3 and 4).
In cores MD98-2172 and SHI9022, the abundance ofNeogloboquadrina dutertrei was high during the LGM, reaching astriking maximum peak in the early deglaciation (Figs. 3 and 4),and decreasing abruptly after ~14.5 ka, then continuing to decreasesteadily and slowly into the Holocene. In core SHI9034, on the otherhand, the abundance of this specieswas high during the LGManddegla-ciation, and decreased in the Holocene (Fig. 5).
The evidence for polar ice-margin retreat occurring between 20and 19 ka indicates that the 19 ka meltwater pulse (MWP), whichrepresents a rapid 10 m rise in sea level from the LGM lowstandsometime between 20 and 19 ka, originated from the Northern Hemi-sphere ice sheets (Clark et al., 2009). In the Southern Hemisphere, anabrupt rise in sea level at 14.5 ka is referred to as MWP-1A (Clark etal., 2009).
Similar to N. dutertrei, the dextrally coiled Neogloboquadrinapachyderma was most abundant during the LGM and in the early de-glaciation but decreased rapidly after ~14.5 ka and was only a minorspecies during the Holocene interval in the three cores (Figs. 3–5).
In cores MD98-2172 and SHI9022, the abundance of Globigerinoidessacculifer was low during the LGM, but it increased gradually since the
. Gray shadings mark the Last Glacial Maximum (LGM, 23–19 ka), the late deglaciationiods (~8–6 ka) respectively.
Fig. 4. The abundance of the important planktonic foraminifera species of core SHI9022. Gray shadings mark the Last Glacial Maximum (LGM, 23–19 ka), the late deglaciation afterthe meltwater pulse (MWP)-1A (14.5–11 ka) and strong Southeast (SE) Monsoon periods (~8–6 ka) respectively.
119X. Ding et al. / Marine Micropaleontology 101 (2013) 115–126
deglaciationMWP-1A ~14.5 ka, and reached amaximumpeak during theearly Holocene, and then declined abruptly around 8–6 ka. After ~5 ka,the abundance of G. sacculifer decreased towards the top of both cores(Figs. 3 and 4). This phenomenon is different from that of core SHI9034,in which G. sacculiferwas most abundant during ~5–3 ka period (Fig. 5).
3.4. Carbonate content
Carbonate content of core MD98-2172 was low during the LGM,then increased noticeably after ~17 ka (Fig. 6) and reached thehighest values of 43.5% during the late Holocene.
The foraminifer fragment content indicates that carbonate disso-lution was lower in the LGM and the deglaciation and higher duringthe Holocene, especially its latter part after ~3 ka (Fig. 6). These re-sults suggest, therefore, that dissolution played a minor role in con-trolling the CaCO3 percentage with the exception of the latter partof the Holocene, when the carbonate content was inversely correlatedwith the foraminifer fragment content.
3.5. Sediment grain size changes
Granulometric analysis performed on the noncarbonate fraction ofMD98-2172 samples indicates that the silt fraction (3–65 μm) is the larg-est contributor to the sediments (≥60% in mass of the carbonate-freefraction), the clay fraction (b3 μm) represents about 30%, and the>65 μm coarse fraction represents only between 0.5% and 6% (Fig. 6).The coarse fraction content was higher in the LGM. The coarse-grain
Fig. 5. The abundance of the important planktonic foraminifera species of core SHI9034. Grathe meltwater pulse (MWP)-1A (14.5–11 ka) and strong Southeast (SE) Monsoon periods
content decreased rapidly across the deglaciation and then increasedgradually again in the Holocene, with a noticeable, striking minimum atabout 9–6.5 ka. The 3–65 μmmedium-grain and b3 μm fine-grain varia-tion displayed a marked anticorrelation, with the former showing highcontent in the LGMand low contents from the deglaciation, and the lattershowing the opposite (Fig. 6).
4. Discussion
4.1. High productivity periods
Δδ13CPF−BF values have been used for qualitative estimates of ex-port productivity in tropical and subtropical oceans, with greatervalues during the period representing enhanced productivity (Jianet al., 2001). The Δδ13CPF−BF values of core MD98-2172 were highduring the LGM and ~7–5 ka in the Holocene (Fig. 2).
G. bulloidesmainly occurs in subpolar regions and is also commonlyencountered in upwelling areas and boundary currents in low-latituderegions where surface productivity is high (Bé, 1977; Duplessy et al.,1981; Prell and Curry, 1981; Brock et al., 1992; Martinez et al., 1998;Pflaumann and Jian, 1999). G. glutinata is known to have a wide latitu-dinal distribution. It can tolerate a rather extensive range of tempera-tures and salinities and is moderately susceptible to dissolution. Itsdistribution resembles that of G. bulloides by having a high abundancein mid to high latitudes and also in upwelling regions of low latitudesthat are characterized by fertile waters (Fairbanks et al., 1982; Thunelland Reynolds, 1984; Martinez et al., 1998; Pflaumann and Jian, 1999;
y shadings mark the Last Glacial Maximum (LGM, 23–19 ka), the late deglaciation after(~8–6 ka) respectively.
Fig. 6. Plots for the planktonic foraminifera δ18O, CaCO3 content, foraminifera fragment content and the >65 μm, 3−65 μm, and b3 μm grain-size for core MD98-2172. Gray shadingsmark the Last Glacial Maximum (LGM, 23–19 ka), the late deglaciation after the meltwater pulse (MWP)-1A (14.5–11 ka) and strong Southeast (SE) Monsoon periods (~8–6 ka)respectively.
120 X. Ding et al. / Marine Micropaleontology 101 (2013) 115–126
Kawahata et al., 2002). The abundances of the two species G. bulloidesand G. glutinata were high during the LGM and at about 8–6 ka in theHolocene, low during the late deglaciation and early Holocene in coresMD98-2172 and SHI9022 (Figs. 3 and 4; Table 2).
N. dutertrei is known to be a tropical to temperate species that showsa tolerance to large sea-surface temperature (SST) and sea-surface salin-ity (SSS) changes and is abundant in active current systems, along conti-nental margins, and in upwelling regions (Fairbanks et al., 1982; Curry etal., 1983; Cannariato and Ravelo, 1997; Kawahata et al., 2002). The abun-dance of N. dutertrei is the highest in high productivity areas with awarm temperature and a low salinity (Bé, 1977; Prell and Curry, 1981;Thunell and Reynolds, 1984; Thiede and Jünge, 1992; Pflaumann et al.,1996; Hilbrecht, 1997; Martinez et al., 1998; Pflaumann and Jian, 1999).The dextrally coiled N. pachyderma, an important constituent of subpolarand transitional assemblages, also occurs in low abundances in the equa-torial Indian Ocean (Bé, 1977). This is attributed to cool-water upwellingin coastal regions of the tropical and subtropical Atlantic (Pflaumann andJian, 1999). In the northern South China Sea, N. pachyderma (dex.) hasbeen found in small numbers along with G. bulloides, representing thecoolest faunal assemblage there (Pflaumannand Jian, 1999). The two spe-ciesN. dutertrei andN. pachyderma (dex.)were themost abundant during
Table 2Average abundance of the productivity indicator species and the depth of the thermocline inHolocene.
Time Core Average abundance
G. bulloides N. dutertrei N
Late Holocene ~6 ka MD98-2172 8.3 8.1SHI9022 8.0 10.8SHI9034 10.0 15.7
Middle Holocene 8–6 ka MD98-2172 10.2 6.2SHI9022 8.3 11.3SHI9034 21.7 18.6
Early Holocene 11–8 ka MD98-2172 8.5 5.2SHI9022 6.4 9.4SHI9034 30.0 18.7
Late deglaciation 14.5–11 ka MD98-2172 6.6 8.1SHI9022 4.7 11.8SHI9034 19.4 23.0
Early deglaciation 19–14.5 ka MD98-2172 8.6 12.7SHI9022 5.5 17.0SHI9034 15.1 23.0
LGM 23–19 ka MD98-2172 9.9 11.3SHI9022 4.9 12.5SHI9034 12.3 22.5
the LGM and the early deglaciation in cores MD98-2172 and SHI9022(Figs. 3 and 4; Table 2).
The highest abundance of the warm-water species G. sacculifer is re-stricted to oligotrophic and mesotrophic conditions. The abundance ofthis species can reach high values in areaswith very low export produc-tion rates (Žarić et al., 2005). In cores MD98-2172 and SHI9022, theabundance of G. sacculifer was low during the LGM and around 8–6 ka(Figs. 3 and 4; Table 2).
In summary, our Δδ13CPF−BF data from core MD98-2172 andchanges of the abundances in the productivity indicator species ofplanktonic foraminifera for cores MD98-2172 and SHI9022 all indi-cate that productivity was high during the LGM and early deglacia-tion, then decreased after the deglaciation MWP-1A event, andbecame low during the Holocene, except a brief increase around8–6 ka in the Timor Sea.
4.2. Changes in the depth of thermocline (DOT)
The DOT controls the vertical distribution of planktonic foraminif-era in the upper water column of the oceans (Bé, 1977; Ravelo et al.,1990). G. ruber and G. sacculifer live in the upper layer of the ocean and
the various intervals for cores MD98-2172, SHI9022 and SHI9034 from the LGM to the
of the productivity indicator species (%) DOT (m)
. Pachyderma (dex.) G. glutinata G. sacculifer Ranging Average
1.4 16.3 7.1 127–179 1552.5 18.8 5.1 147–204 1740.7 6.7 8.2 92–163 1262.3 17.1 7.1 164–197 1832.7 16.6 7.9 175–211 1900.8 8.0 6.0 166–210 1872.2 15.7 9.6 159–190 1752.8 13.3 11.8 155–198 1750.5 7.6 4.7 113–191 1583.7 15.5 6.3 162–193 1852.8 15.2 9.0 158–199 1790.5 7.9 4.0 100–156 1226.1 14.0 3.6 158–196 1734.7 18.2 3.8 132–201 173
12.2 8.6 3.1 91–135 1125.4 19.6 1.9 156–202 1795.7 23.3 1.7 184–197 190
11.9 9.4 1.6 83–175 109
121X. Ding et al. / Marine Micropaleontology 101 (2013) 115–126
are, therefore, mixed-layer dwelling species. N. dutertrei, N. pachyderma,G. bulloides and Globorotalia crassaformis are thermocline-dwellingspecies, living below the mixed layer (Ravelo et al., 1990; Huanget al., 2002; Spooner et al., 2005). In general, mixed-layer-dwellingspecies increase in abundance when the DOT deepens, whereasthermocline-dwelling species increase when it shoals (Ravelo et al.,1990; Huang et al., 2002). Consequently, the abundance of mixed-layer-or thermocline-dwelling planktonic foraminiferal species can be used asa qualitative index of DOT variations (Ravelo et al., 1990; Xu et al., 2006).
The thermocline-dwelling foraminifera in core MD98-2172 weremost abundant during the LGM and early deglaciation, and they de-clined rapidly after the deglaciation MWP-1A interval. Instead, themixed-layer dwellers were most abundant during the late deglacia-tion and Holocene intervals (Fig. 7). This is similar to the changes ofthe adjacent core SHI9016 (Spooner et al., 2005) (Fig. 7). However,for core MD98-2172, the abundance plot of mixed-layer-dwellingforaminifera does not parallel the DOT values estimated using the trans-fer function from Andreasen and Ravelo (1997), especially during theLGM (Fig. 7).
The Mg/Ca temperature difference from surface- and thermocline-dwelling planktonic foraminifera (thermal gradient ΔT(G. ruber−
Pulleniatina obliquiloculata)) is interpreted as changes in the DOT, withlarge ΔT values reflecting a shallow DOT (Anand et al., 2003; Xu et al.,2006). Xu et al. (2006) compared the DOT estimated by using the trans-fer functionwithΔT(G. ruber−P. obliquiloculata) in the Timor Sea during Termi-nation II. The results showed the variations in ΔT were consistent withfauna-based DOT estimates. For the last ~23 ka, DOT values estimatedby using the transfer function for core MD98-2172 and core SHI9022exhibited similar changes to those of ΔT(G. ruber−P. obliquiloculata) for coreMD01-2378 (Xu et al., 2006) (Fig. 7). Especially in core SHI9022, thelatest Holocene DOT at ~176 m obtained by using the transfer functionis close to the modern 18 °C isotherm depth (~170 m) in the TimorSea (Xu et al., 2006). Thus, this may suggest that the use of the transferfunction is a better approach for reconstructing DOT variability thanusing the abundance of the mixed-layer or thermocline-dwelling fora-minifera in the Indonesian area.
The changes in DOT of core MD98-2172 and core SHI9022 exhibithigh-frequency variations but the longer term oscillations are similarto those of ΔT(G. ruber−P. obliquiloculata) of core MD01-2378. Three mainplateaus are recognized at 23–19 ka, 14.5–11 ka, and 8–6 ka, separat-ed by periods of shoaling of the thermocline between 19 and 14.5 ka,11–8 ka, and since 6 ka (Fig. 7, Table 2). However, the data for coreSHI9034 do not parallel the DOT changes in the other two coresMD98-2172 and SHI9022. Apart from a remarkable deepening in
Fig. 7. The abundance of the planktonic foraminifera thermocline and mixed-layer dwellers(DOT) of the cores MD98-2172, SHI9022, and SHI9034. The ΔT(G. ruber−P. obliquiloculata) of the c23–19 ka), the late deglaciation after the meltwater pulse (MWP)-1A (14.5–11 ka) and stronDOTs shoal.
the early Holocene ~10–5 ka, the DOT was shallow throughout coreSHI9034 (Fig. 7, Table 2).
4.3. Indonesian Throughflow and monsoon
4.3.1. Last glacial maximum (23–19 ka)The non-carbonate sediment grain size plots for core MD98-2172
show that the >65 μm and 3–65 μm grain size content displayedpeak values during the LGM; in contrast, the b3 μm fine-grain contentwas relatively low (Fig. 6). These results indicate that sediments werecoarsened with the mass terrestrial discharge owing to the decliningsea level during the LGM.
The high-productivity species abundance of the planktonic forami-nifera in both core MD98-2172 and core SHI9022 and the Δδ13CPF−BF
of core MD98-2172, all indicate that productivity was high during theLGM. This is in accord with previous studies (Martinez et al., 1998,1999; Takahashi and Okada, 2000; Gingele et al., 2001; Holbourn etal., 2005; Spooner et al., 2005), which indicates a strong SEMonsoon ac-tivity or enhanced mixing during the LGM and an increasing supply ofterrigenous nutrition owing to the drop in the sea level. However, dur-ing this time interval, the Austral summer insolation (20°S, Dec.) washigh compared to the early Holocene (Berger, 1978) (Fig. 8),which like-ly resulted in strong NW Monsoon activity and heavy rains over landmasses.
The vegetation over land masses indicates drier climates culminat-ing at the LGM (Barmawidjaja et al., 1993; van der Kaars and Dam,1995; Wang et al., 1999). Studies in Australia by Magee et al. (1995),Magee and Miller (1998), Veeh et al. (2000), van der Kaars and DeDeckker (2002), Hesse and McTainsh (2003), and Hesse et al. (2004)also indicated drier conditions during MIS 3 and 2, suggesting greaterinfluence of the SE Monsoon. Wyrwoll and Miller (2001) and van derKaars and De Deckker (2002) estimated the NW Monsoon “switchingon” in Australia at 14 ka. They presumed that the southward shift ofthe ITCZ in the Austral summer may have been considerably restricted,and the ITCZ lay north of the Banda Sea during glacial times, causing thetradewinds to blow across the Banda Sea (Spooner et al., 2005; Xu et al.,2006) (Fig. 9). This essentially has the effect of setting a “perpetual” SEMonsoon into operation (Barrows and Juggins, 2005).
We note that, unlike today when the SEMonsoon promotes a thinnermixed layer and a shallower thermocline depth, the DOT records for bothcoresMD98-2172 and SHI9022were deeper during the LGM (Fig. 8). TheDOT indicated by ΔT(G. ruber−P. obliquiloculata) of mixed-layer-dwelling spe-cies G. ruber and thermocline-dwelling species P. obliquiloculata in core
of cores MD98-2172 and SHI9016 (Spooner et al., 2005). The depths of the thermoclineore MD01-2378 (Xu et al., 2008). Gray shadings mark the Last Glacial Maximum (LGM,g Southeast (SE) Monsoon periods (~8–6 ka) respectively. The arrows indicate that the
Fig. 8. The ΔT(G. ruber−P. obliquiloculata) shown is for core MD01-2378 (Xu et al., 2008). The depths of the thermocline (DOT) and the abundances of G. bulloides from cores MD98-2172,SHI9022, and SHI9034. The insolation curves for 20°S Dec. (red line) and 30°N Jun. (blue line) (Berger, 1978) are also shown for comparison. Gray shadings mark the Last GlacialMaximum (LGM, 23–19 ka), the late deglaciation after the meltwater pulse (MWP)-1A (14.5–11 ka) and strong Southeast (SE) Monsoon periods (~8–6 ka) respectively. The ar-rows indicate that the DOTs shoal.
122 X. Ding et al. / Marine Micropaleontology 101 (2013) 115–126
MD01-2378 situated at themain ITF outflow(Xuet al., 2008) also showeda similar shift.
Today, the South China Sea (SCS) is a main heat and freshwater con-veyor, as it provides a major pathway for surface waters flowing fromthe west Pacific into the Indonesian Seas (Qu et al., 2009). The exportof fresh surface water from the SCS via the Java Sea to the Bali Sea wasshut off during the LGM, when the Sunda land and large parts of theJava Sea were exposed (De Deckker et al., 2002). Previous studies in-cluding circulation modeling experiments suggested that the ITF wasdominated by surface flow during the LGM, when trade winds intensi-fied, precipitation over Borneo decreased (De Deckker et al., 2002;Partin et al., 2007), and the passage to the SCS was blocked (Tozuka etal., 2007). Furthermore, the higher steric height difference betweenthewestern Pacific and the eastern Indian Ocean owing to a “perpetual”SE Monsoon, and the narrowed main ITF outflow pathways resultingfrom the lowered sea level, would lead to a strong surface water ITFand deepening of DOTs during the LGM (Figs. 8 and 9).
The modern hydrography and environments of the Timor Seashow that a high nutrient content (annual mean ~30 μmol/l) and alow oxygen level (annual mean ~1.6 ml/l) occur below water depth~400 m, but the nutrient content in the water column below ~50 mincreases gradually from 2 to 30 μmol/l along with the water depth(Levitus, 1998; Ding et al., 2006). During the LGM, the average DOTsof the MD98-2172 and SHI9022 core sites were ~180 m or deeper(Table 2); hence intensified mixing of the upper water columnowing to a “perpetual” SE Monsoon wind-driven upwelling mayhave elevated the nutrient content, leading to high productivity.Therefore, it is inferred that, because of the stronger surface waterITF, the high productivity was related to the mixing of the upperwater column due to the wind-driven upwelling, but it is not neces-sarily related to shallow DOTs. Instead, the DOTs were affected bythe ITF more than by the monsoonal wind-driven upwelling in the In-donesian Seas.
4.3.2. Deglaciation (19–11 ka)The DOTs reconstructed from cores MD98-2172 and core
SHI9022 shoaled during the early deglaciation and became stabi-lized after 14.5 ka. This is consistent with the DOT indicated bythe ΔT(G. ruber−P. obliquiloculata) found in core MD01-2378 (Xu et al.,2008) (Fig. 8). After reaching their maximum values in the early degla-ciation, the abundances of the high-productivity speciesN. dutertrei and
N. pachyderma (dex.) decreased abruptly after 14.5 ka. At the sametime, the abundance of the high-productivity species G. bulloidesdropped dramaticallywhereas that of the oligotrophic speciesG. sacculiferrose in both cores MD98-2172 and SHI9022 (Figs. 3 and 4). Furthermore,for core MD98-2172, the Δδ13CPF−BF value also dropped gradually(Fig. 2), the coarse-grain content decreased rapidly, and b3 μm fine-grain content was high across the deglaciation (Fig. 6).
To explain these observations, we propose that the sea level roserapidly after 19 ka, resulting in the widening of the main ITF outflowpathways, which made it possible for warm, low-salinity surfacewater to pour into the eastern Indian Ocean through the Lombokand Ombai Straits, and the Timor Sea, which diverted the surfacewater ITF away from the west Pacific Ocean. Furthermore, the declin-ing productivity indicates a weaker “perpetual” SE Monsoon, whichprobably decreased the steric height difference between the west Pa-cific and the eastern Indian Ocean. Thus, the decreasing water pres-sure gradient between the two ocean basins caused surface waterITF weakening and DOT shoaling during the early deglaciation inthe Timor Sea (Fig. 9).
4.3.3. Early Holocene (11–8 ka)The DOT estimated for cores MD98-2172 and SHI9022 shoaled
during the early Holocene. This phenomenon is even better observedat the core MD01-2378 site for which the DOT was indicated by theΔT(G. ruber−P. obliquiloculata) (Fig. 7). The abundances of foraminifer spe-cies indicative of high productivity conditions in cores MD98-2172and SHI9022 all show a slightly declined productivity during this pe-riod (Figs. 3 and 4). At the same time, the content of the oligotrophicspeciesG. sacculifer in the two cores reached a peak value (Figs. 3 and 4).The >65 μm coarse-fraction content rose again in the early Holocene(Fig. 6).
It was estimated that the Australian NW Monsoon switched on at14 ka (Wyrwoll and Miller, 2001; van der Kaars and De Deckker,2002). Spooner et al. (2005) also suggested that the NW Monsoonseemed to strengthen at 10.3 ka, indicating that the bi-annual mon-soonal systemwas most intense at this time. The strong NWMonsoonand heavy rains may have resulted in increasing freshwater and de-creasing surface productivity during the NW Monsoon season atthat period and carried even coarser sediments from the westernland to the studied area.
Fig. 9. The various scenarios postulated for the ITF during the LGM (23-19 ka), deglaciation (19-11 ka), early Holocene (11-8 ka), middle Holocene (8-6 ka) and late Holocene (last~6 ka). The position of Intertropical Convergence Zone in the Austral summer during the LGM (e) is referred from De Deckker et al. (2002) and Xu et al. (2006).
123X. Ding et al. / Marine Micropaleontology 101 (2013) 115–126
At about 9.5 ka, when the postglacial sea level rise reached −30 mbelow the present sea level (Lambeck and Chappell, 2001; Peltier,2002; Peltier and Fairbanks, 2006; Xu et al., 2008; Linsley et al., 2010),the South China and Indonesian Seas became connected, freshwater ex-port from the SCS to the Java Sea was initiated through the open con-nection of the Karimata Strait, and, together with the strong NWMonsoon and heavy rains, large amounts of freshwater flowed intothe Java Sea from the SCS. The freshwater plug at the southern tip of
the Makassar Strait blocked the warm surface flow, thus initiating theenhanced flow at lower, thermocline depths seen in the modern ITF(Tozuka et al., 2007; Xu et al., 2008; Linsley et al., 2010), and conse-quently the DOT shoaled (Figs. 8 and 9).
4.3.4. Middle Holocene (8–6 ka)During the 8–6 ka period, the DOT at the MD98-2172 and SHI9022
core sites deepened again (Fig. 7), with the abundance of foraminifer
124 X. Ding et al. / Marine Micropaleontology 101 (2013) 115–126
species indicative of high productivity conditions showing a prominentpeak and the content of the oligotrophic species declining obviously atthe two core sites (Figs. 3 and 4). Meanwhile, the Δδ13CPF−BF showeda high value in core MD98-2172 (Fig. 2), which indicated that produc-tivity rose once more. However, for the SHI9034 core situated in thesouth Java Sea region and affected by the upwelling from the SE Mon-soonwind, the DOT deepened gradually in the early Holocene, reachinghigh values (maximumDOT) around ~8 ka (Fig. 8),with the abundanceof the high-productivity species G. bulloides rising in the period andreaching a maximum around ~8 ka (Fig. 8).
The solar insolation curve shows that the Boreal summer values in-creased during the early Holocene (Fig. 8). In the eastern Indian Ocean,the enhancedmarine productivitywas directly related to strengtheningof coastal upwelling during periods of increased Boreal summer insola-tion and associated SE Monsoon strength with a precessional cyclicity(Andruleit et al., 2008; Lückge et al., 2009). The Boreal hemisphere sum-mer insolation controls the SE Monsoon strength. However, in theTimor Sea, the intensity of the SE Monsoon indicated by the high pro-ductivity reached its highest values during 8−6 ka, lagging behind thevariations of Boreal summer insolation. Consequently, in the early Holo-cene ~11–8 ka, the SE Monsoon was insufficiently strong to influencethe main ITF outflow in the Timor Sea, but it had obviously affectedthe Java upwelling area. By analogy to today's seasonal variations,with an increasing SE Monsoon, strong westward currents along thesouthern coast of Java occurred, so that the sea-level steric height differ-ence between the southern and northern coasts of Java was increased,implying that the surface water ITF was boosted in the Java upwellingarea, and the DOT at the SHI9034 core site deepened (Figs. 8 and 9).
During the 8–6 ka period, the strongest SE Monsoon appeared inthe area, with the influence of the SE Monsoon not only restrictedto the Java upwelling area but also expanding into all the IndonesianSeas. As a result the steric height difference between the western Pa-cific and eastern Indian Oceans must have increased. It is also inferredthat the intense SE Monsoon transferred more saline Banda Sea waterinto the southern Makassar Strait, and the freshwater plug at thesouthern tip of the Makassar Strait blocking the warm surface flowfrom the west disappeared. Thus, the surface water ITF intensifiedmomentarily, resulting in a deepened DOT and high productivityover all core sites in the study area (Figs. 8 and 9).
4.3.5. Late Holocene (last ~6 ka)After ~6 ka, theDOT shoaled gradually throughout the late Holocene
in the study area (Fig. 7). The abundance of deep-water-dwelling spe-cies G. menardii and P. obliquiloculata in cores MD98-2172, SHI-9022and SHI-9034 rose clearly during this period. The rise may have beenpartially caused by dissolution (Figs. 3–5). The abundances of thehigh-productivity species N. dutertrei and N. pachyderma (dex.) werelow (Figs. 3–5).
The NWMonsoon gradually strengthened owing to the increase ofthe Austral summer insolation during the late Holocene (Berger,1978) (Fig. 8), which caused heavy rains, with large amounts of fresh-water flowing into the Java Sea from the SCS. The freshwater plug atthe southern tip of the Makassar Strait blocked the warm surfaceflow, and thus the thermocline ITF developed again after ~6 ka(Fig. 9).
Of interest is that a detailed investigation of Termination II(135–128 ka) from core MD01-2378 within the same study area(Xu et al., 2006) further indicated that the vertical structure of theTimor outflow during the MIS 5e sea-level highstand was dominatedby thermocline flow, as occurs today (Gordon et al., 2003).
5. Conclusions
Information obtained from cores MD98-2172, SHI9022, and SHI9034collected in the Timor Sea and Java upwelling area reveal the evolution ofsurface hydrographic, and ITF and monsoon activity in the Timor Sea
since the LGM. We found that the importance of deep versus shallowtransport as well as the intensity of ITF activity did vary through timeand this is related to sea-level variations and monsoon activity over thelast ~23 ka.
Our findings suggest that surface ocean productivity was high dur-ing the LGM, and was related to the behaviors of a “perpetual” SEMonsoon resulting from the ITCZ being deflected northward. Warm,surface water ITF transport was dominated by the higher steric heightdifference between the western Pacific and the eastern Indian Oceansowing to a “perpetual” SE Monsoon and the narrowed main ITF out-flow pathways caused by the lowered sea level.
The warm, surface water ITF transport decreased after ~19 kaMWP in the Timor Sea. This resulted from a sea level rise which wid-ened the main ITF outflow pathways. The warm, low-salinity surfacewater ITF was diverted away from the western Pacific Ocean via thewidened pathways. In addition, the weakened “perpetual” SE Mon-soon likely diminished the steric height difference between the west-ern Pacific and the eastern Indian Oceans. The decrease in waterpressure gradient between the two ocean basins led the surfacewater ITF to weaken. At ~9.5 ka, when the postglacial sea level reached−30 m below its present level, large amounts of freshwater flowedinto the Java Sea from the SCS. The freshwater plug located at thesouthern tip of the Makassar Strait blocked the warm surface flow,thus initiating the thermocline ITF seen in themodern Indonesian Seas.
During the period of ~8–6 ka, the intense SE Monsoon controlledby the high Boreal hemisphere summer insolation increased the stericheight difference between the western Pacific and the eastern IndianOceans, and more saline Banda Sea water transferred into the south-ern Makassar Strait due to an intense SE Monsoon. At this time thefreshwater plug at the southern tip of the Makassar Strait that hadblocked the warm surface flow disappeared. Thus, the surface waterITF intensified momentarily around ~8–6 ka.
Over the last ~6 ka, strongNWMonsoon activity resulted in not onlyheavy rains but also large amounts of freshwater being transportedfrom the SCS to the Java Sea. The freshwater plug at the southern tipof theMakassar Strait blocked thewarm surfaceflow, and consequentlythe thermocline ITF enhanced during the last ~6 ka.
In the Timor Sea, the changes in the vertical profile of the ITF wereinfluenced by the glacio-eustatic sea-level changes that have modifiedthe geometry of the pathways within the Indonesian Seas, as well asby the monsoon activity, which was modulated by the changes in theinsolation with a precessional cyclicity.
Acknowledgments
We are grateful to CNRS and CEA for AMS 14C and LSCE for the geo-chemical analysis; Dr. Marie-France Loutre of Université Catholique deLouvain of Belgium for providing the insolation data and Dr. LipingZhou of Peking University for the discussions. We thank Dr. RichardJordan and two anonymous reviewers for their critical remarks and con-structive suggestions. This work was supported by the K. C. Wang Foun-dation of Centre National de la Recherche Scientifique de France, theNational Natural Science Foundation of China (grant no. 40676034),and the fundamental research project of the State Oceanic Administra-tion of China “the paleoclimate research of the eastern Indian Ocean”.Cores were taken by IMAGES cruise IV of the R/V Marion Dufresne andSHIVA cruise, a joint French-Indonesian marine geological research pro-gram.We express our gratitude to the crews of the survey ships for theirinvaluable help.
References
Anand, P., Elderfield, H., Conte, M.H., 2003. Calibration of Mg/Ca thermometry in plank-tonic foraminifera from a sediment trap time series. Paleoceanography 18 (2),1050. http://dx.doi.org/10.1029/2002PA000846.
Andreasen, D.J., Ravelo, A.C., 1997. Tropical Pacific Ocean thermocline depth recon-structions for the last glacial maximum. Paleoceanography 12 (3), 395–413.
125X. Ding et al. / Marine Micropaleontology 101 (2013) 115–126
Andruleit, H., Lückge, A., Wiedicke, M., 2008. Late Quaternary development of the Javaupwelling system (eastern Indian Ocean) as revealed by coccolithophores. MarineMicropaleontology 69, 3–15.
Barmawidjaja, B.M., Rohling, E.J., van der Kaars, W.A., Grazzini, C.V., Zachariasse, W.J.,1993. Glacial conditions of the northern Molucca Sea region (Indonesia).Palaeogeography, Palaeoclimatology, Palaeoecology 101, 147–167.
Barrows, T.T., Juggins, S., 2005. Sea-surface temperatures around the Australian marginand Indian Ocean during the Last Glacial Maximum. Quaternary Science Reviews24 (7–9), 1017–1047.
Bé, A.W.H., 1977. An ecological, zoogeographical and taxonomic review of recentplanktonic foraminifera. In: Ramsay, A.T.S. (Ed.), Oceanic Micropaleontology, 1. Ac-ademic Press, London, pp. 1–100.
Berger, A., 1978. Long-term variation of daily insolation and Quaternary climaticchanges. Journal of the Atmospheric Sciences 35, 2362–2367.
Bray, N.A., Hautala, S., Chong, J., Pariwono, J., 1996. Large-scale sea level, thermocline,and wind variations in the Indonesian Throughflow region. Journal of GeophysicalResearch 101 (C5), 12239–12254.
Brock, J.C., McClain, C.R., Anderson, D.M., Prell, W.L., Hay, W.W., 1992. Southwest mon-soon circulation and environments of recent planktonic foraminifera in the north-western Arabian Sea. Paleoceanography 7 (6), 799–813.
Cannariato, K.G., Ravelo, A.C., 1997. Pliocene–Pleistocene evolution of eastern tropicalPacific surface water circulation and thermocline depth. Paleoceanography 12(6), 805–820.
Clark, P.U., Dyke, A.S., Shakun, J.D., Carlson, A.E., Clark, J., Wohlfarth, B., Mitrovica, J.X.,Hostetler, S.W., McCabe, A.M., 2009. The last glacial maximum. Science 325, 710–714.
Clark, P.U., Shakun, J.D., Baker, P.A., 2012. Global climate evolution during the last de-glaciation. Proceedings of the National Academy of Sciences of the United Statesof America 1–9. http://dx.doi.org/10.1073/pnas.1116619109.
Creswell, G., Frische, A., Peterson, J., Quadfasel, D., 1993. Circulation in the Timor Sea.Journal of Geophysical Research 98 (C8), 14379–14389.
Curry, W.B., Thunell, R.C., Honjo, S., 1983. Seasonal changes in the isotopic compositionof planktonic foraminifera collected in Panama Basin sediment traps. Earth andPlanetary Science Letters 64 (1), 33–43.
De Deckker, P., Tapper, N.J., van der Kaars, S., 2002. The status of the Indo-Pacific WarmPool and adjacent land at the Last Glacial Maximum. Global and Planetary Change35, 25–35.
Ding, X., Guichard, F., Bassinot, F., Labeyrie, L., Fang, N.Q., 2002. Evolution of heat trans-port pathways in the Indonesian Archipelago during last deglaciation. ChineseScience Bulletin 22, 1912–1917.
Ding, X., Bassinot, F., Guichard, F., Li, Q.Y., Fang, N.Q., Labeyrie, L., Xin, R.C., Adisaputra,M.K., Hardjawidjaksana, K., 2006. Distribution and ecology of planktonic foraminif-era from the seas around the Indonesian Archipelago. Marine Micropaleontology58, 114–134.
Duplessy, J.C., Delibrias, G., Turon, J.L., Pujol, C., Duprat, J., 1981. Deglacial warming of thenortheastern Atlantic Ocean: correlationwith the paleoclimate evolution of Europeancontinent. Palaeogeography, Palaeoclimatology, Palaeoecology 35, 121–144.
Fairbanks, R.G., Sverdlove, M., Free, R., Wiebe, P.H., Bé, A.W.H., 1982. Vertical distribu-tion and isotopic fractionation on living planktonic foraminifera from the PanamaBasin. Nature 198, 841–844.
Fairbanks, R.G., Mortlock, R.A., Chiu, T.C., Cao, L., Kaplan, A., Guilderson, T.P., Fairbanks,T.W., Bloom, A.L., 2005. Marine radiocarbon calibration curve spanning 0 to50,000 years B.P. based on paired 230Th/234U/238U and 14C dates on pristine corals.Quaternary Science Reviews 24, 1781–1796.
Ganachaud, A., Wunsch, C., 2000. Improved estimates of global ocean circulation, heattransport and mixing from hydrographic data. Nature 408, 453–457.
Gingele, F.X., De Deckker, P., Hillenbrand, C.D., 2001. Late Quaternary fluctuations ofthe Leeuwin Current and palaeoclimates on the adjacent landmasses: clay min-erals evidence. Australian Journal of Earth Sciences 48, 867–874.
Gordon, A.L., 1986. Interocean exchange of thermocline water. Journal of GeophysicalResearch 91, 5037–5046.
Gordon, A.L., 2005. Oceanography of the Indonesian Seas and their throughflow.Oceanography 18 (4), 14–27.
Gordon, A.L., Fine, R.A., 1996. Pathways of water between the Pacific and Indian Oceansin the Indonesian Seas. Nature 379, 146–149.
Gordon, A.L., Susanto, R.D., Ffield, A.L., 1999. Throughflow within Makassar Strait. Geo-physical Research Letters 26, 3325–3328.
Gordon, A.L., Susanto, R.D., Vranes, K., 2003. Cool Indonesian Throughflow as a conse-quence of restricted surface layer flow. Nature 425, 824–828.
Hesse, P.P., McTainsh, G.H., 2003. Australian dust deposits: modern processes and theQuaternary record. Quaternary Science Reviews 22, 2007–2035.
Hesse, P.P., Magee, J.W., van der Kaars, S., 2004. Late Quaternary climates of the Australianarid zone. Quaternary International 118–119, 87–102.
Hilbrecht, H., 1997. Morphologic gradation and ecology in Neogloboquadrina pachydermaand N. dutertrei (planktic foraminifera) from core top sediments. Marine Micropale-ontology 31, 31–44.
Hirst, A.C., Godfrey, J.S., 1993. The role of Indonesian Throughflow in a global oceanGCM. Marine Micropaleontology 23, 1057–1086.
Holbourn, A., Kuhnt, W., Kawamura, H., Jian, Z., Grootes, P., Erlenkeuser, H., Xu, J., 2005.Orbitally paced paleoproductivity variations in the Timor Sea and IndonesianThroughflow variability during the last 460 kyr. Paleoceanography 20, PA3002.http://dx.doi.org/10.1029/2004PA001094.
Huang, B.Q., Jian, Z.M., Cheng, X.R., Wang, P.X., 2002. Foraminiferal responses to up-welling variations in the South China Sea over the last 220000 years. Marine Mi-cropaleontology 47, 1–15.
Jian, Z., Huang, B., Kuhnt, W., Lin, H., 2001. Late Quaternary upwelling intensity and eastAsian monsoon forcing in the South China Sea. Quaternary Research 55, 363–370.
Kawahata, H., Nishimura, A., Gagan, M.K., 2002. Seasonal change in foraminiferal pro-duction in the western equatorial Pacific warm pool: evidence from sedimenttrap experiments. Deep-Sea Research II 49, 2783–2800.
Lambeck, K., Chappell, J., 2001. Sea level change through the last glacial cycle. Science292, 679–686.
Laskar, J., Robutel, P., Joutel, F., Gastineau, M., Correia, A.C.M., Levrard, B., 2004. A longterm numerical solution for the insolation quantities of the Earth. Astronomy andAstrophysics 428, 261–285. http://dx.doi.org/10.1051/0004-6361:20041335.
Levitus, 1998. World Ocean Atlas. http://www.ifremer.fr/sismer/references/climatologies/.Linsley, B.K., Rosenthal, Y., Oppo, D.W., 2010. Holocene evolution of the Indonesian
Throughflow and the western Pacific warm pool. Nature Geoscience 3, 578–583.Lückge, A., Mohtadi, M., Rühlemann, C., Scheeder, G., Vink, A., Reinhardt, L., Wiedicke,
M., 2009. Monsoon versus ocean circulation controls on paleoenvironmental con-ditions off southern Sumatra during the past 300,000 years. Paleoceanography24, PA1208. http://dx.doi.org/10.1029/2008PA001627.
Magee, J.W., Miller, G.H., 1998. Lake Eyre palaeohydrology from 60 ka to the present:beach ridges and glacial maximum aridity. Palaeogeography, Palaeoclimatology,Palaeoecology 144, 307–329.
Magee, J.W., Bowler, J.M., Miller, G.H., Williams, D.L.G., 1995. Stratigraphy, sedimentology,chronology and palaeohydrology of Quaternary lacustrine deposits at Madigan Gulf,Lake Eyre, South Australia. Palaeogeography, Palaeoclimatology, Palaeoecology 113,3–42.
Martinez, J.I., Taylor, L., De Deckker, P., Barrows, T.T., 1998. Planktonic foraminiferafrom the eastern Indian Ocean: distribution and ecology in relation to the WesternPacific Warm Pool (WPWP). Marine Micropaleontology 34, 121–151.
Martinez, J.I., De Deckker, P., Barrows, T.T., 1999. Palaeoceanography of the last glacialmaximum in the eastern Indian Ocean: planktonic foraminiferal evidence.Palaeogeography, Palaeoclimatology, Palaeoecology 147, 73–99.
Mix, A.C., Bard, E., Schneider, R., 2001. Environmental processes of the ice age: land,oceans, glaciers (EPILOG). Quaternary Science Reviews 20, 627–657.
Molcard, R., Fieux, M., Syamsudin, F., 2001. The Throughflow within Ombai Strait.Deep-Sea Research I 48, 1237–1253.
Müller, A., Opdyke, B.N., 2000. Glacial–interglacial changes in nutrient utilization andpaleoproductivity in the Indonesian Throughflow sensitive Timor Trough, eastern-most Indian Ocean. Paleoceanography 15, 85–94.
Murray, S.P., Arief, D., 1988. Throughflow into the Indian Ocean through the LombokStrait, January 1985–January 1986. Nature 333, 444–447.
Partin, J.W., Cobb, K.M., Adkins, J.F., Clark, B., Fernandez, D.P., 2007. Millennial-scaletrends in west Pacific warm pool hydrology since the Last Glacial Maximum.Nature 449. http://dx.doi.org/10.1038/nature06164.
Peltier, W.R., 2002. On eustatic sea level history: last Glacial Maximum to Holocene.Quaternary Science Reviews 21, 377–396.
Peltier, W.R., Fairbanks, R.G., 2006. Global glacial ice volume and Last Glacial Maximumduration from an extended Barbados sea level record. Quaternary Science Reviews25, 3322–3337.
Pflaumann, U., Jian, Z., 1999. Modern distribution patterns of planktonic foraminifera inthe South China Sea and western Pacific: a new transfer technique to estimate re-gional sea-surface temperatures. Marine Geology 156, 41–83.
Pflaumann, U., Duprat, J., Pujol, C., Labeyrie, L.D., 1996. SIMMAX: a modern analog tech-nique to deduce Atlantic sea surface temperatures from planktonic foraminifera indeep-sea sediments. Paleoceanography 11 (1), 15–35.
Potemra, J.T., Hautala, S.L., Sprintall, J., 2003. Vertical structure of Indonesian Throughflowin a large-scale model. Deep-Sea Research II 50, 2143–2161.
Prell, W.L., Curry, W.B., 1981. Faunal and isotope indices of monsoonal upwelling:Western Arabian Sea. Oceanologica Acta 4, 91–98.
Qu, T., Song, Y.T., Yamagata, T., 2009. An introduction to the South China Sea throughflow:its dynamics, variability, and application for climate. Dynamics of Atmospheres andOceans 47, 3–14. http://dx.doi.org/10.1016/j.dynatmoce.2008.05.001.
Ravelo, A.C., Fairbanks, R.G., Philander, G., 1990. Reconstructing tropical Atlantic hy-drography using planktonic foraminifera and an ocean model. Paleoceanography5, 409–431.
Schiller, A., Godfrey, J.S., McIntosh, P.C., Meyers, G., Wijffels, S.E., 1998. Seasonal near-surface dynamics and thermodynamics of the Indian Ocean and IndonesianThroughflow in a global ocean General Circulation Model. Journal of PhysicalOceanography 28 (11), 2288–2312.
Siedler, G., Church, J., Gould, J., 2001. Ocean circulation & climate: observing andmodeling the Global Ocean, International Geophsics Series. Academic Press. AHarcourt Science and Technology Company 77, 303–314.
Song, Q., Gordon, A.L., 2004. Significance of the vertical profile of the IndonesianThroughflow transport to the Indian Ocean. Geophysical Research Letters 31,L16307. http://dx.doi.org/10.1029/ 2004GL020360.
Spooner, M.I., Barrows, T.T., De Deckker, P., Paterne, M., 2005. Palaeoceanography ofthe Banda Sea, and Late Pleistocene initiation of the Northwest Monsoon. Globaland Planetary Change 49 (1–2), 28–46.
Sprintall, J., Wijffels, S.E., Molcard, R., Jaya, I., 2009. Direct estimates of the IndonesianThroughflow entering the Indian Ocean: 2004–2006. Journal of Geophysical Re-search 114, C07001. http://dx.doi.org/10.1029/2008JC005257.
Takahashi, K., Okada, H., 2000. The paleoceanography for the last 30,000 years in theSoutheastern Indian Ocean by means of calcareous nannofossils. Marine Micropa-leontology 40, 83–103.
Thiede, J., Jünge, B., 1992. Faunal and floral indicators of ocean coastal upwelling (NWAfrican and Peruvian continental margins). In: Summerhayes, C.P., Prell, W.L.,Emeis, K.C. (Eds.), Upwelling systems: evolution since the Miocene: GeologicalSociety London Special Publication, 64, pp. 47–76.
Thunell, R.C., 1976. Optimum indices of calcium carbonate dissolution in deep-sea sed-iments. Geology 4, 525–527.
126 X. Ding et al. / Marine Micropaleontology 101 (2013) 115–126
Thunell, R.C., Reynolds, L.A., 1984. Sedimentation of Planktonic foraminifera: seasonalchanges in species flux in the Panama Basin. Micropaleontology 30, 243–262.
Tozuka, T., Qu, T., Yamagata, T., 2007. Dramatic impact of the South China Sea on the Indo-nesian throughflow. Geophysical Research Letters 34, L12612. http://dx.doi.org/10.1029/ 2007GL030420.
van der Kaars, W.A., Dam, M.A.C., 1995. A 135,000-year old record of vegetational andclimatic change from the Bandung area, West-Java, Indonesia. Palaeogeography,Palaeoclimatology, Palaeoecology 117 (1–2), 55–71.
van der Kaars, S., DeDeckker, P., 2002. A Late Quaternary pollen record fromdeep-sea coreFR10/95, GC17 offshore Cape Range Peninsula, northwestern Western Australia. Re-view of Palaeobotany and Palynology 120 (1–2), 17–39.
van der Kaars, S., Wang, X., Kershaw, P., Guichard, F., Setiabudi, D.A., 2000. A late Quater-nary palaeoecological record from the Banda Sea, Indonesia: patterns of vegetation,climate and biomass burning in Indonesia and northern Australia. Palaeogeography,Palaeoclimatology, Palaeoecology 155, 135–153.
Veeh, H.H., McCorkle, D.C., Heggie, D.T., 2000. Glacial/interglacial variations of sedi-mentation on the West Australian continental margin: constraints from excess230Th. Marine Geology 166, 11–30.
Wang, X., van der Kaars, W.A., Kershaw, P., Bird, M., Jansen, F., 1999. A record of fire, veg-etation and climate through the last three glacial cycles from the Lombok Ridge coreG6-4, eastern Indian Ocean, Indonesia. Palaeogeography, Palaeoclimatology, Palaeo-ecology 147, 241–256.
Wyrwoll, K.H., Miller, G.H., 2001. Initiation of the Australian summer monsoon14,000 years ago. Quaternary International 83–85, 119–128.
Xu, J., Kuhnt, W., Holbourn, A., Andersen, N., Bartoli, G., 2006. Changes in the vertical pro-file of the Indonesian Throughflow during Termination II: evidence from the TimorSea. Paleoceanography 21, PA4202. http://dx.doi.org/10.1029/2006PA001278.
Xu, J., Holbourn, A., Kuhnt, W., Jian, Z.M., Kawamura, H., 2008. Changes in the thermo-cline structure of the Indonesian outflow during Terminations I and II. Earth andPlanetary Science Letters 273, 152–162.
Žarić, S., Donnar, B., Fischer, G., Mulitza, S., Wefer, G., 2005. Sensitivity of planktic fora-minifera to sea surface temperature and export production as derived from sedi-ment trap data. Marine Micropaleontology 55, 75–105.