Two-phase structure of tropical hydroclimate during Heinrich Stadial 1 and its global implications

Quaternary Science Reviews 222 (2019) 105900

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Two-phase structure of tropical hydroclimate during Heinrich Stadial 1 and its global implications Jie Huang a, b, c, *, Shiming Wan a, b, c, Anchun Li a, Tiegang Li d a

Key Laboratory of Marine Geology and Environment, Institute of Oceanology, Chinese Academy of Sciences, Qingdao 266071, China Laboratory for Marine Geology, Qingdao National Laboratory for Marine Science and Technology, Qingdao 266061, China c Center for Ocean Mega-Science, Chinese Academy of Sciences, Qingdao 266071, China d Key Laboratory of Marine Sedimentology and Environmental Geology, First Institute of Oceanography, State Oceanic Administration, Qingdao 266061, China b

a r t i c l e i n f o

a b s t r a c t

Article history: Received 7 May 2019 Received in revised form 22 August 2019 Accepted 25 August 2019 Available online 6 September 2019

Previous studies have proposed that the two-phase structure of Heinrich Stadial 1 (HS1) is a ubiquitous feature. However, few studies have focused on the internal structure of HS1 over a tropical domain; thus, the interhemispheric ocean-atmosphere teleconnections during HS1 are poorly defined. Here, we present high-resolution sedimentological, geochemical, and palaeoceanographic records from the core CG2 in the southern South China Sea (SCS), off the northern Sunda Shelf. The core continuously spans 9.0 e22.0 ky BP. The records suggest two episodes of increased monsoon rainfall during the pre-HS1 (~19.0 e18.0 ky BP) and in the early HS1 (17.5e16.1 ky BP), along with a weaker monsoon rainfall during the late HS1 (16.1e14.7 ky BP). In contrast, a lower precipitation over the Flores Sea occurred during the early HS1, off the southern Sunda Shelf, followed by an enhanced precipitation runoff during the late HS1 (Muller et al., 2012). Notably, these two different precipitation phases characterise a southward Intertropical Convergence Zone (ITCZ) shift over the Sunda Shelf during HS1. These changes are also consistent with other monsoon records, as well as other available data from the tropical to high latitudes of both hemispheres, supporting the global scope of the twofold HS1. We propose that the two episodes of intense monsoon rainfall over the northern Sunda Shelf during the pre-HS1 and the early HS1 observed in our study are closely linked to two different melting events associated with European ice sheet (EIS) retreat. We further propose that the subsequent maximum Atlantic meridional overturning circulation (AMOC) reduction and the migration of the ITCZ to its southernmost position during the late HS1 are tightly connected with the destabilization of the marine-terminating ice-streams of the Laurentide ice sheet (LIS) in the North Atlantic. Tropical positive feedbacks during HS1, in turn, should have played an important role in sustaining the AMOC reduction and Northern Hemisphere (NH) stadial conditions, by continuously lowering the salinity of the low-latitude currents in the Atlantic Ocean. In addition, the tropical ITCZ appears to have been a key bridge in promoting interhemispheric climate connections, by which North Atlantic cooling was rapidly linked to increased Southern Ocean surface westerlies. This, in turn, drove upwelling and CO2 ventilation from the Southern Ocean, and thereby facilitated deglacial warming. Finally, these findings show that tropical climate variation is an important component of deglacial climate changes; understanding how it responds to millennial events like HS1 may help us better understand global teleconnections. © 2019 Elsevier Ltd. All rights reserved.

Keywords: Sunda shelf Tropical ITCZ HS1 European ice sheet AMOC Last deglaciation

1. Introduction The last deglaciation (ca. 19.0 to 11.0 thousand years before

* Corresponding author. Key Laboratory of Marine Geology and Environment, Institute of Oceanology, Chinese Academy of Sciences, Qingdao 266071, China. E-mail address: [email protected] (J. Huang). https://doi.org/10.1016/j.quascirev.2019.105900 0277-3791/© 2019 Elsevier Ltd. All rights reserved.

present, i.e. 19.0e11.0 ky before present (BP), with “present” defined as 1950) (Clark et al., 2012) is the last major climate transition in the Earth's recent geological history. Therefore, it is crucial for understanding recent climate processes, and for the validation of climate models. Palaeoclimate records (Lamy et al., 2007; Anderson et al., 2009; Barker et al., 2009; Denton et al., 2010; Ayliffe et al., 2013; Denniston et al., 2013) and modelling studies (Chiang

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and Bitz, 2005; Lee et al., 2011; He et al., 2013; Chiang et al., 2014) have confirmed opposing climate trends on millennial timescales between the northern and southern middle-to-high latitudes during this interval. The conventional explanation for these opposing climate trends is the bipolar ocean seesaw (Pedro et al., 2018); in this process, weakening North Atlantic overturning during Northern Hemisphere (NH) stadials reduced northward ocean heat transport and stimulated deep-water formation in the Southern Ocean, thereby warming the Southern Ocean and Antarctica (Crowley, 1992; Broecker, 1998; Stocker and Johnsen, 2003). More recently, an alternate (though potentially complementary) mechanism was proposed, suggesting that atmospheric teleconnections forced bipolar coupling during HS1 and the Younger Dryas (YD); in this process, the southward displacement of the global monsoon system and the Southern Hemisphere (SH) westerly wind belt were in response to NH cooling, thus driving upwelling and CO2 ventilation from the Southern Ocean (Anderson et al., 2009; Denton et al., 2010). As a consequence, the rapid reorganization of the oceanic and atmospheric circulations during the last glacial termination (Termination 1), both linked to millennial-scale cooling in the North Atlantic, may have been pivotal in driving the last deglaciation. Currently, the time of HS1 (17.5e14.7 ky BP, also known as the ‘Mystery Interval’ (Denton et al., 2006)) is of special interest to the scientific community. Seeking a better understanding of the NH and SH records, we differentiate the definitions of Heinrich event 1 (H1, centred at ~16.0e15.5 ky BP according to Carlson and Clark, 2012) and HS1, as described by Barker et al. (2009) or Hodell et al. (2017). Namely, a Heinrich event (e.g. H1) corresponds to a massive discharge of the Laurentide ice sheet (LIS) through the Hudson Strait, whereas a Heinrich stadial (e.g. HS1) is the cold period that contains the Heinrich event. Palaeoclimate reconstructions (Fig. 1A) from Greenland (North Greenland Ice Core Project Members, 2004; Landais et al., 2018), the North Atlantic
not et al., 2006; (Sarnthein et al., 2001; Zaragosi et al., 2001; Me Peck et al., 2006; Naughton et al., 2009; Thornalley et al., 2011; Stern and Lisiecki, 2013; Gil et al., 2015; Toucanne et al., 2015; Hodell et al., 2017; Balmer and Sarnthein, 2018; Ng et al., 2018), tropical and subtropical areas (e.g. Kienast et al., 2006; Dupont et al., 2010; Stager et al., 2011; Broecker and Putnam, 2012; Reeves et al., 2013; Deplazes et al., 2013; Rodríguez-Sanz et al., 2013; Zhang et al., 2014; Dutt et al., 2015; Novello et al., 2017; Crivellari et al., 2018), the Southern Ocean (Anderson et al., 2009), and Antarctica (Monnin et al., 2001; Schmitt et al., 2012; Rhodes et al., 2015; Menviel et al., 2018) have ascribed two distinct climate periods to HS1, giving the event a twofold structure in the global scope. However, our understanding of the feedbacks and north-south teleconnections of the tropical climate variation in this global climate shift within HS1 is limited. Despite the increasing number of studies that report abrupt changes in summer monsoon regimes in both hemispheres during HS1 (e.g. Peterson et al., 2000; Wang et al., 2001, 2004; Partin et al., 2007; Muller et al., 2012; Ayliffe et al., 2013; Denniston et al., 2013), the structure of HS1 over the tropical domain, especially in the western Pacific (i.e. the largest heat and moisture source in the world), is still poorly documented. Naughton et al. (2009) and Broecker and Putnam (2012) hypothesised that one possible cause of the significant mid-HS1 transition is a further southward displacement of the Intertropical Convergence Zone (ITCZ); some additional palaeoclimatic records (e.g. Peterson et al., 2000; Wang et al., 2001; Zhang et al., 2014) have clearly documented such a transition in the strength of the summer monsoon during the mid-HS1. The modelling results suggest that the cooling of the Northern Atlantic, associated with meltwater injection, results in a displacement of the ITCZ towards the warmer SH (Chiang and Bitz,

2005), whereas a southward shift of the ITCZ can strengthen the southern westerly winds (SWW), via the weakening of the Hadley circulation and the SH subtropical jet stream (Lee et al., 2011; Chiang et al., 2014). Thus, defining the exact causes for the abrupt climate changes occurring within HS1 is not only pivotal for assessing the climatic impacts of meltwater release into the highlatitude North Atlantic as a consequence of current global warming, but also for improving our understanding of what changes may occur in the future. Lines of evidence from HS1 within the East Asian Monsoon and Indo-Australian Monsoon systems suggest that the ITCZ was predominantly driven by North Atlantic and Antarctic meltwater events. For instance, speleothem records from Hulu Cave in China (Wang et al., 2001), Gunung Buda Cave in Borneo (Partin et al., 2007), and marine sedimentary records from the western coast of Sumatra (Mohtadi et al., 2014) indicate reduced precipitation during HS1. In contrast, the high terrigenous flux during HS1 in the Flores Sea (Muller et al., 2012) points to enhanced precipitation, in agreement with the Flores speleothem records (Ayliffe et al., 2013). Terrestrial evidence for increased rainfall during HS1 is also found in lake records (Muller et al., 2008) and a speleothem record in northern Australia (Denniston et al., 2013). Records from Haozhu Cave (Zhang et al., 2016) and northwest Australia (Kuhnt et al., 2015) have been interpreted as reflecting SH climate variations. However, there still exists a large gap in high-resolution climate records within HS1 between the regions encompassing East Asia, northwest Borneo, Flores, and northern Australia, severely hindering our understanding of the spatial heterogeneity in HS1 precipitation records in the western Pacific. In this scenario, records off the northern Sunda Shelf (in the southern South China Sea (SCS)) could assist in better understanding the temporal and spatial changes in the western Pacific hydroclimate during the early and late HS1, and their relations to high-latitude forcings and meridional ITCZ dynamics. Previous studies have investigated high-resolution stable oxygen isotopes of planktonic foraminifera, Mg/Ca ratio, mass accumulation rate (MAR), grain size, and clay mineral records over the past 22.0 ky BP along core CG2, which was retrieved from the continental slope of the Sunda Shelf (southern SCS) (Hao et al., 2014; Huang et al., 2016). However, the exact causes of the two abnormal shifts within the terrigenous records in core CG2 prior to and during HS1 have not been resolved. The underlying causes of these shifts could possibly yield important implications for global climate change. In this study, by combining new records of radiocarbon dates and major elements in core CG2, we aim to better understand the western Pacific hydrological changes, and their role in interhemispheric teleconnections prior to and during HS1. We compare our records with other hydroclimate records from both hemispheres, as well as other available data from Greenland, the North Atlantic, the Southern Ocean, and Antarctica. Our findings reveal that two episodes of intense monsoon rainfall during the pre-HS1 and the early HS1 over the northern Sunda Shelf were closely linked to two different melting events associated with the European ice sheet (EIS) retreat. In addition, the tropical ITCZ transition during HS1 occurred in a two-phase structure, with the ITCZ located in its southernmost position during the latter half of HS1, and in lockstep with two-step reduced Atlantic meridional overturning circulation (AMOC) and increased SWW, leading to enhanced venting of CO2 from the deep Southern Ocean. We interpret these patterns to indicate a strong influence of the tropical hydroclimate in driving and/or amplifying interhemispheric teleconnections during HS1 and, consequently, in driving the last deglaciation.

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Fig. 1. (A) Circle symbols denote locations of core CG2 and other sites compared in this study. 1 - North Greenland Ice Core Project (NGRIP) ice core (North Greenland Ice Core Project Members, 2004); 2 - ODP Site 980 (McManus et al., 1999; Benway et al., 2010); 3 - core MD01-2461 from the Porcupine Seabight (Peck et al., 2006); 4 - IODP Site
not et al., 2006; U1308 (Hodell et al., 2017); 5 - core EW9302-2JPC from the Labrador Sea (Marcott et al., 2011); 6 - core MD95-2002 from the Bay of Biscay (Zaragosi et al., 2001; Me Toucanne et al., 2015); 7 - core OCE326-GGC14 from the Laurentian Fan (Gil et al., 2015); 8 - core MD08-3180 from the Azores plateau (Balmer and Sarnthein, 2018); 9 - core MD992331 from the Galician margin (Naughton et al., 2009); 10 - core SU81-18 from the Iberian margin (Bard et al., 2000; Gherardi et al., 2005); 11 e core OCE326-GGC5 from the
n Bermuda rise (McManus et al., 2004); 12 - Great Basin lakes (Broecker and Putnam, 2012); 13 - core MD02-2505 from the San L
azaro Basin (Rodríguez-Sanz et al., 2013); 14 - Pete
Lake (Escobar et al., 2012); 15 - ODP Site 1002 (Peterson et al., 2000); 16 - core MD03-2621 from the Cariaco Basin (Deplazes et al., 2013); 17 - GeoB16224-1 from the continental Itza slope off French Guiana (Crivellari et al., 2018); 18 - core ME0005A-24JC from the eastern equatorial Pacific (Kienast et al., 2006); 19 - core GeoB 3910-2 at the lower continental slope of Northeast Brazil (Dupont et al., 2010); 20 - Speleothems in northeastern Brazil (Wang et al., 2004); 21 - Paix~ao Cave, northeastern Brazil (Stríkis et al., 2015, 2018); 22 - Lapa Sem Fim Cave, central-eastern Brazil (Stríkis et al., 2015, 2018); 23 - Salar de Uyuni (Baker et al., 2001); 24 - Jaragu
a cave, Bonito City (Novello et al., 2017); 25 - Bosumtwi Lake, Ghana (Peck et al., 2004); 26 - Tanganyika lake (Tierney et al., 2008); 27 - core TNO57-21 from the southeast Atlantic Ocean (Barker et al., 2009); 28 - core TN057-13 PC from the Southern Ocean (Anderson et al., 2009); 29 - West Antarctic Ice Sheet Divide (WD) ice core (Rhodes et al., 2015); 30 - European Project for Ice Coring in Antarctica (EPICA) Dronning Maud Land (EDML) ice core (EPICA Community Members, 2006); 31 - EPICA Dome C (EDC) ice core (Monnin et al., 2001); 32 - core SO130-289 KL from the Arabian Sea (Deplazes et al., 2013); 33 - Mawmluh Cave, NE India (Dutt et al., 2015); 34 - Haozhu Cave, China (Zhang et al., 2016); 35 - Qingtian Cave, China (Zhang et al., 2014); 36 - Hulu Cave, China (Wang et al., 2001); 37 - core CG2, this study; 38 - Gunung Buda Cave, Borneo (Partin et al., 2007); 39 - western coast of Sumatra (Mohtadi et al., 2014); 40 - Liang Luar Cave, Flores (Ayliffe et al., 2013); 41 - core VM33-80 from the Flores Sea (Muller et al., 2012); 42 - NW Australia (Kuhnt et al., 2015); 43 - Ball Gown Cave, Australia (Denniston et al., 2013); and 44 - Lynch's Crater, NE Australia (Muller et al., 2008). Dashed lines denote the modern position of the Intertropical Convergence Zone (ITCZ) in July and January, respectively (modified from Zhang et al., 2016). (B) Seasonal changes in regional wind patterns associated with latitudinal displacement of the ITCZ during boreal autumn (modified from Kurita, 2012). (C) Seasonal changes in regional wind patterns associated with latitudinal displacement of the ITCZ during boreal winter (modified from Kurita, 2012). Simplified Pleistocene palaeodrainage systems in B and C are based mainly on the analysis of modern sea-floor bathymetry (modified from Alqahtani et al., 2015). B or C is a part of A, which is marked with white dotted lines in A.

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2. Environmental settings The Sunda Shelf is located close to the Equator, near the West Pacific Warm Pool. It has an area of ca. 125,000 km2, making it the largest epicontinental shelf in the world that is located within a tropical humid climate setting (Hanebuth and Stattegger, 2003; Wang et al., 2014; Alqahtani et al., 2015). During the Last Glacial Maximum (LGM, 26.5e19.0 ky BP (Clark et al., 2009)), the growth of continental glaciers decreased the volume of water in the world's oceans, resulting in a reduction of sea level by ~120e130 m (Hanebuth et al., 2000). Consequently, most of the Sunda Shelf was subaerially exposed, and stretched southward from the Indochina peninsula across the Equator to Java, with the coeval shoreline located close to the present submerged shelf break (Voris, 2000; Hanebuth and Stattegger, 2003). Seabed bathymetry maps and shelf-edge canyons have defined several major rivers on the exposed Sunda Shelf during this sea-level lowstand: the palaeoMekong River, palaeo-Johore River, palaeo-Siam River, palaeoNorth Sunda (or Molengraaff) River, and palaeo-East Sunda River (Voris, 2000; Wang et al., 2014; Alqahtani et al., 2015, Fig. 1B and C). The ITCZ, a narrow band of convective winds and intense precipitation situated close to the Equator, migrates seasonally over the Sunda Shelf, in response to interhemispheric temperature gradients (Schneider et al., 2014; Fraser et al., 2014). During the boreal autumn, the ITCZ begins to migrate southward but remains in the NH, resulting in high precipitation over East China and the northern regions of the Indonesian archipelago, whereas southern Indonesia and Australia experience relatively dry conditions (Fig. 1B). In contrast, the opposite effect occurs during the boreal winter (Fig. 1C). 3. Materials and methods 3.1. Core location and chronological framework The gravity core CG2 (6.3928
N, 110.1542
E; water depth of 1239 m) analysed in this study was collected from the southern slope of the SCS during an SCS survey cruise with R/V KE XUE YI HAO in 2012. The coring site is ~300 km apart from the assumed mouth position of the palaeo-North Sunda (or Molengraaff) River during the glacial sea level lowstand of ~120e130 m (Fig. 1B and C), which was believed to provide abundant sediments to the study area during the last deglaciation, based on our previous research (Huang et al., 2016). The lithology of core CG2 is dominated by brownish-grey to dark grey foraminifer-rich or diatom-bearing nannofossil ooze with terrigenous clay. In this study, we mainly focused on the lower section of core CG2 sediments spanning from ~9.0 ky BP to ~22.0 ky BP (core length from 92 cm to 322 cm; 230 cm long) to determine the hydroclimate changes off the northern Sunda Shelf based on sedimentological, geochemical and palaeoceanographic records. We focused particularly on the preHS1 period and on the internal variability of HS1. The age model for core CG2 in this study was established based on twelve accelerator mass spectrometry (AMS) 14C ages. Among these, six AMS 14C new dates were obtained at Beta Analytic Inc. from Neogloboquadrina dutertrei planktonic foraminifera shells (Table 1). The 14C ages were first corrected for reservoir age, and were then calibrated to calendar age using the IntCal13 calibration curve (Reimer et al., 2013). The current regional reservoir age is 373 ± 33 14C years, as determined by the three locations closet to the core CG2 site (Southon et al., 2002; Dang et al., 2004; Bolton et al., 2016). However, recent studies on the global distribution of reservoir ages have pointed to a large spatial heterogeneity over the last termination, owing to the influences of various local processes on surface mixing and ocean-atmosphere gas exchanges (e.g.

Sarnthein et al., 2015 and references herein). Fortunately, there was a previous study which effectively explored the age reservoir €der et al., 2016) from the variability across the last 26 ky BP (Schro same Indo-Australian monsoon region and with comparable waterdepth levels as our study site, mainly based on the differences between coeval atmospheric and planktic 14C age estimates. Thus, to avoid these problems, the 14C ages from each stratigraphic unit in this study were corrected in light of the ad hoc reservoir age applied €der et al. (2016). Evidently, these age corrections differ by Schro from the global average surface ocean reservoir age of ~400 years (Stuiver and Braziunas, 1993). The reservoir ages are ~350 years higher than the global average reservoir age during the LGM. During HS1, the reservoir ages decreased to 650 14C years, and remained close to this level until the YD. In this regard, reservoir ages exhibited a small reservoir age of 200 14C years in the early Holocene, pointing to intense surface ocean-atmosphere exchange at that time. The final age model was constructed using the age modelling software Clam (Blaauw, 2010; version 2.2) and reported results with 95% confidence intervals (Fig. 2). To further test the robustness of our age model, a Pearson correlation analysis was performed and compared with that of an age model assuming a constant surface 14C reservoir age of 373 years (current regional reservoir age) over the entire deglaciation (Fig. 2). Notably, we observed a significant positive correlation between these two age models (r ¼ 0.9995, p＜0.01), indicating that the age-depth relationship differences stemming from these two different reservoir age corrections in this study are relatively small, and actually do not affect our interpretations. The sedimentation rates in core CG2 also do not differ markedly when age models with constant and variable reservoir age corrections are compared (Fig. 2). The only notable difference occurs in the early part of HS1, when the sedimentation rates (29e93 cm ky1) are marginally smaller than those based on a constant reservoir age correction (30e109 cm ky1). Independently of variable reservoir age correction, the sedimentation rates are low (~13 cm ky1) during the LGM, rising to ~93 cm ky1 during the early HS1, remain steady at ~18 cm ky1 through the late HS1 to the Bølling-Allerød (BA) warming, and decrease to ~11 cm ky1 during the early Holocene. With regard to the MAR variations, it is noteworthy that the same situation as the sedimentation rate also occurs in the early part of HS1, i.e. the maxima value of the MAR (~43 g cm2$ky1) (Fig. 3) was slightly lower than that based on a constant reservoir age correction (~50 g cm2$ky1). 3.2. Geochemical analysis The samples used for geochemical analysis were taken at 2 cm intervals, and 114 samples were obtained and analysed from a lower section of core CG2 sediments. These samples were first wetsieved through a 63 mm sieve. The remaining samples (＜63 mm) were treated with 15% H2O2 at 60
C for 1 h and with 1 N HCl at 60
C for 2 h to remove organic matter, calcite, and Fe-Mn oxides. Finally, the sediments were rinsed with deionised water three times, and were dried at 60
C before being ground into powder for elemental composition analysis. Major element oxides were determined by standard X-ray fluorescence (XRF). The samples were prepared as glass disks, using a Rigaku desktop fusion machine. The analyses were performed on a Rigaku ZSX 100e XRF spectrometer at the State Key Laboratory of Isotope Geochemistry, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences (GIGCAS). The calibration lines used in quantification were produced by a bivariate regression of data from 36 reference materials, encompassing a wide range of silicate compositions. The calibrations incorporated matrix corrections based on the empirical TraillLachance procedure, and the analytical uncertainties were mostly

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Table 1 Calibration of 14C dates was performed using Clam software (Blaauw, 2010) with the IntCal13 calibration curve (Reimer et al., 2013) (see part 3 for details). Reservoir ages €der et al., (2016). A constant reservoir age correction of 373 years (Southon et al., 2002; Dang et al., 2004; Bolton et al., 2016) is also applied applied in this study are from Schro for data comparison. Depth (cm)

Sample ID Foraminifera species

92e94

Beta390235 126e128 Beta422856 140e142 OS-98793 166e168 Beta422857 198e200 Beta422858 210e212 Beta404729 216e224 OS-98797 224e230 Beta422860 234e238 Beta422861 244e246 Beta404730 254e258 Beta422863 292e300 OS-98800

Conventional AMS14C age (yr BP)

Age model based on constant reservoir age correction

Age model based on variable reservoir age correction

References

Reservoir correction

Calendar age (cal yr Reservoir BP) correction

Calendar age (cal yr BP) 9079(9019～9253) Huang et al. (2016) 12138(12009～ This study 12316) 13035(12966～ Hao et al. 13091) (2014) 14480(14291～ This study 14718) 16242(16110～ This study 16347) 16682(16581～ Huang et al. 16796) (2016) 16970(16872～ Hao et al. 17062) (2014) 17130(17036～ This study 17213) 17248(17159～ This study 17343) 17359(17267～ Huang et al. 17458) (2016) 17616(17513～ This study 17805) 20131(19874～ Hao et al. 20325) (2014)

Mixed planktonic foraminifera N. dutertrei

8340 ± 30

373

8871(8725～9005) 200

10930 ± 30

373

N. dutertrei

11750 ± 35

373

N. dutertrei

12990 ± 40

373

N. dutertrei

14120 ± 40

373

Mixed planktonic foraminifera N. dutertrei

14350 ± 40

373

14600 ± 50

373

N. dutertrei

14710 ± 50

373

N. dutertrei

14760 ± 40

373

Mixed planktonic foraminifera N. dutertrei

14850 ± 40

373

15020 ± 50

373

N. dutertrei

17450 ± 80

373

12471(12356～ 12536) 13263(13207～ 13367) 14960(14773～ 15089) 16593(16460～ 16723) 17016(16923～ 17118) 17293(17202～ 17375) 17440(17351～ 17515) 17543(17470～ 17624) 17638(17566～ 17722) 17878(17779～ 18016) 20571(20280～ 20754)

600 600 600 600 600 600 600 600 600 600 750

€der et al., 2016; black line) and a constant reservoir age correction of 373 years (Southon et al., Fig. 2. Age models of core CG2, based on variable reservoir age correction (Schro 2002; Dang et al., 2004; Bolton et al., 2016; blue line), were built using the “classical” (non-bayesian) age modelling software Clam (Blaauw, 2010; version 2.2). All radiocarbon ages were calibrated after reservoir correction (see the main text for details). The age-depth models are smooth splines with a common smoothing parameter of 0.3 (Blaauw, 2010) generated through 1000 iterations. Any model with age-depth reversals was removed. Calculations were performed at 95% confidence. The goodness-of-fits are 7.88 and 10.75, respectively (see Blaauw, 2010 for more information on this parameter). Grey-shaded areas indicate associated error ranges (2s). The purple-shaded area indicates sedimentation rates based on variable reservoir age correction, whereas the grass green-shaded area indicates sedimentation rates for a constant reservoir age correction. (For interpretation of the references to colour in this figure legend, the reader is referred to the Web version of this article.)

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Fig. 3. Downcore variations of contents of three grain-size populations, mean grain size, dry bulk density, MAR, d18Oseawater (Huang et al., 2016; Hao et al., 2014), and Si/Al ratios (this study) records in core CG2 from 22.0 to 9.0 ky BP. Detrital fluxes from 232Th and residual calculations from core VM33-80 sediments in the Flores Sea are also shown for comparison (Muller et al., 2012). The chronology of core CG2 has been established by 12 accelerator mass spectrometry (AMS) 14C ages (indicated by black pentagrams). Light-grey shadings mark Last Glacial Maximum (LGM), early Heinrich stadial 1 (HS1) and pre-HS1. Dark-grey shading marks late HS1.

between 1 and 5% (Li et al., 2006). 4. Results and discussion 4.1. Changes in the monsoon rainfall intensity over the northern Sunda Shelf prior to and during Heinrich Stadial 1 (HS1) Several authors have used Si/Al ratios, grain size and MAR records as proxies for terrigenous inputs in the SCS region (e.g. Wehausen and Brumsack, 2002; Steinke et al., 2003; Liu et al., 2005; Wan et al., 2007; Sun et al., 2008; Huang et al., 2011; Liu et al., 2016). Si mostly comes from alumino-silicates and quartz in the southern SCS area, as biogenic opal is of minor importance (average ~2%) throughout the studied core intervals as observed by the smear identification (10 cm interval), whereas Al is a major component of all clay minerals. Thus, the Si/Al ratio in this study mainly reflects the relative contribution of coarse- and fine-grained materials in the terrigenous components of the study area sediments, and high Si/Al ratios are caused by high quartz contents relative to clay contents, i.e. the dominance of silt-sized particles over clay-sized particles (Calvert et al., 1993). The MAR values should be mainly controlled by terrigenous inputs, as carbonate mineral is of minor importance (average ~6%) throughout the studied core intervals, as determined by unpublished X-ray diffraction analysis (10 cm interval). Thus, although calcareous matter may affect at minor scale the MAR values, we assumed that their contribution can be neglected in the context of this study. In general, the variations of Si/Al ratios in core CG2 have a good downcore correlation with those of mean grain size, dry bulk density, and MAR values (Fig. 3), and these correlations are likely associated with the contemporaneous changes in sea level and monsooninduced fluvial input (Steinke et al., 2003; Huang et al., 2016). Here, we also present local seawater oxygen isotopes (d18Oseawater) (Hao et al., 2014) attributed to changes in surface water salinity. Lighter d18Oseawater values are caused by enhanced monsoon rainfall

intensity on land, and the subsequent larger river discharge. Two peaks in mean grain size, dry bulk density, Si/Al ratio, and d18Oseawater values are detected from core CG2 prior to HS1 (~19.0e18.0 ky BP) and during the early HS1 (17.5e16.1 ky BP) (Fig. 3). The increasing terrigenous input episodes may be the result of a lowered sea level and/or an intensified summer monsoon (Steinke et al., 2003; Huang et al., 2016). However, the sea level increased by 10 m in the Sunda Shelf during the pre-HS1 (Hanebuth et al., 2009). The increasing terrigenous inputs detected during this episode might be, therefore, the result of intensified monsoon rainfall, rather than lowered sea level. The strengthening of the summer monsoon is further supported by the Chinese stalagmite records from Hulu Cave (Wu et al., 2009). Similarly, sea level curves for the Sunda Shelf also indicate a slow rise in sea level during the early HS1 (Geyh et al., 1979; Hesp et al., 1998; Hanebuth et al., 2000; Yokoyama et al., 2000, Fig. 4L). Thus, the enhanced terrigenous inputs observed during this period might be closely associated with contemporaneous intensified monsoon rainfall, rather than a lowered sea level. In contrast to the strong sedimentary and hydrological signals during the early HS1, there is no evidence of a similar response from core CG2 in the late HS1 (16.1e14.7 ky BP), along with weak and stable signals in the mean grain size, dry bulk density, Si/Al ratio, MAR, and d18Oseawater values (Fig. 3). This signature may indicate a weaker hydrological response over the northern Sunda Shelf during the late HS1 than that during the early HS1. 4.2. A southward Intertropical Convergence Zone (ITCZ) shift over the Sunda Shelf during HS1 The comparison of our data (northern Sunda Shelf) with that of the southern site VM33-80 (Flores Sea) (Muller et al., 2012) shows a contrasting two-phase pattern within HS1 (Fig. 3). The significant increases in the 232Th-derived detrital input and residual detrital fluxes of core VM33-80 from the Flores Sea reveal that large

J. Huang et al. / Quaternary Science Reviews 222 (2019) 105900

amounts of detrital material were transported into the Flores Sea during the latter half of HS1 (16.1e14.7 ky BP), which is related to enhanced precipitation runoff from surrounding landmasses (Kalimantan, southern Sumatra, Java, and parts of Sulawesi). Conversely, lower values of the 232Th-derived detrital input and residual detrital fluxes were observed in the early HS1 (17.5e16.1 ky BP), and this lack of a detrital signature may indicate a period of significantly lower precipitation over the Flores Sea (off the southern Sunda Shelf) during the early HS1 as compared with during the late HS1 (Muller et al., 2012, Fig. 3). In general, these two different precipitation phases (referring to the above tropical hydrological response from the northern to southern Sunda Shelf) highlight a plausible meridional migration of a higher precipitation pattern that may have characterised the rapid southward ITCZ shift within the interval of HS1. 4.3. Global scale of the two-stage tropical hydroclimate during HS1 The two-phase structure of tropical hydrological records during HS1 observed over the Sunda Shelf in this study is also observed in NH records from eastern China Hulu Cave speleothem H82 (Wang et al., 2001; Wu et al., 2009; Southon et al., 2012, Fig. 4B), northern Borneo Snail Shell Cave speleothem SCH02 (Partin et al., 2007, Fig. 4F), the western USA Great Basin lakes (Broecker and Putnam, 2012, Fig. 4H), and Cariaco Basin ODP Site 1002 (Peterson et al., 2000, Fig. 4G). As shown in Fig. 4B and F, there is a pronounced shift towards heavy oxygen at 16.1 ky BP, suggesting that China and Borneo's monsoons were stronger during the first half of HS1 than during the second half (Wang et al., 2001; Wu et al., 2009; Southon et al., 2012; Partin et al., 2007). Similarly, the deglacial wet period in the Great Basin lakes was confined to the latter half of HS1 and, at least in New Mexico's Estancia Basin, the first half of HS1 was a period of desiccation (Broecker and Putnam, 2012, Fig. 4H). Fig. 4G also shows that during the time interval prior to 16.1 ky BP, the reflectance was lower, suggesting an increased input of debris from rivers and, hence, wetter conditions in the northern part of Venezuela during the first half of HS1. In contrast, the highreflectance interval between 16.1 and 14.7 ky BP represents a time of reduced debris input from Venezuelan rivers (Peterson et al., 2000). Similarly, records from the SH, including west Flores Liang Luar Cave speleothem LR06-C5 (Ayliffe et al., 2013, Fig. 4I), northern Australia Ball Gown Cave speleothems BGC-6 and BGC-14 (Denniston et al., 2013, Fig. 4J), and northeastern Brazil speleothems (Wang et al., 2004, Fig. 4K), also exhibit these features. However, the signal has a sign opposite to those from the north, i.e. with a relatively weak monsoon rainfall during the first half of HS1, followed by a general increase in monsoon rainfall during the latter half of HS1. These aforementioned interhemispheric anti-phase relationships in monsoon rainfall suggest that the tropical ITCZ migrated southward within the interval of HS1, and possibly achieved its maximum southward displacement during the latter half of HS1; then, it shifted northward again until 14.7 ky BP during the final stages of HS1. Similar interhemispheric anti-phase patterns were also observed during the YD period (Griffiths et al., 2009). Notably, these two-stage trends of the hydrological cycle within HS1 are also globally similar in nature with other monsoon records from Western China (Zhang et al., 2014, 2016), Northeastern India (Dutt et al., 2015), the Arabian Sea (Deplazes et al., 2013), Western Sumatra (Mohtadi et al., 2014), Australia (Muller et al., 2008; Kuhnt et al., 2015), the western USA (Rodríguez-Sanz et al., 2013), Central America (Escobar et al., 2012; Deplazes et al., 2013), the Eastern Pacific (Kienast et al., 2006), Eastern South America (Dupont et al., 2010; Stríkis et al., 2015, 2018; Crivellari et al., 2018), Western South America (Baker et al., 2001; Novello et al., 2017), and Africa (Stager et al., 2011; Tierney et al., 2008; Peck et al., 2004) monsoon

7

domains. A summary of the proxy response to the two-phase structure of HS1 and the locations of the sites are shown in Fig. 1A and listed in Table 2. 4.4. Trigger mechanisms Climate models that forced cooling in the North Atlantic by imposing a surface freshwater flux that shut down the AMOC's northward oceanic heat transport indicated an anomalous northward atmospheric heat flux, causing the ITCZ to shift to the warmer hemisphere; in the case of HS1, the shift was to the SH (Chiang and Bitz, 2005). Based on the aforementioned discussions in Sections 4.1, 4.2, and 4.3, we could hypothesise that during both events, i.e. the one prior to HS1 and the other during the early HS1, the ITCZ was probably located further north, although starting to migrate toward south; in the late HS1, the ITCZ was expected to shift further south to the SH. To explore the potential mechanisms of the tropical hydrological response over the northern Sunda Shelf during the pre-HS1 and the early HS1, we compare our results with Eurasian fluvial discharge and ice-rafted detritus (IRD) records from the North Atlantic. Notably, fluvial discharge from the Bay of Biscay, recorded by a
not et al., 2006, branched and isoprenoid tetraether index (Me Fig. 5A) and high-resolution XRF Ti/Ca and Fe/Ca records (Toucanne et al., 2015), shows two peaks of meltwater from the Eurasian glaciers and ice sheets, from ~19.0 to 16.5 ky BP. In addition, the eastern IRD record at ODP Site 980 from Feni Drift (McManus et al., 1999; Benway et al., 2010, Fig. 5A), which is expected to be sensitive to ice rafting from the EIS, also displays two peaks of IRD grain counts during the same period. These combined observations indicate that the two different melting episodes associated with the EIS retreat should be interrupted by a stable period of the EIS with no melting, which is well-expressed in our monsoon rainfall records (a weak monsoon rainfall episode between two episodes of intense monsoon rainfall between ~19.0 and 16.5 ky BP). This implies that an early tropical hydrological cycle over the northern Sunda Shelf during the pre-HS1 and the early HS1 should be closely linked to Eurasian-sourced freshwater forcing. Several studies have shown that the early meltwater discharge
not et al., 2006; Toucanne et al., 2008, (Zaragosi et al., 2001; Me 2015) and IRD deposition (Grousset et al., 2001; Peck et al., 2006) prior to the H1 event in the central and eastern North Atlantic was derived mainly from the EIS. Peck et al. (2006) suggested that instability and meltwater forcing from the Northwest EIS temporarily weakened the AMOC prior to the H1 event, illustrating that even modest ice sheets can have a disproportionate impact on deep-water circulation if the ice sheet is close to the source of the deep-water formation. By investigating the links between the EIS ice-margin fluctuations, Channel River meltwater discharges, and AMOC rates, Toucanne et al. (2015) provided direct evidence that the EIS played a crucial role in the abrupt reorganizations of the global climate system that accompanied the end of the last glacial period. On this basis, Hodell et al. (2017) supported a complex history for HS1 with a reduction in the AMOC during the early part (~19.0e16.1 ky BP), possibly driven by earlier melting of the EIS, whereas the LIS assumed a greater role during the latter half (~16.1e14.7 ky BP). Recently, based on a compilation of sedimentary 231Pa/230Th records in the western and deep high-latitude Atlantic, a similar conclusion was also obtained by Ng et al. (2018) through studying the timing of the AMOC shift on a millennial timescale, in addition to its relationship with the timing of ice sheet and climate changes during the last glacial termination. Evidently, all these studies suggest that the EIS could have played a critical role in the climatic reorganization that accompanied the last deglaciation.

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J. Huang et al. / Quaternary Science Reviews 222 (2019) 105900

Fig. 4. Two-phase structure of HS1 from the tropical to high latitudes. (A) NGRIP d18O (North Greenland Ice Core Project Members, 2004) (black line). (B) Hulu Cave speleothem H82 d18O (Wang et al., 2001; Wu et al., 2009; Southon et al., 2012) (dark brown line). (C) Mean grain size (Huang et al., 2016) (red line) and ratio of Si/Al (blue line, this study) from core CG2. (D) MAR (Huang et al., 2016) (dark blue line) and d18Oseawater values (Hao et al., 2014) (orange line) at core CG2. (E) Detrital fluxes from 232Th (red line) and residual calculations (blue line) in the Flores Sea sediments (Muller et al., 2012). (F) Stalagmite SCH02 d18O records from the northern Borneo (Partin et al., 2007) (red line). (G) Colour reflectance (550 nm) of the Cariaco Basin sediments from ODP Hole 1002C (Peterson et al., 2000) (dark green line). (H) The time period covering the “Big Dry” interval in the Great Basin of the western USA (Broecker and Putnam, 2012) (bold black line). (I) Liang Luar Cave, Flores (Ayliffe et al., 2013) (magenta line). (J) Ball Gown Cave, Australia (Denniston et al., 2013) (red line). (K) Growth phases of speleothems from the northeastern Brazil suggest strengthening of the South American summer monsoon (Wang et al., 2004) (purple diamonds). (L) Relative sea level data from the Bonaparte Gulf (Yokoyama et al., 2000), the Sunda Shelf (Hanebuth et al., 2000), Singapore (Hesp et al., 1998) and the Malacca Strait (Geyh et al., 1979). The chronology of core CG2 has been established by 12 AMS 14C ages (indicated by black pentagrams). Light blue vertical bars indicate the timing of the Younger Dryas (YD), HS1, and LGM, respectively. The peach vertical bar indicates the period of Bølling-Allerød (BA). (For interpretation of the references to colour in this figure legend, the reader is referred to the Web version of this article.)

J. Huang et al. / Quaternary Science Reviews 222 (2019) 105900

9

Table 2 Summary of proxy response to the two phases of Heinrich Stadial 1 (HS1) (Early and Late HS1) used in this study. Three main subdivisions of ‘Cryosphere’, ‘Ocean’, and ‘Atmosphere’ are shown in the left. The location, proxy type, reconstructed environmental parameter, change in the reconstructed parameter, and reference are provided for each record. The amount of plus (minus) signs indicate the strength of the increase (decrease) in the reconstructed parameter for each HS1 phase. A cross represents an unappreciable change in the reconstructed parameter. Early HS1 is placed after Late HS1 for visual consistence with Figs. 4 and 5. Abbreviations: Ice-rafted detritus (IRD); sea surface salinity (SSS); Atlantic meridional overturning circulation (AMOC). Type

Site Region no.

Cryosphere 1

Greenland NGRIP

1

Greenland NGRIP

2

North Atlantic North Atlantic North Atlantic North Atlantic North Atlantic North Atlantic North Atlantic North Atlantic North Atlantic Antarctic

3 4 5 6 7 8 9 10 29

Ocean

Site name

ODP Site 980 MD01-2461 IODP Site U1308 EW9302-2JPC MD95-2002 OCE326-GGC14 MD08-3180 MD99-2331 SU81-18 WD ice core

30

Antarctic

EDC ice core

31

Antarctic

EDML ice core

10

North Atlantic North Atlantic Southern Ocean Southern Ocean North Atlantic

SU81-18

11 27 28 45

Atmosphere 12 13

OCE326-GGC5 TNO57-21 TN057-13 PC West and highlatitude North Atlantic

Subtropical Great Basin Lakes Subtropical MD02-2505

14

Tropical


n Itz
Pete a Lake

15

Tropical

ODP Site 1002

16

Tropical

MD03-2621

17

Tropical

GeoB16224-1

18

Tropical

ME0005A-24JC

19

Tropical

GeoB 3910-2

20

Tropical

21

Tropical

22

Tropical

23

Tropical

Location

75.10
N, 42.32
W 75.10
N, 42.32
W 55.48
N, 14.70
W 51.75
N, 12.92
W 49.88
N, 24.24
W 48.80
N, 45.08
W 47.45
N, 8.53
W 43.07
N, 55.83
W 38.00
N, 31.00
W 42.15
N, 9.69
W 37.77
N, 10.18
W 79.47
S, 112.09
W 75.00
S, 0.00
W 75.10
S, 123.40
E

Proxy

Parameter

d18O 17

d-excess, O-excess IRD (lithics/gram)

Late HS1 Early HS1 16.1 e14.7 ky BP

17.5 e16.1 ky BP

Temperature

e

＋

Evaporation conditions IRD

＋

e

＋

＋＋

References

North Greenland Ice Core Project Members (2004) Landais et al. (2018)

Lithological and geochemical records Ca/Sr

IRD

＋

＋＋＋

McManus et al., 1999; Benway et al. (2010) Peck et al. (2006)

IRD

＋＋

e

Hodell et al. (2017)

Detrital CaCO3 counts

IRD

＋＋

＋

Marcott et al. (2011)

Pediastrum, BIT-index, Ti/Ca, Fe/ Ca, εNd IRD abundances

Meltwater runoff IRD

e

＋＋

＋

＋＋


not et al. Zaragosi et al. (2001); Me (2006); Toucanne et al. (2015) Gil et al. (2015)

d O, planktic C reservoir age

Meltwater

＋＋

＋＋

Balmer and Sarnthein, (2018)

IRD concentrations

IRD

＋＋

e

Naughton et al. (2009)

IRD counts, magnetic susceptibility CH4 concentration

IRD

＋＋＋

＋

Bard et al. (2000)

Methane production Temperature

＋＋

＋

Rhodes et al. (2015)

＋

＋＋

EPICA Community Members, 2006

18

14

18

d O CO2 concentration

Atmospheric CO2

＋

＋＋

Monnin et al. (2001)

37.77
N, 10.18
W 33.70
N, 57.58
W 41.10
S, 7.80
E 53.20
S, 5.10
E e

231

Pa/230Th

e

Gherardi et al. (2005)

230

ee

e

McManus et al. (2004)

Polar foraminifera species

AMOC strength AMOC strength Temperature

ee

231

＋＋

＋＋

Barker et al. (2009)

Opal fluxes

Upwelling

＋＋＋

＋

Anderson et al. (2009)

231

AMOC strength

ee

e

Ng et al. (2018)

40.67
N, 117.67
W 25.00
N, 112.00
W 16.92
N, 89.83
W 10.71
N, 65.17
W 10.68
N, 64.97
W 6.66
N, 52.08
W 0.02
N, 86.46
W

Radiocarbon ages

Lake size

＋

e

Broecker and Putnam, (2012)

d18OSW-IVC

SSS

e

＋

Rodríguez-Sanz et al. (2013)

d O, d C

Precipitation

ee

e

Escobar et al. (2012)

Color reflectance

Discharge

e

＋

Peterson et al. (2000)

Total reflectance

Rainfall

e

＋

Deplazes et al. (2013)

ln(Fe/Ca), BIT index, brGDGT concentration Percentage organic carbon

Terrigenous input Marine primary production Precipitation

ee

＋＋

Crivellari et al. (2018)

＋＋

＋

Kienast et al. (2006)

＋＋

＋

Dupont et al. (2010)

＋



Wang et al. (2004)

d O

Speleothem growth Precipitation

＋＋

＋

Stríkis et al. (2015), 2018

d18O

Precipitation

＋＋

＋＋

Stríkis et al. (2015), 2018

Natural g-radiation

Effective moisture

＋＋

＋

Baker et al. (2001)

4.25
S, 36.35
W Northeastern Brazil 10.17
S, 40.83
W ~o Cave Paixa 12.62
S, 41.03
W Lapa Sem Fim Cave 16.15
S, 44.61
W Salar de Uyuni 20.25
S, 67.50
W

Pa/

230

Pa/

18

Th

Th

13

Palynological and geochemical records 230 Th ages 18

(continued on next page)

10

J. Huang et al. / Quaternary Science Reviews 222 (2019) 105900

Table 2 (continued ) Type

Site Region no.

24

Tropical

25

Tropical

26

Tropical

32

Subtropical

33

Subtropical

34

Subtropical

35

Subtropical

36

Subtropical

37

Tropical

38

Tropical

39

Tropical

40

Tropical

41

Tropical

42

Tropical

43

Tropical

44

Tropical

Site name

Jaragu
a cave

Location

21.08
S, 56.58
W Bosumtwi lake 6.50
N, 1.42
W Tanganyika lake 6.70
S, 29.83
E SO130-289 KL 23.12
N, 66.50
E Mawmluh Cave 25.26
N, 91.88
E Haozhu Cave 30.68
N, 109.98
E Qingtian Cave 31.33
N, 110.37
E H82 (Hulu Cave) 32.50
N, 119.00
E CG2 6.39
N, 110.05
E Borneo SCH02 4.00
N, 114.00
E Western Sumatra 0.78
S, 114.01
E Liang Luar LR06-C5 8.53
S, 120.43
E VM33-80 7.86
S, 121.79
E SO185-18506 15.31
S, 119.50
E Ball Gown Cave 17.03
S, 125.00
E Lynch's Crater 17.62
S, 146.17
E

Proxy

Parameter

d18O

Late HS1 Early HS1 16.1 e14.7 ky BP

17.5 e16.1 ky BP

References

＋

＋＋

Novello et al. (2017)

Magnetic hysteresis

Monsoon strength Aridity

e

＋

Peck et al. (2004)

dDleaf wax

Precipitation

＋

e

Tierney et al. (2008)

Total reflectance

Rainfall

e

＋

Deplazes et al. (2013)

Monsoon strength Monsoon strength Monsoon strength Monsoon strength Monsoon strength Monsoon strength SSS

e

＋

Dutt et al. (2015)

e

＋

Zhang et al. (2016)

e

＋

Zhang et al. (2014)

e

＋

Wang et al. (2001)

e

＋

This study

e

＋

Partin et al. (2007)

e

＋

Mohtadi et al. (2014)

＋

e

Ayliffe et al. (2013)

＋

e

Muller et al. (2012)

＋

e

Kuhnt et al. (2015)

＋

e

Denniston et al. (2013)

＋

e

Muller et al. (2008)

18

d O d18O d18O 18

d O Sedimentological, geochemical and palaeoceanographic records d18O

d18OSW 18

d O Detrital fluxes, residual calculations ln(K/Ca) 18

d O Ash yield, Si/Al, Cyperaceae/ Poaceae, d15N

At the onset of the last deglaciation, insolation-induced melting played an important role in suppressing the AMOC (Clark et al., 2004; He et al., 2013, Fig. 5A). Meltwater that came from the continental interior of Europe (i.e. melting of terrestrial-terminating ice-streams, not from ocean-terminating ice-streams) during the pre-HS1 and the early HS1 potentially contributed to the progressive weakening of the AMOC. This triggered a reduction in northward oceanic heat transport and resulted in NH cooling, finally leading to a steepening of the interhemispheric sea surface temperature gradient (Clement and Peterson, 2008, Fig. 5D). Such greater interhemispheric temperature contrasts at this time could have resulted in a southward displacement of the ITCZ in the tropical Pacific and Atlantic (Wu et al., 2009; Deplazes et al., 2013; Zhang et al., 2014). As expected with an ITCZ located over the northern Sunda Shelf and close to the Equator, the mean grain size and Si/Al ratios of core CG2 sediments all increased to their maximum levels during the pre-HS1 and the early HS1, indicating higher terrigenous and freshwater (lighter d18Oseawater values) inputs from the palaeo-North Sunda (or Molengraaff) River (Fig. 5E and F). During the second phase of HS1, the d18O values determined in Greenland ice cores were more negative (North Greenland Ice Core Project Members, 2004, Figs. 4A and 5B), along with significant changes occurred in the ice d-excess and 17O-excess records (Landais et al., 2018), indicating colder temperatures, and warmer and wetter conditions of Greenland moisture sources, than those in the former stage. As indicated by the pronounced increase in the IRD CaCO3 grain counts from core EW9302-2JPC off Newfoundland (Marcott et al., 2011, Fig. 5A), the largest North Atlantic iceberg discharge and ensuing meltwater discharge from the LIS was

Monsoon strength Monsoon strength Monsoon strength Monsoon strength Precipitation

triggered by enhanced extensive subsurface warming in the Labrador Sea caused by the first phase of AMOC reduction (AlvarezSolas et al., 2010; Marcott et al., 2011). It was suggested to have caused a further weakening of the AMOC (reaching a minimum at approximately 16.0 ky BP) (Naughton et al., 2009; Stern and Lisiecki, 2013; Toucanne et al., 2015; Hodell et al., 2017; Ng et al., 2018), consistent with the timing of the maximum North Atlantic 231 Pa/230Th values (McManus et al., 2004; Gherardi et al., 2005; Ng et al., 2018, Fig. 5C). Subsequent peak cooling in the North Atlantic resulted in the NH summer ITCZ being located in its southernmost position (Muller et al., 2012; Deplazes et al., 2013; Zhang et al., 2014, 2016, Figs. 4E and 5G). As a result, the net consequence during the second phase of HS1 was generally drier conditions in the northern-low to middle latitudes, and wetter conditions in the southern-low latitudes (Fig. 4).

4.5. Implications for global climate changes The ITCZ mean position changes are considered to play a significant role in Atlantic-to-Pacific humidity transport (Broecker et al., 1990; Zaucker and Broecker, 1992). When the ITCZ had a northerly location at the very beginning of the last deglaciation, the trade winds blowing across the tropical Atlantic caused a net transport of freshwater into the Pacific, and consequently, the waters of the North Atlantic basin became very salty. As these warm, salty waters circulated north to the sub-polar regions of the North Atlantic, evaporation caused the surface waters to cool, resulting in the formation of very cold, salty surface waters, which eventually became dense enough to sink in the region south of Greenland and in the Norwegian Sea, forming the North Atlantic

J. Huang et al. / Quaternary Science Reviews 222 (2019) 105900

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Fig. 5. Tropical hydroclimate plays a key role in promoting interhemispheric climate connection. (A) Branched and isoprenoid tetraether (BIT) index records in the Bay of Biscay,
not et al., 2006; Toucanne et al., 2015) (red line), North Atlantic ice-rafted debris records (McManus et al., 1999; Benway et al., 2010; interpreted as Eurasian fluvial discharge (Me Marcott et al., 2011) (blue and black lines), and 21 June insolation at 65
N (Laskar et al., 2004) (dark blue line). (B) NGRIP d18O (North Greenland Ice Core Project Members, 2004) (dark blue line). (C) North Atlantic 231Pa/230Th records from core OCE326-GGC5 (McManus et al., 2004) (red and green lines), SU81-18 (Gherardi et al., 2005) (purple line) and compiled western and deep high-latitude Atlantic (Ng et al., 2018) (black line). (D) Northern Hemisphere (blue line) and Southern Hemisphere (red line) proxy temperature stacks (Shakun et al., 2012). (E) Mean grain size (Huang et al., 2016) (red line) and ratio of Si/Al (blue line, this study) from core CG2. (F) MAR (Huang et al., 2016) (black line) and d18Oseawater values (Hao et al., 2014) (orange line) at core CG2. (G) Detrital fluxes from 232Th (red line) and residual calculations (blue line) in the Flores Sea sediments (Muller et al., 2012). (H) Biogenic opal flux in the Southern Ocean, interpreted as a proxy for changes in upwelling south of the Antarctic Polar Front (Anderson et al., 2009) (red line). (I) Polar foraminiferal species abundance in marine sediment core TN057-21 from the South Atlantic sector of the Southern Ocean (Barker et al., 2009) (grey line). (J) Atmospheric CO2 concentrations in the EPICA Dome C Antarctica ice core (Monnin et al., 2001) (blue circles). (K) EDML d18O, a proxy for Atlantic-sector Antarctic temperature (EPICA Community Members, 2006) (black line). The chronology of core CG2 has been established by 12 AMS 14C ages (indicated by black pentagrams). Vertical bars as indicated for Fig. 4. (For interpretation of the references to colour in this figure legend, the reader is referred to the Web version of this article.)

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Deep Water (NADW) (Broecker et al., 1990; Peterson et al., 2000; Leduc et al., 2007). In contrast, a southerly ITCZ position (e.g. during the first phase of HS1) led to the orogenic blocking of moisture transport by the Andes (Peterson et al., 2000). This fresh water export to the Pacific decreased and overall preferentially returned to the Atlantic Ocean, particularly through the Amazon drainage basin (Arz et al., 1998). This process is suggested to have lowered the salinity of low-latitude currents in the Atlantic Ocean, which eventually further weakened the AMOC (Broecker et al., 1990; Peterson et al., 2000; Leduc et al., 2007). A further southward displacement of the tropical ITCZ (e.g. during the second phase of HS1) was expected to act as a positive feedback to sustain the AMOC reduction and NH stadial conditions until ca. 14.7 ky BP, by continuously lowering the salinity of the low-latitude currents in the Atlantic Ocean. In this scenario, these findings support the notion that the tropical climate variability was not just limited to a passive response to the abrupt climatic changes observed in the North Atlantic region, and that tropical positive feedbacks should have played an important role in the periods of these unstable climate conditions. In addition, the tropical ITCZ appears to have been a key bridge in promoting north-south climate connection and rapidly transmitting the North Atlantic climate signal to the Southern Ocean (Anderson et al., 2009; Denton et al., 2010). Such a connection, which seems to be crucial to transfer abrupt climatic changes between low latitudes and middle-to-high latitudes in the SH, could also have played a key role in the Southern Ocean CO2-degassing, via the southward intensification of the SWW during HS1 (Lamy et al., 2007; Anderson et al., 2009; Skinner et al., 2010; Montade et al., 2015; Menviel et al., 2018). During the first phase of HS1, the initial weakening of the AMOC induced a southward shift of the ITCZ (still located in the NH but very close to the Equator), which could have acted to strengthen the SH westerlies, as indicated by modelling studies (Lee et al., 2011; Chiang et al., 2014). Such an intensification of the SH westerlies, and the associated enhanced Southern Ocean deep convection, could have promoted the release of CO2 by increasing the rate of vertical mixing of CO2-rich deep waters in the Southern Ocean (Menviel et al., 2018), thereby contributing to the first rise in atmospheric CO2 concentration during that period (Fig. 5J). During the second phase of HS1, a further weakening of the AMOC occurred and was associated with the disintegration of the LIS (Toucanne et al., 2015; Hodell et al., 2017; Ng et al., 2018), which would have acted to displace the tropical ITCZ to its southernmost position. A further intensification of the SH westerlies in response to this further southward displacement of the ITCZ enhanced Southern Ocean convection, and led to another abrupt atmospheric CO2 increase (Fig. 5J). This resulted in a peak in the Southern Ocean opal flux (Anderson et al., 2009, Fig. 5H) and stronger Antarctic Intermediate Water (AAIW) and Antarctic Bottom Water (AABW) formations (Menviel et al., 2018), consistent with the reduced ventilation ages in the South Atlantic (Skinner et al., 2010) and the Pacific (Siani et al., 2013). In general, this two-phase rapid southward displacement of the tropical ITCZ during HS1, which linked North Atlantic cooling to increased Southern Ocean surface westerlies via atmospheric teleconnections, was an integral part of the sequence of events that led to rising atmospheric CO2 concentrations (Fig. 5J) and global warming (Fig. 5I and K) during the last termination, and ultimately drove the last deglaciation.

combined sedimentological, geochemical, and palaeoceanographic records show that precipitation was considerably higher during the pre-HS1 (~19.0e18.0 ky BP) and the early HS1 (17.5e16.1 ky BP), whereas a lower precipitation was detected during the late HS1 (16.1e14.7 ky BP). Off the southern Sunda Shelf, an enhanced monsoon rainfall is considered to have persisted in the Flores Sea during the late phase of HS1 (16.1e14.7 ky BP), whereas a weak monsoon rainfall was observed during the early HS1 (17.5e16.1 ky BP) (Muller et al., 2012). Notably, these two different precipitation phases from the northern to southern Sunda Shelf indicated a plausible meridional migration of a higher precipitation pattern, characterizing a rapid southward ITCZ shift within the interval of HS1. The two-stage trend within HS1 is also consistent with other monsoon records from Eastern and Western China, Northeastern India, Arabian Sea, Borneo, Flores, Western Sumatra, Australia, the western USA, Central America, the Eastern Pacific, Eastern and Western South America, and Africa monsoon domains, as well as other available data from Greenland, the North Atlantic, the Southern Ocean, and Antarctica, highlighting the global scope of this twofold HS1. At the onset of the last deglaciation, two melting episodes associated with EIS retreat during the pre-HS1 and the early HS1 resulted in an initial AMOC decline and the corresponding southward shift of the ITCZ, consistent with two episodes of intense monsoon rainfall over the northern Sunda Shelf during the same period. Subsurface ocean warming in the Labrador Sea caused by this AMOC slowdown triggered episodes of iceberg rafting from the LIS during the late HS1, leading to the second phase of maximum AMOC reduction and the migration of the ITCZ to its southernmost position, consistent with increased rainfall in the southern low latitudes. Such a southward shift of the ITCZ during HS1, in turn, should have acted as a positive feedback to sustain the AMOC reduction and NH stadial conditions until ca. 14.7 ky BP, by continuously lowering the salinity of low-latitude currents in the Atlantic Ocean. In addition, this signal of ITCZ southward displacement during HS1 should have been rapidly transmitted to the Southern Ocean through increasing Southern Ocean surface westerlies, finally enhancing Southern Ocean convection and leading to the abrupt atmospheric CO2 increase and, thus, deglacial warming. Under this scenario, the tropical ITCZ appears to have played a key role in promoting a rapid north-south climate connection during HS1. In general, our new and existing highresolution tropical hydroclimate records have provided new insights into the conditions that led to this two-phase structure of HS1, and the critical role that the tropical hydrological cycle played in the sequence of events that drove the last deglaciation. As such, they have helped us better understand these interhemispheric teleconnections.

Acknowledgement We acknowledge the captains and crews of KE XUE YI HAO for their support during the SCS Survey Cruise in 2012. We thank Editor Antje Voelker and two anonymous reviewers for their constructive € der, Sebastian Beil, comments. We are also grateful to Jan F. Schro Kenji M. Matsuzaki, and Peng Zhang for their constructive discussions, and to Xinyu Wang for her contribution to the geochemical analysis. This work was supported by the National Natural Science Foundation of China (41406064, 41622603).

5. Conclusions Appendix A. Supplementary data Our results suggest that two different climate phases occurred over the Sunda Shelf during the period of HS1, and are referred to as the early and late phases of HS1. Off the northern Sunda Shelf, our

Supplementary data to this article can be found online at https://doi.org/10.1016/j.quascirev.2019.105900.

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