INTRODUCTION
High-density radiocarbon (14C) dating and improved age-depth modeling methods are advancing our understanding of recent paleoclimate variability (e.g., Hajdas et al. Reference Hajdas, Ascough, Garnett, Fallon, Pearson, Quarta, Spalding, Yamaguchi and Yoneda2021). One important development is the increased number of high-precision measurements performed on a daily basis. For example, the single stage Accelerator Mass Spectrometer (AMS) installed at the Atmosphere and Ocean Research Institute (AORI), the University of Tokyo has already analyzed over 16,000 samples since 2013 (Yokoyama et al. Reference Yokoyama, Hirabayashi, Goto, Okuno Ji, Haraguchi, Ratnayake and Miyairi2019a; Yokoyama et al. Reference Yokoyama, Miyairi, Aze, Sawada, Ando, Izawa, Ueno, Hirabayashi, Fukuyo and Ota2022).
However, while analytical precision and throughput has greatly improved, the issue of material suitability is of paramount importance (e.g., Bronk Ramsey Reference Bronk Ramsey2009). For example, organic carbon in marine sediments may be derived from multiple sources, complicating interpretation. In coastal settings, terrestrial material discharged from rivers may be relatively old while organic matter produced by plankton will generally be contemporaneous with deposition. The marine organic matter may be subject to a temporally and spatially uncertain reservoir age (e.g., Nakamura et al. Reference Nakamura, Yokoyama, Maemoku, Yagi, Okamura, Matsuoka, Miyake, Osada, Teramura and Adhikari2012; Yokoyama et al. Reference Yokoyama, Miyairi, Aze, Yamane, Sawada, Ando, de Natris, Hirabayashi, Ishiwa and Sato2019b; Fukuyo et al. Reference Fukuyo, Clark, Purcell, Parton and Yokoyama2020; Heaton et al. Reference Heaton, Köhler, Butzin, Bard, Reimer, Austin, Bronk Ramsey, Grootes, Hughen and Kromer2020), while the residence time of terrestrial organic matter is similarly unknown (Carvalhais et al. Reference Carvalhais, Forkel, Khomik, Bellarby, Jung, Migliavacca, Saatchi, Santoro, Thurner and Weber2014; Jiang et al. Reference Jiang, Zhou, Tu, Luo, Ding, Zhu, Liu, Liu, Zhang and Shen2022).
Planktonic foraminifera utilize seawater for calcification and are therefore also affected by the same reservoir effect as the contemporaneously produced marine organic matter, but as rainfall increases, river runoff from proximal rivers supplies an increasingly larger fraction of relatively older (due to above-mentioned residence time) terrestrial carbon to the sediments. This increases the offset of bulk organic matter from foraminiferal radiocarbon (Nakamura et al. Reference Nakamura, Yokoyama, Maemoku, Yagi, Okamura, Matsuoka, Miyake, Osada, Adhikari and Dangol2016; Ishiwa et al. Reference Ishiwa, Yokoyama, Obrochta, Uehara, Okuno Ji and Miyairi2021). Hence, relative changes in the offset between radiocarbon of marine surface dissolved inorganic carbon (DIC) and bulk organic radiocarbon can be used as a proxy for precipitation variability on land. Here, we evaluate radiocarbon offsets offshore of Indonesia to infer past changes in local and global climate.
The regional climate in the Indonesian Archipelago is characterized by seasonal change in precipitation (Spooner et al. Reference Spooner, Barrows, De Deckker and Paterne2005) that is related to the position of the Intertropical Convergence Zone (ITCZ) because tropical convergence causes precipitation (Xie and Arkin Reference Xie and Arkin1997). The seasonal migration distance of the ITCZ over the Indonesian Archipelago is the largest on Earth today (van der Kaars et al. Reference van Der Kaars, Wang, Kershaw, Guichard and Setiabudi2000). During the austral summer, the ITCZ migrates southward, resulting in high monsoonal precipitation over the Indonesian Archipelago (Hobbs Reference Hobbs1998; Figure 1). The ITCZ returns to northerly position (10ºN to 15ºN) during the austral winter, resulting in lower precipitation driven by the dry southeast monsoon (Hobbs Reference Hobbs1998). As the meridional position of the ITCZ reflects the thermal balance between the northern and the southern hemispheres (Chiang and Bitz Reference Chiang and Bitz2005; Donohoe et al. Reference Donohoe, Marshall, Ferreira and McGee2013; Chiang et al. Reference Chiang, Lee, Putnam and Wang2014), an understanding of past precipitation variations in the Indonesian Archipelago will also provide information on global climate change.
Seasonal climatic changes over the Indonesian Archipelago today during both the austral summer and winter. The shaded band indicates the approximate location of the ITCZ. Arrows indicate prominent monsoon wind direction. Blue dots correspond to the sites of cores (1) MD05-2970, (2) SO185-18506, (3) SO185-18479 all in the Timor Sea, (4) V33-80 offshore Flores Island and (5) the location of Snail Shell cave on Borneo Island. (Please see online version for color figures.)
MATERIALS AND METHODS
Regional Setting and Sampling
The Indonesian Throughflow (ITF) transports surface and thermocline waters from the Pacific Ocean to the Indian Ocean. The ITF provides relatively cool and low salinity water to the Indian Ocean with its amount reaching 7.5 Sv (Sv = 106 m3/s) in the Timor Sea (Gordon et al. Reference Gordon, Sprintall, van Aken, Susanto, Wijffels, Molcard, Ffield, Pranowo and Wirasantosa2010). The ITF is the only pathway between the Pacific Ocean and the Indian Ocean because the water depth in the South China Sea and Arafura Sea is shallow (less than 20 m). The Indonesian Archipelago located in the Indo-Pacific Warm Pool. High sea surface temperature in this region invigorates atmosphere circulation, which in turn plays an important role in global climate (Chiang Reference Chiang2009).
Core MD05-2970 was recovered from the eastern Timor Sea (9º25′S, 130º60′E) at a water depth of 437 m below sea level in June 2005 during Cruise MD 148 of the Marion Dufresne. The 28.92-m-long core was subsampled at a 2.5-cm interval.
Radiocarbon Dating
Radiocarbon analysis was performed on TOC (51 samples) and planktonic foraminifera (G. ruber and T. sacculifer; 8 samples) at the Atmosphere and Ocean Research Institute, The University of Tokyo, using a single stage AMS (Yokoyama et al. Reference Yokoyama, Hirabayashi, Goto, Okuno Ji, Haraguchi, Ratnayake and Miyairi2019a; Yokoyama et al. Reference Yokoyama, Miyairi, Aze, Sawada, Ando, Izawa, Ueno, Hirabayashi, Fukuyo and Ota2022). TOC samples (1 mg C) were graphitized following the procedures described in Yokoyama et al. (Reference Yokoyama, Miyairi, Aze, Sawada, Ando, Izawa, Ueno, Hirabayashi, Fukuyo and Ota2022), which is an automated line using an elemental analyzer (Elementar, vario MICRO cube).
Samples for foraminfers were freeze dried then were passed through a >63 μm mesh to isolate the sand fraction containing mostly foraminifera. The >125 µm fraction was examined under a reflected light microscope and approximately 50–100 individuals of T. sacculifer were picked with the aim of obtaining at least 2 mg of calcite. If the number of T. sacculifer in the sample was too low to obtain the needed mass of calcite, specimens of G. ruber were also picked until 2 mg was obtained. Both foraminifera species are surface dwelling but G. ruber prefers a slighly shallower habitat. The foraminifera were treated by first evolving to CO2 gas using H3PO4, and then the CO2 gas was graphitized using H2 as a reducing agent and Fe as a catalyst (Yokoyama et al. Reference Yokoyama, Miyairi, Matsuzaki and Tsunomori2007).
Radiocarbon ages were calibrated to calendar years using MatCal (Lougheed and Obrochta Reference Lougheed and Obrochta2016), the Marine20 calibration curve (Heaton et al. Reference Heaton, Köhler, Butzin, Bard, Reimer, Austin, Bronk Ramsey, Grootes, Hughen and Kromer2020) for foraminifera age dating and the Intcal20 calibration curve (Reimer et al. Reference Reimer, Austin, Bard, Bayliss, Blackwell, Bronk Ramsey, Butzin, Cheng, Edwards and Friedrich2020) for TOC age dating. Following previous work, a local reservoir effect was not applied because the surface ocean was stratified (Sarnthein et al. Reference Sarnthein, Grootes, Holbourn, Kuhnt and Kühn2011; Kuhnt et al. Reference Kuhnt, Holbourn, Xu, Opdyke, De Deckker, Röhl and Mudelsee2015). However, the assumption of a constant marine reservoir age is problematic before 18 ka. Calibrated ages older than 18 ka are considered maximum ages (Sarnthein et al. Reference Sarnthein, Grootes, Holbourn, Kuhnt and Kühn2011). Although the TOC samples contain a mixture marine and terrestrial carbon, the IntCal20 calibration curve is used for the assumption of dominance of terrestrial carbon in the sediment. The IntCal20 curve produces older ages than the Marine20 curve (Heaton et al. Reference Heaton, Köhler, Butzin, Bard, Reimer, Austin, Bronk Ramsey, Grootes, Hughen and Kromer2020). Therefore, the use of the Marine20 to calibrate TOC samples would decrease the TOC age and thus also the difference between the TOC and foraminifera age (hereafter ΔTOC-foram). This is because TOC ages are older than foraminifera ages. Changing the calibration curve affects neither the timing nor direction of changes in ΔTOC-foram. Age modeling was performed in a deterministic 105 iteration Monte Carlo routine, undatable, that considers depth uncertainty and a Gaussian accumulation rate uncertainty between adjacent dates (Lougheed and Obrochta Reference Lougheed and Obrochta2019). Sedimentation uncertainty scaling was set to 0.1 and bootstrapping was set to 30%. Separate foraminifera and TOC age models were created using the above parameters.
RESULTS
The obtained age models for both planktonic and TOC radiocarbon dating and the calibrated calendar year ranges are shown in Appendix A and B. The minimum and maximum calendar ages of foraminifera range between 10.2 ka and 27.4 ka with an average sedimentation rate 17 cm/kyr. The minimum and maximum calendar ages of TOC dating range between 5.0 ka and 29.6 ka. The model based ΔTOC-foram exhibits minima values of ∼0 year offsets at around 18 ka, and is negative before 24 ka (Figure 2A). Offset increase to ∼3000 years in the intervening intervals. ΔTOC-foram calculated from paired measurements from the same horizon exhibits a similar pattern as the model-based offset (Figure 2A). δ13C of TOC samples ranges between –25 and –15‰. δ13C values of TOC are relatively low from 28 ka to 17 ka, are relatively high from 16 ka to 11 ka. As all samples analyzed are taken at a 2.5 cm interval, each sample represents 300 years at the points of the highest sedimentation rate. Hereafter, we will discuss millennial scale climate change using ΔTOC-foram.
Display of (A) the age difference between the TOC and foraminifera age from the same horizon in core MD05-2970 (thin red line with dots) and from the age models (brown thinker line). (B) δ13C values of TOC samples measured by accelerator mass spectrometry.
DISCUSSION
Age Models and Variations inΔTOC-foram
As the mentioned ΔTOC-foram is consistent with the ΔTOC-foram calculated from paired measurements (Figure 2A), we consider only the higher-resolution, modeled offsets in this section. Planktonic foraminifera tests are formed in seawater and therefore solely from marine DIC. However, total organic carbon in sediments is a mixture of terrigenous and marine organic material, with the terrigenous material subject to multiple carbon pathways and residence times. Thus, we generally consider the planktonic foraminifer age depth model to be more reliable and use it as the basis for the following discussion.
The negative ΔTOC-foram before 24 ka is consistent with previous work suggesting a non-stationary reservoir age at this time (Sarnthein et al. Reference Sarnthein, Grootes, Holbourn, Kuhnt and Kühn2011). The ΔR in the Timor Sea appears to have been relatively constant near 0, yet it was significantly larger prior to 18 ka, likely more than 1000 years, possibly due to ocean circulation changes and lack of stratification (Sarnthein et al. Reference Sarnthein, Grootes, Holbourn, Kuhnt and Kühn2011). Thus, we avoid making paleoclimatic interpretations prior to 18 ka, In particular, the negative values prior to ∼24 ka suggest a substantial increase in reservoir age, a major decrease in terrestrial organic matter residence time, or a combination of both.
The interpretation that ΔR was constant after 18 ka suggests that ΔTOC-foram change after this time is unrelated to ocean circulation changes. While variations in ocean productivity could affect ΔTOC-foram, the relatively constant δ13C of TOC samples suggests that stable productivity and little. vegetation changes across the Indonesian Archipelago (Dubois et al. Reference Dubois, Oppo, Galy, Mohtadi, van Der Kaars, Tierney, Rosenthal, Eglinton, Lückge and Linsley2014). δ13C values obtained by AMS may reflect fractionation in the ion source, as opposed to IRMS measurements. This could cause higher variability in our data, but as the values are relatively constant, we believe the interpretation of the δ13C values is relatively straightforward. Thus, we interpret variations in ΔTOC-foram as reflecting changes in the residence time of terrestrial organic carbon.
We further suggest that varying precipitation amount drives changes in TOC residence time (e.g., Carvalhais et al. Reference Carvalhais, Forkel, Khomik, Bellarby, Jung, Migliavacca, Saatchi, Santoro, Thurner and Weber2014; Jiang et al. Reference Jiang, Zhou, Tu, Luo, Ding, Zhu, Liu, Liu, Zhang and Shen2022; Appendix C). Turnover times of terrestrial carbon is mainly influenced by precipitation over the Indonesian Archipelago (Carvalhais et al. Reference Carvalhais, Forkel, Khomik, Bellarby, Jung, Migliavacca, Saatchi, Santoro, Thurner and Weber2014), with higher precipitation being associated with shorter turnover times. The soil is easily eroded and transported to the Timor Sea resulting in the younger TOC of material sampled from the ocean floor. On the other hand, turnover times of terrestrial carbon become longer when the precipitation is low because soil less effectively eroded. Therefore, we expect an antiphased relationship between ΔTOC-foram and precipitation amount. Although terrigenous material input to the Sunda Trench is limited today (Omura et al. Reference Omura, Ikehara and Arai2017), terrestrial carbon contents of coastal sediment cores are ∼20% (Zhou et al. Reference Zhou, Martin and Müller2019). Organic carbon in the sediments recovered from the Bengal Fan at a water depth of 240 m was stored for more than 1000 years (French et al. Reference French, Hein, Haghipour, Wacker, Kudrass, Eglinton and Galy2018). Considering that our sediment core was recovered from a depth of 437 m, it is reasonable to interpret that terrestrial carbon input likely influenced ΔTOC-foram variability at this site.
The Meridional Shift of the ITCZ During the Deglaciation
During the last deglaciation (18–11.7 ka), the thermal balance between the northern and southern hemispheres changed dramatically causing a dynamic meridional migration of the ITCZ (e.g., Yokoyama et al. Reference Yokoyama, Suzuki, Siringan, Maeda, Abe-Ouchi, Ohgaito, Kawahata and Matsuzaki2011; De Deckker et al. Reference De Deckker, Barrows and Rogers2014) A portion of migration was driven by partial collapse of northern hemisphere ice sheets (Chiang and Bitz Reference Chiang and Bitz2005; Yokoyama and Purcell Reference Yokoyama and Purcell2021). The northern hemisphere cooled because of the weakened Atlantic Meridional Overturning Circulation (AMOC), allowing more heat to be related by the southern hemisphere during Heinrich Stadial 1 (HS1; 18.0–15.0 ka) and the Younger Dryas (YD; 12.9–11.7 ka; McManus et al. Reference McManus, Francois, Gherardi, Keigwin and Brown-Leger2004; Chiang and Bitz Reference Chiang and Bitz2005; Obrochta et al. Reference Obrochta, Crowley, Channell, Hodell, Baker, Seki and Yokoyama2014). A resulting teleconnection is movement of the ITCZ (and its associated precipitation band) southward over the Indonesian Archipelago (Yokoyama et al. Reference Yokoyama, Purcell, Marshall and Lambeck2006; Safaierad et al. Reference Safaierad, Mohtadi, Zolitschka, Yokoyama, Vogt and Schefuß2020). In contrast, a vigorous AMOC during the Bølling–Allerød warming interval (B/A; 15.0–12.9 ka; McManus et al. Reference McManus, Francois, Gherardi, Keigwin and Brown-Leger2004) reduced the thermal balance between the hemispheres, warmed the northern hemisphere, and shifted the ITCZ northward away from the Indonesian Archipelago. Thus, precipitation over the southern margin of the Indonesian Archipelago decreased during B/A (Kuhnt et al. Reference Kuhnt, Holbourn, Xu, Opdyke, De Deckker, Röhl and Mudelsee2015).
During HS1, decreased ΔTOC-foram suggests high precipitation over the Timor Sea and adjacent landmasses and implies a southward shift of the ITCZ resulting in wet conditions over the Timor Sea (Figure 3A). As the northern hemisphere cooled during HS1 due to weakened AMOC (Figure 3G; McManus et al. Reference McManus, Francois, Gherardi, Keigwin and Brown-Leger2004), the ITCZ likely shifted southward due to the changing thermal balance between both hemispheres (Chiang and Bitz Reference Chiang and Bitz2005; Donohoe et al. Reference Donohoe, Marshall, Ferreira and McGee2013; Chiang et al. Reference Chiang, Lee, Putnam and Wang2014). This is consistent with a relatively low ΔTOC-foram that implies wet conditions during HS1. Wet conditions over the Timor Sea and adjacent landmasses are also consistent with paleoclimatic records from the Indonesian Archipelago region (Kuhnt et al. Reference Kuhnt, Holbourn, Xu, Opdyke, De Deckker, Röhl and Mudelsee2015).
The age difference between the TOC and foraminifera age from the same horizon of core MD05-2970 (thin red line with dots) and that from the age models (brown thicker line). (B) core SO185-18506 (ln(K/Ca)) indicating changes in terrigenous flux in the Timor Sea (Kuhnt et al. Reference Kuhnt, Holbourn, Xu, Opdyke, De Deckker, Röhl and Mudelsee2015). (C) core SO185-18479 (ln(K/Ca)) indicating changes in terrigenous flux from the Timor Sea (Kuhnt et al. Reference Kuhnt, Holbourn, Xu, Opdyke, De Deckker, Röhl and Mudelsee2015). (D) core V33-80 232Th flux indicating changes in terrigenous flux in the Flores Sea (Muller et al. Reference Muller, McManus, Oppo and Francois2012). (E) Snail Shell cave δ18O indicating precipitation changes over Borneo Island (Muller et al. Reference Muller, McManus, Oppo and Francois2012). (F) Titanium content of the sediment core recovered from ODP Site 1002 indicating changes in terrigenous flux in the Cariaco Basin (Haug et al. Reference Haug, Hughen, Sigman, Peterson and Röhl2001). (G) core OCE326-GGC5 231Pa/230Th indicating the strength of the Atlantic Meridional Overturning Circulation from the western subtropical Atlantic Ocean (McManus et al. Reference McManus, Francois, Gherardi, Keigwin and Brown-Leger2004). Shaded areas correspond to Heinrich Stadial 1 (18–14.7 ka), the Bølling–Allerød warming interval (14.7–12.9 ka) and the Younger Dryas (12.9–11.7 ka).
The K/Ca ratio of cores SO185-18506 and SO185-18479 recovered from the western Timor Sea is high during HS1 indicating high terrigenous input to the western Timor Sea caused by the increased precipitation amount (Figure 3B, C; Kuhnt et al. Reference Kuhnt, Holbourn, Xu, Opdyke, De Deckker, Röhl and Mudelsee2015). Both the 232Th flux and the residual flux (non-biogenic components) of core V33-80 recovered from offshore Flores Island increased due to enhanced precipitation runoff into the Flores Sea from the surrounding landmass (Figure 3D; Muller et al. Reference Muller, McManus, Oppo and Francois2012). Terrigenous input of core ST08, retrieved from offshore and to the southwest of Sumba Island, also exhibited an increase as a result of higher precipitation (Ardi et al. Reference Ardi, Maryunani, Yulianto, Putra and Nugroho2020). The southward ITCZ and the associated wet northwest monsoon are causes of increased precipitation in the Flores Sea region (Muller et al. Reference Muller, McManus, Oppo and Francois2012), consistent with low ΔTOC-foram. This is consistent with the speleothem δ18O record from northern Borneo Island, which shows a trend toward more positive values during HS1. This reflects dry conditions in that region due to a southward shift of the ITCZ away from Borneo Island (Figure 3E; Partin et al. Reference Partin, Cobb, Adkins, Clark and Fernandez2007). Therefore, the precipitation band associated with the ITCZ was located between northern Borneo Island and the Timor Sea during HS1.
Although dry conditions are revealed by grain size analysis and terrigenous input proxy variability of core GeoB 10053-7 recovered from south of Java Island, it is also consistent with the southward displacement of the ITCZ during HS1 (Mohtadi et al. Reference Mohtadi, Oppo, Steinke, Stuut, Pol-Holz, Hebbeln and Lückge2011). The precipitation band associated with the ITCZ was located further south of the Java Sea. Modeling results are also consistent with this inferred southward position of the ITCZ during HS1 (Du et al. Reference Du, Russell, Liu, Otto-Bliesner, Oppo, Mohtadi, Zhu, Galy, Schefuß and Yan2023). In contrast, no major changes in annual rainfall were observed in northwestern Australia (De Deckker et al. Reference De Deckker, Barrows and Rogers2014), suggesting that the degree of the southward migration of the ITCZ and the associated precipitation change was insufficiently large to change the local climate in this region. Although the effects of HS1 have not been identified in the Australian region (De Deckker et al. Reference De Deckker, Barrows and Rogers2014), the northern hemisphere cooling during HS1 may have pushed the ITCZ southward because the one of significant factors of the ITCZ movement is thermal balance between both hemispheres.
The ΔTOC-foram increase was substantial during the B/A interval implying that the ITCZ shifted northward with a prominent dry southeast monsoon. The southern hemisphere, especially around Antarctica, cooled (as seen through the Antarctic Cold Reversal; ACR; Lemieux-Dudon et al. Reference Lemieux-Dudon, Blayo, Petit, Waelbroeck, Svensson, Ritz, Barnola, Narcisi and Parrenin2010) during the northern hemisphere warming (B/A). The shifting thermal balance would have pushed the ITCZ northward. However, the ΔTOC-foram did not remain high during the B/A interval because of the exposure of the Sunda Shelf (at ∼14 ka) which forced a shift of the atmospheric circulation over the Indonesian Archipelago (Du et al. Reference Du, Russell, Liu, Otto-Bliesner, Gao, Zhu, Oppo, Mohtadi, Yan and Galy2021). The low 232Th flux and residual flux in the Flores Sea (Muller et al. Reference Muller, McManus, Oppo and Francois2012), and the low K/Ca ratio in the western Timor Sea (Kuhnt et al. Reference Kuhnt, Holbourn, Xu, Opdyke, De Deckker, Röhl and Mudelsee2015), are all consistent with decreased precipitation during the B/A (Figure 3A, B, C, D). In contrast, increased precipitation during the B/A was recorded in the Borneo speleothem (Figure 3E; Partin et al. Reference Partin, Cobb, Adkins, Clark and Fernandez2007). All these records are consistent with the northward shift of the ITCZ because of AMOC recovery (Figure 3G) and associated vigorous thermal exchange between both hemispheres as recorded in the Atlantic Ocean (McManus et al. Reference McManus, Francois, Gherardi, Keigwin and Brown-Leger2004). The relatively high titanium input to the Cariaco Basin (Haug et al. Reference Haug, Hughen, Sigman, Peterson and Röhl2001) off the coast of Venezuela, also supports a northward movement of the ITCZ during the B/A over the western Atlantic indicating that the ITCZ moved similarly in both the eastern and the western Pacific in this interval (Figure 3F).
AMOC was reduced and the meridional thermal exchange was inhibited during the YD (Matsumoto and Yokoyama Reference Matsumoto and Yokoyama2013). Although AMOC remained stronger during the YD than during HS1 (Figure 3G; McManus et al. Reference McManus, Francois, Gherardi, Keigwin and Brown-Leger2004), our ΔTOC-foram was low and the K/Ti ratio in the Timor Sea increased during the YD interval suggesting precipitation had increased because of the southward shift of the ITCZ which is similarly observed during HS1 (Figure 3A, B, C; Kuhnt et al. Reference Kuhnt, Holbourn, Xu, Opdyke, De Deckker, Röhl and Mudelsee2015). The positive shift of the speleothem δ18O record from Borneo Island and the low terrestrial material input to the Ciriaco Basin, both imply a precipitation decrease. This supports the interpretation that the ITCZ was in a southward location globally during the YD (Figure 3E, F; Haug et al. Reference Haug, Hughen, Sigman, Peterson and Röhl2001; Partin et al. Reference Partin, Cobb, Adkins, Clark and Fernandez2007). On the other hand, the 232Th flux and residual flux in the Flores Sea remained steady during the YD (Figure 3D; Muller et al. Reference Muller, McManus, Oppo and Francois2012), possibly because either a reduced meridional movement of the ITCZ as a result of AMOC production, or perhaps chronological uncertainty hinders interpretation of the record.
CONCLUSION
By performing extensive radiocarbon dating of both the TOC and planktonic foraminifera in a coastal marine core, it was possible to reconstruct precipitation changes driven by the ITCZ movement in the Timor Sea region. During HS1 and the YD, the Timor Sea region was wet because of the southerly position of the ITCZ. Conversely, the northward shift of the ITCZ resulted in dry conditions in the Timor Sea during the B/A. The ITCZ movement during the last deglaciation was consistent with the thermal balance/shift between both hemispheres that was mainly controlled by AMOC strength. As precipitation is a key factor of climate change, precipitation reconstruction using the extensive radiocarbon dating may be a useful tool to determine paleoclimate changes in different regions over the Indonesian Archipelago.
SUPPLEMENTARY MATERIAL
To view supplementary material for this article, please visit https://doi.org/10.1017/RDC.2024.13
ACKNOWLEDGMENTS
The work presented here is partly supported by the Japan Society for the Promotion of Science (JSPS) through Grants-in-Aid Nos. 20H00193 and 15KK0151 to Y.Y. This paper is a contribution to INQUA commission on Coastal and Marine Processes and the PAGES PALSEA program. We also express our gratitude to Quan Hua and anonymous reviewers for their constructive comments that improved the quality of this manuscript.
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