Last Glacial - Holocene climate variability in the Atlantic sector of the Southern Ocean

Quaternary Science Reviews 135 (2016) 115e137

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Last Glacial - Holocene climate variability in the Atlantic sector of the Southern Ocean Wenshen Xiao a, b, *, Oliver Esper a, Rainer Gersonde a a b

Alfred-Wegener-Institut, Helmholtz-Zentrum für Polar- und Meeresforschung, Bremerhaven D-27568, Germany State Key Laboratory of Marine Geology, Tongji University, Shanghai 200092, China

a r t i c l e i n f o

a b s t r a c t

Article history: Received 28 September 2015 Received in revised form 10 December 2015 Accepted 22 January 2016 Available online xxx

The Southern Ocean plays a major role in the glacial/interglacial global carbon cycle. However, there is a substantial lack of information from its Antarctic Zone south of the Polar Front (PF) to understand key climate processes (e.g., sea ice variability, productivity changes, CO2 source region, shifts of the Southern Westerly Wind) active in this region during the glacial/interglacial transition, due to the limited highresolution sediment records from this area. To close this gap, we investigated high resolution diatom records from a series of sediment cores from the Atlantic and Western Indian sectors of the Southern Ocean between the modern PF and the Winter Sea Ice (WSI) edge. Summer Sea Surface Temperature (SSST) and sea ice information spanning the past 30 thousand years were derived from diatom transfer functions and indicators, which augment comprehensive information on past surface ocean conditions and related ocean and atmospheric circulation, as well as opal deposition. These complementary lines of evidences also provide important environmental boundary conditions for climate simulations understanding the past climate development in the high latitudes Southern Ocean. Our reconstructions show that the Last Glacial (LG) SSSTs south of the modern PF are 1e3
C colder than modern conditions, WSI expanded to the modern PF. Our data suggests effective carbon export in the Antarctic Zone during the LG. Deglacial two steps of warming support the bipolar seesaw mechanism. Antarctic Zone is an important source region for the CO2 deglacial increase. The warming was more suppressed towards south, due to continuous ice discharge from Antarctica. The SSSTs exceeded modern values during the early Holocene optimum, when WSI extent probably retreated south of its modern position. The southern boundary of maximum opal deposition zone may have shifted to south of 55
S in the Bouvet Island area at this time. The mid-late Holocene cooling with WSI re-expanding to the Bouvet Island area, probably related to enhanced cold-water export by the Weddell Gyre from the developing cavity under the West Antarctic Ice Shelf. The cooling also suggests a northward shift of the Southern Westerly Wind, at least its southern boundary. © 2016 Elsevier Ltd. All rights reserved.

Keywords: Southern Ocean Diatoms Sea Surface Temperatures Sea ice Last Glacial Maximum Termination I Holocene

1. Introduction Well-dated paleoclimate records document Earth climate variability at a large range of time scales. They provide insights in (1) past climate dynamics including forcing, amplification and propagation physical and biological mechanisms, and (2) the range of climate variability under different climate boundary conditions. These information also provide an opportunity to test the ability of numerical models to simulate realistic climate development

* Corresponding author. State Key Laboratory of Marine Geology, Tongji University, Shanghai 200092, China. E-mail address: [email protected] (W. Xiao). http://dx.doi.org/10.1016/j.quascirev.2016.01.023 0277-3791/© 2016 Elsevier Ltd. All rights reserved.

beyond conditions described by instrumental data. The past ca. 30 ka (ka: thousand years) include several particularly interesting climatic end-members and transitions: (1) the Last Glacial Maximum (LGM), a cold climate end member marked by a sea level ca 120e140 m below present (Clark et al., 2009), (2) the LGMHolocene transition (Termination I), which documents a dramatic shift in climate conditions triggered by orbital forcing and punctuated by warm and/or cold conditions on both hemispheres that displays a non-synchronous pattern (EPICA members, 2006; Denton et al., 2010), and (3) the present interglacial (Holocene), which shows spatial and latitudinal variability in climate development (Jansen et al., 2007). The LGM-Holocene climate conditions were reconstructed by marine and land records primarily from the

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Northern Hemisphere and low latitude regions. These are complemented by a series of ice cores recovered from Greenland, Antarctica and low latitude glaciers with centennial to annual resolution (e.g., Thompson et al., 2000; NGICP members, 2004; EPICA members, 2006; WAIS Divide Project members, 2013). Corresponding information for the southern high latitude regions largely relies on Antarctic ice core records, which can be well dated and allow for accurate inter-hemispheric correlation (e.g., EPICA members, 2006; Lemieux-Dudon et al., 2010). Marine records from the Australian-New Zealand region as well as the Atlantic and Pacific sectors of the Southern Ocean show a similar climate pattern to Antarctic ice cores but mainly rely on records from areas north of the modern Polar Front (PF) (e.g., Kaiser et al., 2005; Barker et al., 2009; Bostock et al., 2013). The Southern Ocean represents the largest water body of the world oceans, and by connecting the Atlantic, Indian and Pacific Oceans, is a crucial conduit for the redistribution of heat and nutrients between the various ocean basins. During glacial periods, the Southern Ocean is considered to be an important carbon sink, not only in the Subantarctic Zone (SAZ) but also in the seasonal sea ice zone (SIZ) to the south, due to enhanced bioproductivity and strong stratification (Martínez-García et al., 2014; Abelmann et al., 2015). During the deglaciation, these regions become carbon sources, due to wind driven upwelling (Anderson et al., 2009). In the Antarctic Zone south of the Polar Front, the variability of seasonal sea ice strongly affects the albedo as well as the ocean/atmosphere exchange of heat and gas, thermohaline circulation and primary productivity (Thomas and Dieckmann, 2010). The variability in sea ice cover/extent also changes the sea surface water temperature gradient that has a direct impact on the overlying atmospheric circulation (e.g., Southern Westerly Wind, SWW, Kohfeld et al., 2013), which in turn modulates the position of the upwelling region and related opal deposition and CO2 degassing (Anderson et al., 2009). In the Atlantic sector, the sea ice distribution is strongly affected by the cold Weddell Gyre (Jullion et al., 2014), which represents an interface between the West Antarctic Ice Sheet (WAIS) and the open ocean South Atlantic. The generation of cold surface and bottom waters in the Weddell Sea is connected to the presence of large sub-ice cavities beneath the shelf ice (Nicholls et al., 2009). Thus, the reconstruction of cold water expansion history helps understanding the ice-ocean interaction under past climate conditions. Antarctic ice core sea salt sodium (þNass) records give general sea ice areal coverage information of the Southern Ocean, but cannot determine the actual sea ice € thlisberger et al., 2010). Model extension in a specific area (Ro simulations show difficulties in reproducing the LGM Southern Ocean sea ice distribution either in extent or seasonality, which calls for robust proxy data to understand the sea ice dynamics in the Southern Ocean (Roche et al., 2012; Goosse et al., 2013). Given the important role of the Southern Ocean in many climate-relevant processes, several key questions arise, including (1) how does the high latitude Southern Ocean climate development fit into the context of global deglaciation mechanisms; (2) what is the spatial/temporal pattern of sea ice extent in the South Atlantic Antarctic Zone; (3) where is the upwelling region for opal production located, and what is its impact on the carbon cycle; and (4) what is the Weddell Gyre cold water expansion history and what are the implications for its interaction with the WAIS. Despite the climatic significance of the Antarctic Zone in the Atlantic sector, a limited number of records is available from this region. These include a few records from areas adjacent to Bouvet Island (e.g., Bianchi and Gersonde, 2004; Anderson et al., 2009; Divine et al., 2010), the Scotia Sea (e.g., Allen et al., 2005, 2011; Collins et al., 2012, 2013) and the Antarctic Peninsula region (e.g., Shevenell et al., 2011; Pike et al., 2013). This information is however not

sufficient to infer the spatial variability of Sea Surface Temperatures (SSTs) and sea ice distribution. Due to the poor preservation of carbonate in the Antarctic Zone (Diekmann, 2007), little paleoceanographic information can be obtained by assemblage analysis or geochemical proxies based on calcareous microfossils, as widely applied in low latitude oceans. In contrast, diatoms represent the dominant biogenic components in Antarctic sediments, and thus have great potential for deciphering past ocean conditions (Crosta et al., 1998a; Zielinski et al., 1998; Esper and Gersonde, 2014a,b). In this paper we augment the knowledge on Summer Sea Surface Temperature (SSST), sea ice extent, and Winter Sea Ice (WSI) concentration in the high latitude Southern Ocean during the past 30 ka, using diatom-based transfer function reconstructions (Esper and Gersonde, 2014a,b) with 12 sediment cores recovered from the present Permanent Open Ocean Zone (POOZ) between the PF and the WSI edge, between 40
480 E (Western Indian sector) and 44
060 W (Scotia Sea, Western Atlantic sector), an area sensitive to cold water expansion from the Weddell Gyre, and, potentially, northward movements of the WSI edge during glacial times (Fig. 1). The new SSST and sea ice records are compared with published records of a similar nature to generate a spatial picture across the Atlantic sector of the Southern Ocean. A comparison with climate records obtained from other Southern Ocean sectors, as well as continental ice cores will provide a broader, global context for southern high-latitude Last Glacial (LG) to Holocene climate variability. Our new records will provide valuable information to address the questions mentioned above, and support the proper development and testing of numerical simulations of past climate conditions and related modulation and forcing mechanisms, e.g., for the present interglacial (e.g., Renssen et al., 2005, 2010; Brovkin et al., 2008; Elsig et al., 2009) and the last glacial/interglacial transition (Termination I) (e.g., Stocker, 2003; € hler et al., 2011). Ko 2. Study area The Southern Ocean circulation is characterized by the largescale Antarctic Circumpolar Current (ACC), which is subdivided by a series of oceanic fronts (Fig. 1). From north to south, they are the Subtropical Front (STF), the Subantarctic Front (SAF), the Polar Front (PF), and the Southern ACC Front (Orsi et al., 1995). More recent physical oceanographic studies indicate that the Southern Ocean fronts have more complex structures, with a series of branches (northern-, mean-, and southern branches) regarded as “sub-front” (Sokolov and Rintoul, 2009). However, on the geological time scale, oceanographic and climatic signals would be most strongly affected by the mean positions of the fronts. The PF represents the most important frontal system in the ACC, as it separates the cold and fresh Antarctic surface water from the relatively warm and salty subAntarctic water (Orsi et al., 1995; Sokolov and Rintoul, 2009). The PF also forms an ecological and sedimentary water-mass boundary, with calcareous oozes predominate north of the PF, whereas diatomaceous oozes form the Circum-Antarctic Opal Belt to the south between the PF and WSI edge (Burckle and Cirilli, 1987; Diekmann, 2007). South of Africa, warm surface water is imported from the Indian Ocean to the South Atlantic Ocean by the Agulhas Current through the so-called Warm Water Route (WWR), which modulates the heat budget of the South Atlantic (Seidov and Maslin, 2001). In the south, the Weddell Gyre exports cold Weddell Sea water to the open ocean. The generation of cold Weddell Sea water masses is linked to the presence of cavities beneath the Filchner-Ronne ice shelf and sea ice in front of the ice shelf itself (Nicholls et al., 2009). This process is crucial for the extent of the cold surface water expansion and the related sea ice field in the Southern Ocean Atlantic sector (Comiso, 2003), and also impacts on

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Fig. 1. Locations of studied and discussed core sites (Table 1). New records are marked with white circles and bold numbers. WSI/SSI: modern average winter/summer sea ice edge (Comiso, 2003); locations of oceanic fronts are according to Orsi et al. (1995). ACC: Antarctic Circumpolar Current; PF: Polar Front; SAF: Subantarctic Front; STF: Subtropical Front; AC: Agulhas Current; BC: Brazil Current; MC: Malvinas Current; BMC: Brazil-Malvinas Confluence; SWW: Southern Westerly Wind.

the production of Antarctic Bottom Water (AABW) (Nicholls et al., 2009). Below the cold surface layer, the Circumpolar Deep Water (CDW) shoals towards the south (Orsi et al., 1995), replenishing nutrients and heat to the surface ocean. Atmospheric circulation overlying the Southern Ocean is dominated by the Southern Westerly Wind (SWW), which represents an important driver of the large scale Southern Ocean circulation, and related heat, salt, and nutrient transport (Wunsch, 1998). Today the position of SWW varies seasonally, with more northward migration during winter and southward situated during summer. Its core zone is at ca. 50e55
S (Schneider et al., 2003). The SWW induced upwelling supplies nutrient (including dissolved silicon) to the surface ocean, which favors diatom productivity and the subsequent opal deposition (Anderson et al., 2009). The upwelling also potentially releases CO2 stored in abyssal ocean to the atmosphere (Toggweiler et al., 2006; Denton et al., 2010). Thus, the location and strength of SWW have big impact on Southern Ocean opal deposition and global carbon cycle.

PS1651-1, PS1652-2 (ANT-VI/3, 1988, Fütterer, 1988), PS2102-2 (ANT IX/4, 1991, Bathmann et al., 1992); and ODP1093/PS1654-2 and ODP1094/PS2090-1 (Bianchi and Gersonde, 2004) collected close to Bouvet Island; and piston core PS2606-6 (ANT-XI/4, 1994) collected at the Conrad Rise in the Western Indian Ocean (Jacot Des Combes et al., 2008). The sediment cores were generally sampled and analyzed for their diatom content at a spacing of 10 cm. Exceptions are gravity and box cores PS1649-2 and PS1649-1, respectively, studied at 5 cm intervals. For Core PS2102-2, previously investigated by van Beek et al. (2002), we have increased the resolution of studied samples to 5 cm in the Holocene section and to 2.5 cm in the LG to Termination I interval. The down-core diatom assemblage variation in cores ODP1093/PS1654-2, ODP1094/PS2090-1, and PS2606-6 were previously presented by Bianchi and Gersonde (2004), and Jacot Des Combes et al. (2008), respectively. Here we present new information on SSST, WSI and SSI, using the newly developed diatom based transfer functions (Esper and Gersonde, 2014a,b).

3. Materials and methods

3.2. Sample preparation and diatom census

3.1. Materials

The sample treatment and the preparation of permanent mounts for light microscopy followed the standard procedure established at AWI-Bremerhaven (Gersonde and Zielinski, 2000). Counts on diatom species were carried out on the permanent mounted slides of acid-cleaned materials. Following the taxonomy described by Zielinski and Gersonde (1997), the diatom taxa were counted at the magnification of 1000 with a Zeiss Laborlux microscope with apochromatic optics. The counting procedure and definition of counting units followed those of Schrader and Gersonde (1978). A minimum of 400 specimens was counted in each sample. To assess the dissolution bias on the original diatom assemblage, three preservation states (good, moderate, and poor)

The high-resolution LG-Holocene records investigated in this study (Table 1, Fig. 1) are located in the present POOZ of the Atlantic and Western Indian sectors of the Southern Ocean, between the modern WSI edge and the PF. This region is characterized by intensive biosiliceous productivity, mainly composed of diatoms, and regarded as the “circum-Antarctic opal belt” (Diekmann, 2007). The studied cores includes PS67/197-1 and PS67/219-1 recovered during R/V Polarstern cruises ANT-XXII/4 (2005) from the Scotia Sea (Schenke and Zenk, 2006); spliced sediment sequences from gravity and box corer deployments at site PS1649; gravity cores

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Table 1 Coordinates and water depths/elevation, coring devices and data source of studied and discussed cores. Station

Latitude

Longitude

Water depth (m)

Coring device

Data source

Marine cores PS1649-1 PS1649-2 PS1651-1 PS1652-2 PS2102-2 PS67/197-1 PS67/219-1 PS2606-6 ODP1093 PS1654-2 TN057-17PC1 ODP1094 PS2090-1 TN057-13PC4 PS1768-8 ODP1089 PS2821-1 TN057-21 ODP1090 PS2489-2 PS2495-3 PS2498-1 MD07-3076 KC73 TPC36 TPC63 TPC78 TPC286 MD07-3133 MD07-3134 ODP1233 GeoB3313-1 GeoB6211 GeoB13862-1 GeoB6308 ODP1098 JPC 10 GGC5

54
54
53
53
53
55
57
53
49
50
50
53
53
53
52
40
40
41
42
42
41
44
44
52
52
53
55
61
57
59
41
41
32
38
39
64
64
33


54.880 S 54.630 S 37.850 S 39.840 S 4.380 S 8.240 S 13.220 S 13.90 S 58.580 S 9.50 S S 10.80 S 10.70 S 120 S 35.580 S 56.180 S 56.580 S 80 S 52.40 S 52.40 S 16.50 S 9.20 S 4.460 S 9.20 S 36.20 S 55.80 S 330 S 41.40 S 260 S 250 S 0.010 S S 30.310 S 1.110 S 18.10 S 51.1620 S 530 S 420 N

3
17.180 E 3
18.460 E 3
51.370 E 5
5.970 E 4
59.140 W 44
6.280 W 42
28.020 W 40
48.10 E 5
51.920 E 5
43.30 E 6
E 5
7.80 E 5
7.980 E 5
60 E 4
28.560 E 9
53.640 E 9
53.280 E 7
490 E 8
55.10 E 8
58.40 E 14
29.40 W 14
13.70 W 14
12.470 W 41
10.70 W 46
52.70 W 48
2.40 W 45
1.20 W 40
8.40 W 43
270 W 41
280 W 74
26.990 W 74
270 W 50
14.560 W 53
47.70 W 55
57.90 W 64
12.480 W 64
120 W 57
350 W

2446 2427 2075 1963 2390 3837 3619 2545 3624 3744 3700 2807 2819 2800 3299 4620 4575 4981 3699 3794 3134 3783 3770 3760 3553 3956 3840 3467 3101 3663 838 852 657 3588 3620 1010 905 4550

BC GC GC GC GC PC PC PC OD PC PC OD PC PC GC OD PC PC OD PC GC GC PC KC PC PC PC PC PC PC OD GC GC GC GC OD PC PC

this study this study this study this study this study; van Beek et al., 2002 this study; Xiao et al., 2016 this study; Xiao et al., 2016 this study; Jacot Des Combes et al., 2008 this study; Bianchi and Gersonde, 2004 this study; Bianchi and Gersonde, 2004 Nielsen et al., 2004; Divine et al., 2010 this study; Bianchi and Gersonde, 2004 this study; Bianchi and Gersonde, 2004 Anderson et al., 2009; Divine et al., 2010 Zielinski et al., 1998; Esper and Gersonde, 2014a,b Cortese et al., 2007 Cortese et al., 2007 Barker et al., 2009, 2010 Martínez-García et al., 2009 Martínez-García et al., 2009 Gersonde et al., 2003 Gersonde et al., 2003 Skinner et al., 2010 Allen et al., 2005 Allen et al., 2011 Collins et al., 2012, 2013 Collins et al., 2012, 2013 Collins et al., 2012, 2013 Weber et al., 2012, 2014 Weber et al., 2012, 2014 Kaiser et al., 2005; Lamy et al., 2007 Lamy et al., 2002 Voigt et al., 2015 Voigt et al., 2015 Voigt et al., 2015 Domack et al., 2001; Shevenell et al., 2011; Pike et al., 2013 Etourneau et al., 2013 McManus et al., 2004

75
75
79
80
64
75


60 S S 28.060 S 0.60 S 12.100 S 60 N

123
210 E 0
112
5.190 W 119
30.60 W 57
41.100 W 42
19.20 W

Ice cores EDC EDML WDC Byrd JRI NGRIP

Elevation (m) 3233 2891 1766 1530 1542 2917

Jouzel et al., 2007; Monnin et al., 2001, 2004; Lambert et al., 2008 EPICA members, 2006, Fischer et al., 2007 WAIS Divide Project members, 2013 Ahn and Brook, 2008 Mulvaney et al., 2012 NGICP members, 2004

Coring devices: BC-box corer; GC-gravity corer; PC-piston corer; KC-Kasten corer; OD-ocean drilling.

of the diatom samples were estimated according to Esper et al. (2010), based on the relationship of weakly and heavily silicified species, the dissolution stage of the valves, and the quantification of fragmentation. 3.3. Summer Sea Surface Temperature and sea ice estimation The SSSTs were estimated using the Imbrie and Kipp transfer function method (IKM). The IKM is an application of principal component regression to generate a single calibration formula between diatom counts and the environmental variables (Imbrie and Kipp, 1971). IKM is an effective method for diatom based SSST reconstruction in the Southern Ocean due to the strong relationship between diatom assemblage and SSST, analyzed by statistic methods of Detrended Correspondence Analysis and Canonical Correspondence Analysis (Esper and Gersonde, 2014a). The SSSTs of cores PS1768-8, ODP1093/PS1654-2, and ODP1094/ PS2090-1 were previously estimated using the IKM calibrated with a diatom reference data set of 93 surface samples from the Atlantic and Western Indian sectors of the Southern Ocean

(Zielinski et al., 1998; Bianchi and Gersonde, 2004). The SSSTs of these cores were recalculated, together with the SSST estimation of other cores studied in this paper, using the updated reference data set and calculation (Esper and Gersonde, 2014a). The new reference data set (D336/29/3q) includes diatom assemblage data obtained from 336 sites from the Southern Ocean, considering the abundance pattern of 29 diatom taxa or taxa groups, and using 3 factor mode analysis. This data set deals well with high assemblage variability in SSST estimation (Esper and Gersonde, 2014a). The referenced hydrographic dataset for SSST calibration represents values measured at 10 m water depth as presented by Olbers et al. (1992), a dataset not strongly affected by Southern Ocean warming during the last decades. The estimation of past surface ocean temperature is restricted to summer. This considers the results from sediment trap experiments, which showed significant diatom flux to the sea floor is restricted to austral summer in the Southern Ocean (Abelmann and Gersonde, 1991; Fischer et al., 2002). Log transformation of the diatom abundance data was applied in order to reduce the dominance of Fragilariopsis kerguelensis and enhance the weight of less abundant taxa in calculation. The standard error

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of the estimates' regression is 0.86
C. A software package by Sieger et al. (1999) was used for the calculations. The calculated communalities quantify how well each assemblage fits within the simplified “three factors” model on a scale between 0 and 1. They generally vary between 0.8 and 0.9, suggesting high accuracy and reliability of the estimation according to the reference data set. Estimation of the WSI extent followed the method proposed by Gersonde and Zielinski (2000): relative abundance >3% of the two diatom sea-ice indicator species, Fragilariopsis curta and Fragilariopsis cylindrus, combined as F. curta group, indicate the presence of WSI. Relative abundances >3% of Fragilariopsis obliquecostata, a species restricted to waters colder than 1
C (Zielinski and Gersonde, 1997), along with a distinct drop in biogenic sedimentation rate, is interpreted to indicate proximity to the summer sea ice (SSI) edge (Gersonde and Zielinski, 2000). The past WSI concentration (WSIC) was calculated by the Modern Analog Technique (MAT) transfer function. This method searches for the best analogs based on similarity indexes between the surface sediment reference and the down-core assemblage, and uses the average for the modern conditions observed at these best analog sites as an estimate for past environmental conditions, dealing best with the nonlinear distribution of sea ice. The MAT calculation is based on a 274 surface sediment sample diatom data set from the Atlantic, Pacific and Western Indian sectors of the Southern Ocean (D274/28/4an), considering the abundance pattern of 28 diatom taxa or taxa groups and calculated with 4 analogs, and showing high fidelity (Root Mean Square Error of Prediction 5.52%) in estimating the nonlinear distributed Southern Ocean WSIC (Esper and Gersonde, 2014b). The SSI concentration was not calculated due to the relatively large statistical errors in the estimation, resulting from poor opal preservation under the perennial sea ice cover (Esper and Gersonde, 2014b). 4. Stratigraphy The accurate dating of Southern Ocean sediment records represents a major challenge but is required in order to put them into a detailed time frame together with other records obtained from marine, ice and land archives. For the dating of the Holocene and LG records, the radiocarbon method calibrated for the past 50,000 years (Fairbanks et al., 2005; Reimer et al., 2013) is the most commonly used method. However, evidence suggests that besides differences in regional reservoir effect, reservoir shifts have occurred in the past (Hughen et al., 2004; 2006), as a result changes in ocean circulation and ventilation, and introduces a time-variable factor in the conversion between 14C and calendar ages. This holds particularly true for the Southern Ocean where strong reservoir effect may exceed 1000 years (Bard, 1988) and the termination and glacial reservoir corrections may be much larger than today (Sikes et al., 2000; Siani et al., 2013). Additionally, in Southern Ocean sediments, AMS 14C dating and oxygen isotope measurements based on calcareous hard-parts of foraminifers is hampered by the scarcity or lack of biogenic carbonate. To generate age models of highest possible accuracy and consistency, we combined 75 available and 58 new AMS 14C data from 10 cores derived from planktonic foraminifers and organic carbon in bulk samples (Table 2), together with available oxygen and carbon isotope records and information derived from the regional pattern of magnetic susceptibility, diatom species abundance. The transfer function derived SSST records were also tentatively used for regional correlation among closely spaced cores. 4.1. AMS

14

C chronology

New AMS

14

C ages dated on organic carbon from both humic

119

acid and residue fractions were obtained from the bulk sediment samples of cores PS1649-1/PS1649-2 and PS1652-2 at the Leibniz Labor für Altersbestimmung und Isotopenforschung in Kiel (KIA), Germany. Similar analysis was conducted on sediments of cores PS67/197-1 and PS67/219-1 at the Poznan Radiocarbon Laboratory in Poland (Poz) (Xiao et al., 2016). The datings on the organic carbon humic acid fraction for Core PS2606-6 were partly presented in Jacot Des Combes et al. (2008), whereas additional datings are presented here. The AMS 14C ages of other cores are summarized from previous studies (Table 2). The measured residue ages are less than 100 to more than 2000 years older than the humic acid ages (Table 2). Generally, in the Bouvet Island area, the residue ages are comparable to the humic acid ages during the Holocene but differ largely during the LG and the Termination 1. This suggests that, the residue fraction is contaminated by re-deposition of old stable organic carbon of probable terrigeneous origin; and the contamination is less in the Holocene than in the LG and the deglaciation. This is in agreement with the higher diatom ooze deposition in the Holocene sections and the presence of more terrigeneous material in the older sections of the cores. The foraminiferal 14C ages in the western Bouvet Island area Core PS2102-2 are similar to the humic acid 14C ages in the eastern cores (PS1652-2, PS1649-1/PS1649-2 and PS2090-2/ODP1094) at intervals with similar SSST and diatom species fluctuation pattern. In cores PS1654-2/ODP1093, the foraminiferal ages are similar to humic acid ages at close core depth intervals (Table 2). This indicates that ages dated on both materials are comparable, and the same reservoir age can be applied. Therefore, the age models of cores from Bouvet Island area are based on the foraminifera and humic acid ages. The exception to this is in one sample (PS1654-2; 1068e1072 cm), where the humic acid age is significantly older than the residue age. The younger age was conservatively taken for the core age model (Bianchi and Gersonde, 2004). The AMS 14C dates of Core PS2102-2 were obtained from two laboratories (van Beek et al., 2002) (Table 2). We note that the AAR (Laboratory Department of Physics and Astronomy University of Aarhus) ages are systematically ca. 300 years younger than the KIA ages. The AAR age at 300e302.5 cm constrains the start of the early Holocene optimum at the core site (Table 2; Suppl. Fig. 1g), and results in the extremely elevated sedimentation rate for this interval (Fig. 2g). This age could also be affected by an inter-lab systematic error, and subsequently the calibrated calendar age could be biased towards a younger age by ca. 300 years. However, in comparison with other dates in the adjacent cores (PS1649-2, PS1654-2), this age is still reasonable, thus is kept for the core stratigraphy. One date at the start of the early Holocene optimum in Core PS1652-2 (627e630 cm) is slightly younger than in the other cores. Here we conservatively transfer 2 dates from core PS1649-2, marking the early Holocene optimum, into the correspondent intervals in Core PS1652-2 by using the diatom assemblage fluctuation and SSST correlation. For Core PS2606-6, the humic acid ages are generally 2e3 ka younger than the residue ages in the Holocene and 1e2 ka younger in the LG. However, in contrast to the Bouvet Island area cores, residue ages were used to develop the stratigraphy of Core PS26066. The correlation of diatom species and SSST fluctuation patterns to the Bouvet Island area cores (especially to cores PS2090-1/ ODP1094 at similar latitudes and relative location compared to the PF) show better consistency with the residue ages of Core PS2606-6 to the humic acid ages in the Bouvet Island area cores. This is also supported by the matching of the residue ages and the foraminifera age of this core (Table 2). The reason for the significantly younger humic acid ages in this core is yet unclear. Conventional 14C ages were converted to calendar ages using the

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Table 2 Determination of AMS 14C ages of studied cores, together with other datings from previous studies in the area. Values used to establish the age models are marked bold. The calendar ages are calibrated using the program CALIB 7.0.2 with Marine 13 calibration curves (Stuiver and Reimer, 1993; Reimer et al., 2013). A constant reservoir correction of 1100 years (DR ¼ 700 years) was applied in the Bouvet Island area and Western Indian sector cores, and 1300 years (DR ¼ 900 years) in the Scotia Sea cores. Carbon sources for dating are: Reresidue fraction; Hehumic acid fraction; Feplanktonic foraminifera N. pachydermasin. Core

Depth (cm)

Lab id

C Source

14

PS1649-1

2e5 2e5 31e34 31e34 37e40 37e40 56e59 56e59 84e87 84e87 121e124 121e124 151e154 151e154 6.5e10 6.5e10 198e202 198e202 315e319 315e319 406.5e410.5 406.5e410.5 406.5e410.5 406.5e410.5 486.5e490 486.5e490 627e630 627e630 711e715 711e715 762e766 762e766 801e805 801e805 22e27 95e100 95e100 316e321 316e321 441e445 441e445 460e464 533e538 538e542 538e542 621e626 621e626 627e632 748e752 748e752 905e909 905e909 1009e1013 1009e1013 1068e1072 1068e1072 8e9.5 53e55 77e79 141.5e143.5 0e4 0e4 48e52 48e52 249e253 249e253 460e464 460e464 550e554 550e554

KIA 36296 KIA 36296 KIA 36297 KIA 36297 KIA 36298 KIA 36298 KIA 36299 KIA 36299 KIA 36300 KIA 36300 KIA 36301 KIA 36301 KIA 36302 KIA 36302 KIA 36303 KIA 36303 KIA 36304 KIA 36304 KIA 36305 KIA 36305 KIA 36306 KIA 36306a KIA 36306b KIA 36306* KIA 36307 KIA 36307 KIA 36308 KIA 36308 KIA 36309 KIA 36309 KIA 36310 KIA 36310 KIA 36311 KIA 36311 KIA14938 KIA15510 KIA15510 KIA15511 KIA15511 KIA15512 KIA15512 KIA15514 KIA15515 KIA14941 KIA14941 KIA14945 KIA14945 KIA14939 KIA14946 KIA14946 KIA14942 KIA14942 KIA14943 KIA14943 KIA14944 KIA14944 AAR-1535 AAR-1536 AAR-1537 AAR-1538 KIA14947 KIA14947 KIA14948 KIA14948 KIA14949 KIA14949 KIA14950 KIA14950 KIA14955 KIA14955

R H R H R H R H R H R H R H R H R H R H R H H H R H R H R H R H R H F R H R H R H F F R H R H F R H R H R H R H F F F F R H R H R H R H R H

4495 2500 8775 7445 9735 10,100 10,560 10,415 13,940 13,350 19,320 17,900 19,070 19,690 1855 1815 4640 4415 6335 6255 8285 7955 8265 8110 8875 8835 10,215 10,055 22,260 20,120 20,360 18,310 26,310 17,850 3930 8120 7320 9525 8930 10,200 9820 10,270 10,530 10,460 10,450 12,730 10,410 10,920 11,790 11,670 12,970 13,070 13,440 13,700 15,860 17,320 2850 9540 10,510 12,850 2580 1890 3955 3355 7660 7240 13,110 9040 10,755 9960

PS1649-2

PS1652-2

PS1654-2

PS1768-8

PS2090-1

C age (a)

Error (a)

Lower cal range (a BP)

Upper cal range (a BP)

Median probability (a BP)

±40 þ30/¡25 ±50 ±40 ±50 ±55 ±60 þ55/¡50 ±80 ±55 þ210/200 ±200 þ280/270 þ310/300 ±30 ±30 ±25 ±35 ±35 ±35 ±35 ±35 ±40 ±40 ±45 ±45 þ55/50 ±55 þ390/370 ±140 þ280/270 ±80 þ1410/1200 ±130 ±30 ±50 ±50 þ60/55 ±60 ±70 ±70 ±60 ±55 ±70 ±70 ±60 ±80 ±50 ±70 þ80/¡70 ±80 ±90 ±80 þ120/¡110 ±120 ±230 ±75 ±110 ±140 ±150 ±30 ±35 ±45 ±50 ±45 ±50 ±70 ±80 ±60 ±80

3617 1272 8400 7175 9545 10,138 10,632 10,474 15,083 13,954 21,538 19,739 21,016 21,709 647 621 3832 3533 5930 5863 7957 7632 7936 7779 8532 8468 10,225 10,067 24,430 22,521 22,528 20,523 26,436 19,871 2939 7767 7002 9368 8581 10,200 9630 10,250 10,595 10,493 10,482 13,328 10,387 11,099 12,498 12,216 13,503 13,610 14,041 14,274 17,645 18,980 1550 9251 10,400 13,310 1342 658 2936 2176 7391 6914 13,713 8628 10,850 9793

3854 1411 8661 7373 9892 10,398 11,020 10,775 15,624 14,430 22,487 20,755 22,400 23,189 765 730 4022 3762 6154 6050 8141 7811 8131 7958 8854 8765 10,516 10,365 26,125 23,316 23,846 20,995 31,892 20,545 3152 7983 7255 9623 8955 10,530 10,124 10,582 10,969 10,905 10,892 13,645 10,857 11,334 12,743 12,676 13,935 14,074 14,800 15,276 18,272 20,102 1911 9824 11,118 13,936 1514 818 3211 2498 7556 7182 14,078 9120 11,198 10,241

3740 1338 8530 7275 9721 10,239 10,804 10,616 15,313 14,150 22,062 20,263 21,740 22,450 699 670 3924 3634 6036 5939 8041 7715 8020 7876 8664 8606 10,368 10,194 25,409 22,892 23,197 20,750 29,327 20,207 3043 7882 7150 9484 8754 10,355 9870 10,437 10,760 10,670 10,658 13,471 10,611 11,209 12,620 12,510 13,730 13,846 14,364 14,876 17,956 19,565 1738 9507 10,748 13,611 1428 729 3077 2348 7464 7056 13,891 8897 11,055 10,061

226

Ra age (a BP)

Reference 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 2 2 2 2 2 2 2 2 2 2 2 2 2 2 2 2 2 2 2 2 2 2 1 3 3 3 2 2 2 2 2 2 2 2 2 2

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121

Table 2 (continued ) Core

PS2102-2

PS67/197-1

PS67/219-1

PS2606-6

Depth (cm)

Lab id

C Source

14

675e679 675e679 720e724 720e724 766e770 766e770 826e830 826e830 826e830 826e830 20e23 26.5e27.5 70.5e71.5 72e75 207.5e210 210.5e211.5 266.5 300e302.5 0e4 0e4 270e273 270e273 470e473 470e473 542e545 542e545 1e5 1e5 240e243 240e243 275e278 275e278 541e544 541e544 624e627 624e627 12e16 12e16 12e16 126e130 126e130 126e130 126e130 197e200 226e230 226e230 226e230 226e230 273e277 273e277 304e308 304e308 304e308 316e320 316e320 355e359 355e359 355e359 355e359 425e429 425e429 516e520 516e520

KIA14951 KIA14951 KIA14956 KIA14956 KIA14952 KIA14952 KIA14953 KIA14953a KIA14953b KIA14953* KIA29590 KIA9300 KIA9299 AAR-1530 AAR-1532 KIA10019 KIA9298 AAR-1533 Poz-32738 Poz-32667 Poz-32740 Poz-32669 Poz-32849 Poz-32671 Poz-32743 Poz-32693 Poz-32873 Poz-32807 Poz-32814 Poz-32804 Poz-32855 Poz-32853 Poz-32863 Poz-32862 Poz-32865 Poz-32864 KIA 16758a KIA 16758b KIA 16758* KIA 16759a KIA 16759b KIA 16759* KIA 16759 KIA 23807 KIA 16760a KIA 16760b KIA 16760* KIA 16760 KIA 16761 KIA 16761 KIA 16762a KIA 16762b KIA 16762* KIA 23790 KIA 23790 KIA 16763a KIA 16763b KIA 16763* KIA 16763 KIA 16764 KIA 16764 KIA 16765 KIA 16765

R H R H R H R H H H F F F F F F F F R H R H R H R H R H R H R H R H R H R R R R R R H F R R R H R H R R R R H R R R H R H R H

11,890 11,130 15,355 11,700 13,740 12,660 19,890 15,750 16,050 15,900 1890 2095 3345 3050 6650 7110 10,150 10,440 7260 4470 16,720 13,350 17,700 16,020 20,580 17,370 6520 5730 14,740 12,590 10,700 13,500 20,550 17,250 23,420 20,930 2920 3100 3010 8090 8535 8315 6210 9975 10,180 10,270 10,225 7110 11,410 7700 12,460 12,175 12,320 12,970 9860 13,715 13,330 13,525 12,960 17,400 16,650 21,860 20,340

* 14C ages obtained by averaging values from two analyses. Reference. 1. This study. 2. Bianchi and Gersonde, 2004. 3. Gersonde et al., 2003. 4. van Beek et al., 2002. 5. Xiao et al., 2016. 6. Jacot Des Combes et al., 2008.

C age (a)

Error (a)

Lower cal range (a BP)

Upper cal range (a BP)

Median probability (a BP)

±80 ±120 ±55 ±100 ±70 ±110 þ120/110 þ190/180 ±110 þ190/¡180 ±30 ±40 ±60 ±75 ±70 ±90 ±60 ±140 ±40 ±40 ±80 ±110 ±140 ±160 ±120 ±130 50 50 80 70 70 110 360 250 270 360 ±45 ±25 ±45 ±90 ±50 ±70 ±130 ±130 ±110 ±50 ±110 ±110 ±70 þ170/160 ±70 ±55 ±70 ±90 ±60 þ69/55 ±60 þ69/¡60 ±130 þ90/¡80 ±100 þ130/120 ±210

12,560 11,200 17,139 12,186 14,677 13,204 22,408 17,361 17,896 17,546 661 845 2144 1783 6220 6680 10,177 10,281 6715 3360 18,506 13,652 19,449 17,522 22,894 19,008 5905 4963 15,896 12,993 10,545 13,814 22,422 18,723 25,876 22,738 1698 1945 1809 7657 8181 7928 5606 9664 10,114 10,269 10,162 6636 11,844 7174 13,097 12,772 12,906 13,489 9718 14,579 13,918 14,162 13,428 19,442 18,589 24,539 22,614

12,846 11,996 17,546 12,719 15,267 13,674 22,948 18,310 18,445 18,458 802 1048 2519 2172 6526 7135 10,467 11,035 6941 3564 18,848 14,190 20,144 18,308 23,555 19,711 6160 5277 16,434 13,332 10,967 14,640 24,006 19,868 27,062 24,454 1934 2117 2052 8019 8389 8233 6183 10,413 10,621 10,570 10,648 7160 12,447 7819 13,385 13,106 13,274 13,952 10,143 15,238 14,383 14,943 14,016 19,965 18,996 25,387 23,682

12,686 11,571 17,357 12,526 15,009 13,416 22,652 17,821 18,169 18,006 727 943 2336 1975 6365 6893 10,300 10,655 6828 3462 18,685 13,931 19,787 17,906 23,220 19,378 6020 5128 16,166 13,179 10,722 14,117 23,190 19,253 26,380 23,621 1821 2033 1924 7849 8302 8075 5895 10,060 10,344 10,440 10,389 6895 12,124 7495 13,241 12,947 13,105 13,729 9945 14,964 14,125 14,524 13,718 19,678 18,801 25,016 23,174

226

Ra age (a BP)

855 2255 2335 6630 6700

± ± ± ± ±

35 75 75 220 220

9560 ± 320

Reference 2 2 2 2 2 2 2 2 2 2 1 4 4 4 4 4 1 4 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 1 1 1 1 1 1 6 1 1 1 1 6 1 6 1 1 1 1 6 1 1 1 1 1 6 1 6

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program CALIB 7.0.2 with calibration curve Marine 13 (Stuiver and Reimer, 1993; Reimer et al., 2013). In previous studies, the carbon reservoir effects applied in the calibration were according to Bard (1988) based on latitudinal variations (Gersonde et al., 2003; Bianchi and Gersonde, 2004; Jacot Des Combes et al., 2008). A regional reservoir for mid-late Holocene of ca. 1100 yrs has been derived in Core PS2102-2 by comparison between AMS 14C dating on planktonic foraminifera (N. pachydermasin.) and 226Ra-in-barite dating (van Beek et al., 2002). Such comparison also indicates a significantly high reservoir age (ca. 1900 yrs) during the early Holocene, which can be an overestimate resulting from changes in sedimentation rates and large errors in 226Ra measurements for the early Holocene interval (van Beek et al., 2002). Although the carbon reservoir may be variable through the LG to Holocene (Skinner et al., 2010), we tentatively applied a constant reservoir effect of 1100 yrs (DR ¼ 700 yrs) in the cores from Bouvet Island area and the Western Indian sector. This may result in slightly older estimation by 14C calibration during the LG. For cores from the Scotia Sea, we applied the widely accepted reservoir age of 1300 yrs (DR ¼ 900 yrs) for this area (Domack et al., 2001; Pugh et al., 2009), according to studies on living organisms (Gordon and Harkness, 1992; Berkman and Forman, 1996) (Table 2). 4.2. Regional core correlation The individual AMS 14C core chronologies represent the basic chronological framework for each core. More detailed core stratigraphies were established by regional correlation (Suppl. Fig. 1), enabling us to transfer age control points within cores and construct detailed regional age models (Fig. 2). Based on the assumption that the climate signals preserved in the geological records show similar changes in areas with similar oceanographic conditions (frontal system, sea ice coverage, etc.), a series of parameters were used for regional core correlation. These include diatom species composition, estimated SSSTs, and magnetic susceptibility (MS). The MS records presented in this study are taken from the PANGAEA database (http://www.pangaea.de). Besides the AMS 14C chronology, the stratigraphies of cores PS67/197-1 and PS67/219-1 are also based on their MS correlation to Antarctic ice core dust records (Suppl. Fig. 2; Xiao et al., 2016), as this method has been proven to be reliable in Scotia Sea sediment stratigraphy (Pugh et al., 2009; Weber et al., 2012; Xiao et al., 2016). The time series of the Scotia Sea cores represent the highest time resolution for the LG interval amongst the records analyzed in this paper. Due to the lack of glacial 14C dates, detailed LG stratigraphies of the Bouvet Island area and Western Indian sector cores were made by correlation to Core PS67/197-1 diatom composition fluctuations. 5. Results 5.1. Sedimentation rates Biogenic opal is a major component for Holocene sediments south of the modern PF in the circum-polar opal belt, and to a lesser extent for LG sediments (Diekmann, 2007). Thus, changes in sedimentation rates partly mirror changes in biogenic opal deposition in this area. Low LG sedimentation rates occur in all cores from the Bouvet Island area (Fig. 2). The initial rise in sedimentation rates occurs at 14e15 ka in the northern cores of the Bouvet Island area and the Western Indian sector. In contrast, the high LG to Termination 1 sedimentation rates in the Scotia Sea cores are attributed to the high input of terrigeneous material in glacial times from the nearby Patagonia and Antarctic Peninsula source regions (Diekmann et al., 2000; Diekmann, 2007).

The early Holocene is marked by dramatically elevated sedimentation rates at ca. 10e11 ka, prominent in cores from the Bouvet Island area and the Western Indian sector, and less pronounced in the Scotia Sea cores (Fig. 2). This interval corresponds to sediments composed of diatom ooze, maximum estimated SSSTs, and the almost absence of sea ice diatoms (e.g., F. obliquecostata, F. curta group) together with a maximum occurrence of open-ocean species (e.g., Fragilariopsis kerguelensis, Azpeitia tabularis) in the cores (Fig. 3aej), probably mirroring the distinct southward shift of the opal deposition zone. The high sedimentation rates of the early Holocene Bouvet Island area and the Western Indian sector cores are constrained by AMS 14C dating in different cores (Table 2, Fig. 2). Raw 14C ages for this interval were ca. 10 ka, which is close to the 14C plateau at around 10.3e10.4 ka, and spans about 300 years in calibrated age (Reimer et al., 2013). Considering the possible change in the Holocene reservoir effect (Skinner et al., 2010), this remarkably high sedimentation rate could be partly biased by the 14C plateau and changes in reservoir effect. Sedimentation rates drop dramatically after the optimum, concurrent with the increase of cold-water/sea-ice diatom occurrences. A second Holocene high sedimentation period occurs between 7 and 9 cal. ka at the northern sites of the Bouvet Island area (ODP1093 and ODP1094). Mid-late Holocene sedimentation rates are however lower than those observed during the early Holocene, indicating a northward shift of the opal deposition zone during the mid-late Holocene. Sedimentation rates generally decrease from north to south in the eastern Bouvet Island area, indicating reduced sedimentation under the influence of sea ice. In contrast, the sedimentation rates of the Scotia Sea cores are significantly influenced by terrigeneous input, showing higher LG values than during the Holocene. 5.2. Diatom composition Diatom compositions of the studied cores show a similar evolutionary pattern through the LG to the Holocene (Fig. 3). Eucampia antarctica and cold water related species (e.g., F. curta group; F. obliquecostata; Actinocyclus actinochilus) dominate the glacial diatom assemblage. Along with the deglaciation starting from ca. 18 ka, cold-water species decrease to low numbers and are replaced by open ocean species (e.g., Fragilariopsis kerguelensis and Thalassiothrix antarctica). The abundance of Azpeitia tabularis, a species most frequently encountered north of the PF (Esper and Gersonde, 2014a), increases remarkably during the early Holocene. During the mid-late Holocene, F. kerguelensis still dominates the diatom assemblage, while the F. curta group strongly increases in the Bouvet Island area, particularly at the southern sites, where its abundances are close to its LG values (Figs. 3aec; 4a). The F. curta group increase is less significant in the Scotia Sea cores, and the abundances of this species remain at trace values in the Western Indian sector core (Fig. 3j). The diatom assemblage in the mid-late Holocene also differs from the one observed during the LG by the very low abundance of E. antarctica. Chaetoceros resting spores (RS) exhibit remarkably different abundance patterns between the Scotia Sea cores and the cores from the Bouvet Island area and the Western Indian sector (Fig. 5). They show relatively low occurrence in the LG Scotia Sea sediments and high values during the Termination I and the Holocene, whereas in cores from the other two areas, they show high abundance in the LG sediments and strongly decrease in numbers towards the Holocene. At all sites, diatoms are much better preserved during the Holocene than during the LG, as indicated by the higher abundance of weakly silicified species during the Holocene, and higher dissolution/valve fragmentation during the LG. Selective dissolution plays

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123

Fig. 2. Age models and sedimentation rates of studied cores from the Bouvet Island area (aeg); Scotia Sea (hei); and Western Indian sector (j). Pluses: calibrated 14C ages (Table 2); circles: regional core correlation by diatom relative abundance, SSST variability (Suppl. Fig. 1); squares and triangles in core PS67/197-1 and PS67/219-1: correlation between magnetic susceptibility signal and EDML dust proxies, and Holocene Scotia Sea ash layer, respectively (Suppl. Fig. 2) (Xiao et al., 2016).

an important role in the preservation of the diatom frustules, especially in the southernmost Core PS1649-2, as suggested by the lower occurrence of the weakly silicified sea ice species F. curta group than in the northern cores, and a lower occurrence of the same group in LG sediments compared to Holocene sediments.

5.3. SSST and sea ice reconstructions The LG interval marks the coldest period in all the records, characterized by lowest SSSTs and highest occurrence of diatom sea ice indicators and estimated WSIC. In the central and northern Scotia Sea, the SSSTs ranged between 0.2 and 0.4
C, respectively (Fig. 3h-i). WSI extended to at least 55
S in the northern Scotia Sea, with concentration of around 80%. In the Bouvet Island area, lowest SSSTs of about 0.2
C occur in the southernmost Core PS1649-2

(55
S), and increase northwards to about 1.3
C in cores ODP1093/PS1654-2 (50
S) (Fig. 6a). SSSTs are about 0.4
C in the western Bouvet Island area Core PS2102-2. The estimated SSSTs of ca. 0.2
C in Core PS1649-2 may be biased towards warmer temperatures due to selective opal dissolution, indicated by the lower abundance of weakly silicified cold-water related species than in the northern cores (Fig. 4a). Diatom sea ice indicators suggest, that the LG WSI extent reached 50
S in the Bouvet Island area (ODP1093/PS1654-2). WSIC reached around 80% in the southern, and 50% in the northern Bouvet Island area, respectively (Fig. 4a). In the Western Indian sector Core PS2606-6, LG SSSTs were about 0.3
C, WSI extended to the core site with concentrations ranging from 50 to 80%. Glacial SSI may have reached 53.6
S (PS1651-1/ PS1652-2) in the Bouvet Island area, but did not extend to the Western Indian sector core site.

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Fig. 3. Relative abundances of diatom species, estimated SSSTs and WSIC of studied cores from the (aeg) Bouvet Island area; (hei) Scotia Sea; and (j) Western Indian sector, plotted against age. Dashed lines denote the modern SSSTs at 10 m water depth at the core sites (Locarnini et al., 2013).

W. Xiao et al. / Quaternary Science Reviews 135 (2016) 115e137

Fig. 3. (continued).

125

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W. Xiao et al. / Quaternary Science Reviews 135 (2016) 115e137

Fig. 3. (continued).

Two steps of warming (Termination Ia and Ib) characterize the deglacial process. The SSST rise and strong decreases in relative abundance of diatom cold water species and sea ice indicators in all cores since ca. 18 ka indicate a rapid decline of the WSI and SSI field along the first phase of deglacial warming (Termination Ia). The Scotia Sea records show highest resolution of the deglacial warming process. The SSSTs reached a temporal optimum at 16e14 ka of ca. 1.2
C and 1.6
C in the central and northern Scotia Sea cores, respectively (Fig. 3hei). WSI retreated in the Scotia Sea, showing strongly variable, 10e80% concentration, which could have been affected by sea ice advection from the Pacific sector. In the Bouvet Island area, the SSSTs show a distinct warming, more pronounced in the north, reaching a transient optimum at ca. 14 ka in the Bouvet Island area of 0.7e4
C from south to north. WSI retreated rapidly to about 10% concentration in the northern cores. The higher WSIC in the western (PS2102-2), and southern (PS1649-2) Bouvet Island area cores than in the northern cores (Fig. 4) suggest spatial difference of cold water influence from the Weddell Gyre. In the Western Indian sector core, SSSTs culminate at 3.1
C at about 14 ka, and the WSI concentration reduced to about 12%. Besides the Scotia Sea cores, this deglacial optimum is well documented in cores ODP1093/PS1654-2, ODP1094/PS2090-1 and PS2602-6, but not in other Bouvet Island area cores. This may be due to lower sample resolution for the deglacial interval in these cores. The timing of this optimum generally corresponds to the Antarctic Isotope Maximum (AIM) 1 peaking at ca. 14.5 ka, reflecting the Southern Ocean warming in response to reduced northward heat transport during North Atlantic stadials (EPICA members, 2006). After Termination Ia, a cooling is recorded in all cores at 12e14 ka, generally corresponding to the Antarctic Cold Reversal

(ACR) (EPICA members, 2006; Pedro et al., 2011, 2015). The SSSTs dropped to 0.6e0.7
C in the Scotia Sea cores, with the WSIC returning to about 80%. In the Bouvet Island area, the amplitudes of cooling ranged between 0.5 and 1.5
C across 55 to 50
S (Fig. 6a), accompanied by a slight rebound of diatom sea ice indicators and WSIC (Fig. 4a). In the Western Indian sector core, the SSSTs also dropped to 1.7
C with WSIC increasing to ca. 30%. After the temporal cooling, a second deglacial warming phase (Termination Ib) leads to the Antarctic early Holocene optimum (AHO) between 11 and 9 ka (Fig. 6). The SSSTs rose to 2 and 3.1
C in the central and northern Scotia Sea cores, respectively, together with WSIC decreased to about 10%. In the Bouvet Island area cores, SSSTs increased up to 2e4.9
C at 55e50
S in the Bouvet Island area, diatom sea ice indicators decreased to trace values, along with WSIC reduction to less than 15%. The peak SSST in the southernmost Core PS1649-2 occurred in one sample (2.9
C) and it is higher than the SSSTs in the northern Core PS1652-2 for the same time interval (Figs. 3a; c; 6a). This is probably due to selective dissolution, with lower relative abundance of sea ice species (e.g., F. curta group) in this sample than in the northern core. This period coincides with the highest sedimentation rates and the accumulation of almost pure diatom ooze in the Bouvet Island area cores. In the Western Indian sector core, SSSTs rose to 3
C, and sea ice was almost absent. Following the AHO, a cooling trend started at 8e9 ka, and the colder than AHO conditions persisted during the mid-late Holocene. The SSSTs dropped to about 0.8 and 1.4
C in the central and northern Scotia Sea cores, respectively; WSIC show strong variability centered at ca. 50%. In the Bouvet Island area, a rapid decrease of 2.2
C in SSSTs to 0.7
C occurred in the southernmost Core PS1649-2, and smaller amplitude of 1.2
C in the northernmost

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127

Fig. 4. Relative abundances of summer (F. obliquecostata, dark blue) and winter (F. curta group, light blue) sea ice (SSI and WSI) indicators, and winter sea ice concentration (WSIC) (gray). More than 3% (Dashed line) of the sea ice indicator species imply presence of sea ice. (a). North-south transect in the Bouvet Island area; (b). East-west transect across Western Indian and Atlantic sectors. Atmospheric CO2 records are from EDC and Byrd ice cores (Monnin et al., 2001, 2004; Ahn and Brook, 2008). EDML and WDC sea salt sodium fluxes (ssNaþ) represent changes in overall spatial sea ice coverage (Fischer et al., 2007; WAIS Divide Project members, 2013). AHO: Antarctic early Holocene Optimum; ACR: Antarctic Cold Reversal; AIM 1: Antarctic Isotope Maximum 1 (EPICA members, 2006). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

cores ODP1093/PS1654-2 to 3.7
C. Diatom WSI indicators strongly increased in the southern Bouvet Island area cores together with increased WSIC, while they remained low in the northern most cores ODP1093/PS1654-2. Much higher WSIC was calculated for intervals of 2e3 ka and 6e8 ka in Core PS1652-2 than the near-by cores PS1651-1 and ODP1094/PS2090-1, due to better preservation of lightly silicified sea ice diatoms (e.g., F. curta group) (Fig. 3bec). In the Western Indian sector core, SSSTs decreased to ca. 1.8
C, along with a slight increase of WSIC to about 15%. The strong increase in WSI indicators suggests a significant re-expansion of the WSI field in the southern Bouvet Island area, and to a lesser extent in the Western Indian sector. In contrast, SSI indicators maintained at trace values, suggesting SSI extent remained south of the studied sites.

6. Discussion 6.1. Last Glacial South of the modern PF, our records show persistent cold conditions, with SSSTs close to 0
C during the LG. Slight SSST drops between 30 and 27 ka to full glacial conditions are best recorded in the high resolution Scotia Sea (PS67/197-1), and Western Indian sector (PS2606-6) records (Fig. 6c), and mirrors the atmospheric cooling as recorded in Antarctic ice cores (EPICA members, 2006). The LG SSSTs are colder than Holocene conditions by 1e3
C at sites south of the modern PF. This is significantly lower than the 4e8
C reconstructed for the LG-Holocene in the subantarctic and subtropical zones (Gersonde et al., 2005; MARGO members, 2009; Ho

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Fig. 5. Relative abundance of Chaetoceros RS in cores from the Western Indian sector, Bouvet Island area and Scotia Sea, compared to Scotia Sea Iceberg Rafted Debris flux (IBRD, Weber et al., 2014) and EDML nssCa2þ record as an indication of dust flux (Fischer et al., 2007).

et al., 2012; Marr et al., 2013). The small amplitude LG cooling in our Antarctic Zone records is due to the limit represented by the sea water freezing temperature. Compared to modern conditions, the LG WSI expanded northward by ~5
latitude to 50
S in the Bouvet Island area, and it also extended north of our Scotia Sea site PS67/197-1 (55
S) (Fig. 4a; b), probably to ca. 52
S (KC73) (Allen et al., 2005). This reconstruction is consistent with a previous study suggesting that the LG WSI extended to the modern PF location (Gersonde et al., 2005). The diatom SSI indicator suggests the sporadic occurrence of permanent sea ice cover in the southern Bouvet Island area (PS1649-2). Although both our Scotia Sea records show slightly less than 3% of the SSI indicator F. obliquecostata, other records in this area indicate, that SSI might have extended to ~56
S (TPC78, close to PS67/ 197) in the northern Scotia Sea (Allen et al., 2011; Collins et al., 2012, 2013). Such difference may come from preservation differences between cores. The diatom sea ice indicator data suggest a much more extended SSI field during the LG. In contrast, maximum SSI extent was assumed to be similar to its modern position south

of 60
S in the Atlantic sector during the LGM (Crosta et al., 1998b). The difficulties of analyzing the SSI extent and concentration are mainly due to strong opal dissolution in the permanent sea ice covered area, where little information is available for reconstructions of sea ice extent based on siliceous microfossils (Gersonde et al., 2005; Esper and Gersonde, 2014b). The Western Indian sector Core PS2606-6 and northern Bouvet Island cores ODP1093/PS1654-2 and PS1768-8 show maximum WSIC at 18e25 ka, the time interval of lowest SSSTs (Figs. 4a, b; 6a, c). This period is assumed to mirror the maximum north/northeastward cold water expansion from the Weddell Gyre during full glacial times, and is in agreement with the maximum extent of Patagonian glaciers (McCulloch et al., 2005; Kaplan et al., 2008; Hein et al., 2010). This time interval also generally corresponds to the lowest sea level with maximum icesheet growth between ca. 19 and 26.5 ka (Peltier and Fairbanks, 2006; Clark et al., 2009). Our results support the idea that the coldest time interval in the Southern Ocean was prior to the LGM, a time period conventionally defined at 19e23 ka (Gersonde et al., 2003, 2005). WSI indicators show a slightly different pattern with their maximum abundance at 23e29 ka. Our data infer the presence of a LG northward-extended cold water belt. The LG SSST gradients are small between the northern and central Scotia Sea (~0.2
C difference across 2
latitude between PS67/197-1 and PS67/219-1) (Fig. 3h, i), and between 50 and 55
S (~1
C: ODP1093 and PS1649-2) in the Bouvet Island area (Fig. 6a). Compared to previous South Atlantic SST reconstructions (Fig. 6a), a stronger gradient was present in the Subantarctic Zone (~11
C difference across 9
latitude between ODP1089 and ODP 1093). This supports a LG northward shift of SWW (Toggweiler and Russell, 2008; Kohfeld et al., 2013), at least its southern boundary. The northward shifted SWW, in combination with the buoyancy loss due to expanded sea ice cover to the Bouvet Island area (Ferrari et al., 2014; Watson et al., 2015), would have reduced the nutrient supply to the surface water in the south by reducing winddriven upwelling (Fig. 4a), thus suppressing diatom production and opal deposition (Trull et al., 2001; Anderson et al., 2009). Such a scenario is supported by the LG low sedimentation rates and low diatom concentration in our Bouvet Island area cores, similar to reduced opal flux in the adjacent core TN057-13, suggesting northward displacement of opal deposition to at least north of 50
S (Bradtmiller et al., 2009; Anderson et al., 2009). The LG expanded sea ice field in the South Atlantic would have strongly reduced the surface ocean heat transport from the Pacific Ocean in to the South Atlantic via the Drake Passage. The strong northward cold water expansion would also push the whole ACC oceanic frontal system along with the SWW northward, and also result in a reduced Agulhas heat leakage from the Indian to the Atlantic Oceans (Franzese et al., 2006; Bard and Rickaby, 2009). This may result in the observed stronger cooling in the LGM tropical Atlantic than in the tropical Indian and Pacific Oceans (MARGO members, 2009). The expanded sea ice field would allow for enhanced brine rejection during sea ice formation that contributes to the formation of glacial AABW (Watson and Naveira Garabato, 2006). The extended AABW formation area would in turn push the admixing zone of North Atlantic Deep Water (NADW) and CDW farther north than today (Curry and Oppo, 2005; Marchitto and Broecker, 2006). High amounts of Chaetoceros RS occurred in the Bouvet Island area (30e40%) cores and decreased towards both the Scotia Sea (5e10%) and the Western Indian sector (~10%) during the LG (Fig. 5). Chaetoceros RS are considered to be an indicator of carbon export (Abelmann et al., 2006) and the observed pattern is generally in agreement with the Antarctic dust concentration record

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Fig. 6. SSST records from Southern Ocean Atlantic and western Indian sectors compared to the EDML d18O, James Ross Island ice core temperature anomaly, and WDC d18O. (a) Bouvet Island area east transect; (b) Bouvet Island area west transect; (c) Longitudinal transect of Western Indian sector, Bouvet Island area and Scotia Sea. For data references see Table 1.

(Fischer et al., 2007; Lambert et al., 2008). Sea ice acts as an effective transporter of iron and enhances its bioavailability in the ocean (Abelmann et al., 2006). The strongly expanded sea ice field in the Bouvet Island area provided readily bioavailable iron from Patagonian wind borne dust, which settled on sea ice and fertilized the upper ocean during the melting season. In the Western Indian sector core, at the maximum eastward expansion of WSI and far end of Patagonian dust plume, Chaetoceros RS are less abundant than in the Bouvet Island area cores. In contrast, in the Scotia Sea cores, despite their closer location to the Patagonian dust source, the LG Chaetoceros RS represent only 5e10% of the diatom assemblage, much less than in the cores to the east. This may be linked to the northward shifted SWW, where the main dust transport track was probably located further away from the Scotia Sea. More importantly, the possible occurrence of permanent sea ice cover (Allen et al., 2011; Collins et al., 2012, 2013) and northward shifted SWW (Toggweiler et al., 2006) might have lowered diatom productivity and shortened the blooming season of Chaetoceros RS in the Scotia Sea. The north/northeastward extension of the Chaetoceros RS zone towards the modern POOZ indicates an expanded area of carbon export in the South Atlantic south of the modern PF (Abelmann et al., 2006, 2015). The intensively extended area for brine rejection during sea ice formation resulted in efficiently transport of atmospheric CO2 to the abyss, and being kept in the stratified deep ocean (Hillenbrand and Cortese, 2006; Bouttes et al., 2009, 2010; Ferrari et al., 2014).

6.2. Termination 6.2.1. Start of the deglacial warming, Termination Ia A slight SSST increase is recognized in the Scotia Sea cores between 20 and 18 ka, but not in the Bouvet Island area cores probably due to their lower resolution. Such timing may support the early deglacial d18O decrease starting from ~22 to 21 ka as recorded in the WAIS Divide ice core (WDC), located at relatively low altitude and primarily recording the climate signal in warm seasons (WAIS Divide Project members, 2013). In contrast, the East Antarctic ice cores show deglacial warming starting at 18 ka (EPICA members, 2006; Jouzel et al., 2007; Stenni et al., 2011). Such lagged warming signal is biased towards cold seasons (Laepple et al., 2011; Küttel et al., 2012). In addition, the signals registered in WDC ice core at lower altitude better represent the conditions over adjacent high latitude Southern Ocean, compared to the high altitude ice cores, which receive moisture largely from lower latitude oceans (Delaygue et al., 2000; Sodemann and Stohl, 2009). The earlier deglacial warming in WDC is in agreement with the initial sea ice decrease signal, in response to spring insolation increase over the adjacent ocean surface (WAIS Divide Project members, 2013) (Figs. 6; 7m). Such sea ice reduction is not reflected in the WSIC but recognized in the decrease of sea ice indicators in the Scotia Sea cores. This may indicate shortened WSI duration, and thus enhanced seasonality at this time. This interval also shows elevated Scotia Sea Chaetoceros RS abundance (Figs. 3hei; 5), corresponding

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to the increase of iceberg rafted debris (IBRD) flux in the Scotia Sea at 20e19 ka (Weber et al., 2014). The timing corresponds to initial deglacial melt water input from Antarctic Peninsula Ice Sheets (Heroy and Anderson, 2007). A major step of SSST rise and sea ice reduction occurred at about 18 ka (Figs. 4; 5). This is also recorded by lighter d18O and heavier d13C in planktonic foraminifera in Bouvet Island area indicating surface water freshening and improved ventilation (Bianchi and Gersonde, 2004). The timing of Termination Ia is in agreement with that of the Australian-New Zealand sector of the Southern Ocean (Bostock et al., 2013), and East Antarctic ice core temperature records (Lemieux-Dudon et al., 2010). It also coincides with the retreat of the grounded ice from the Antarctic Peninsula outer shelf  Cofaigh et al., 2014) and various southern mid-latitude glaciers (O (Alloway et al., 2007; Rodbell et al., 2009; Hein et al., 2010) at 17e18 ka. The abrupt SSST rise is concurrent with the deglacial atmospheric CO2 rise (Monnin et al., 2001) and improved bottomsurface water ventilation (Barker et al., 2010; Skinner et al., 2010) (Fig. 7eef, k). This interval coincides with the Heinrich Stadial 1 (H1), which is characterized by strong reduction of NADW formation, thus weakened Atlantic Meridional Overturning Circulation (AMOC) (McManus et al., 2004; Gherardi et al., 2009) (Fig. 7aeb). The latter would have resulted in warm water transport from the tropics to the North Atlantic being reduced, and heat accumulated in the South Atlantic (Seidov and Maslin, 2001). Thus, our records support the “bipolar seesaw” mechanism (Rahmstorf, 2002; Stocker, 2003). Compared to the subantarctic and subtropical Atlantic and Pacific records (Lamy et al., 2007; Barker et al., 2009; Bostock et al., 2013) (Fig. 7c), our data suggest a smaller amplitude of warming, together with delayed WSIC reduction towards the south. This implies a quick response to heat accumulation in the subantarctic and subtropical regions, and more continuous influence of cold water from the Weddell Gyre in the southern Bouvet Island area. 6.2.2. Antarctic Isotope Maximum 1 and Antarctic Cold Reversal Termination Ia reached a transient optimum at around 14 ka in the Western Indian sector (PS2606-6) as well as in the northern Bouvet Island area (PS1654-2/ODP1093), and was characterized by a strong increase of the warmer water species A. tabularis in both cores. An abrupt warming is also recorded in cores PS2090-1/ ODP1094 and adjacent Core TN057-13 (Divine et al., 2010), and to a lesser extent in other Bouvet Island area cores, partly due to lower sample resolution (e.g., the northerly located Core PS1768-8). Sea ice indicators and WSIC records clearly suggest a rapid decline of sea ice field in all cores, which allowed more heat transport from the Pacific and Indian Oceans to the South Atlantic via the water routes. The timing generally corresponds to the Antarctic Isotope Maximum (AIM) 1 as recorded in Antarctic ice cores (EPICA members, 2006). This signal is delayed until the early Holocene in other Bouvet Island area cores, suggesting a spatial difference, with more persistent cold water influence in the western and southern Bouvet Island area. This warming was accompanied by a southward shift in SWW, enhanced wind-driven upwelling and the southward displacement of opal deposition (Anderson et al., 2009).

Fig. 7. SSST records from (g) Western Indian sector, (h) Bouvet Island area and (i) Scotia Sea compared to (a) NGRIP d18O; (b) North Atlantic AMOC intensity (red: 232Th based values and blue: 238U based values); (c) South Pacific SST; (d) South Atlantic

planktic foraminifera cold water species variation; (e) bottom-surface water (red: benthic-planktic; blue: projection age model), and (f) bottom water-atmosphere ventilation history (spline-smoothed); (j) Antarctic Palmer Deep SST; (k) EDC and Byrd CO2 records; (l) EDML d18O and JRI temperature anomaly; (m) summer and spring insolation changes at 54
S. YD: Younger Dryas; H1: Heinrich stadial 1; B/A: BøllingAllerød warm interval. For data references see Table 1. (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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In Antarctic ice cores, Termination Ia leading to AIM 1 peaks at ca. 14.5 ka, earlier than our marine records by a few hundred years. The determination of accurate timing of the AIM 1 in our marine records is hampered by low sample resolution and uncertainties in changes of carbon reservoir effect during the deglaciation. During the inception of the deglaciation, the wind-driven upwelling released the 14C depleted CO2 from the ocean to the atmosphere, corresponding to the abrupt decrease of atmospheric D14C and significantly enhanced ocean ventilation (Anderson et al., 2009; Barker et al., 2010; Skinner et al., 2010; Siani et al., 2013). This may have reduced the surface ocean reservoir age, thus possibly introduced bias in the accurate timing in our records. In addition, our dating on organic carbon of the glacial-deglacial interval may potentially be biased towards younger ages. In fact, chronologies constrained by planktonic foraminiferal dating, as in the nearby Core TN057-13, closely approximate the age for this transition as recorded in the Antarctic ice core records (Anderson et al., 2009; Divine et al., 2010). Besides, the ice core chronologies themselves also have 200e300 years error for this time interval (LemieuxDudon et al., 2010; Stenni et al., 2010, 2011). The SSSTs derived from the high resolution Scotia Sea records (PS67/197-1 and PS67/219-1) show rather stable conditions between 16 and 14 ka, most similar to the pattern observed in the d18O records from the Atlantic sector ice cores EDML and Dome F, but contrast to the one recorded in the Indian and Pacific sector ice cores EDC, Vostock and Talos Dome, which show continuous warming till ca. 14.5 ka (Stenni et al., 2011). Considering the different moisture source regions for ice cores, we suggest that the Scotia Sea may be an important moisture source region for the Atlantic sector ice cores, thus may give a hint to the atmospheric circulation situation and related water mass and sea ice distribution during the Termination. Following Termination Ia, the deglacial warming was interrupted by a SSST drop of ~0.3e1.3
C and a sea ice indicator/WSIC increase between 12 and 14 ka in our records (Figs. 4; 6). Such a cooling is also recorded by the foraminiferal assemblages at the South Atlantic Subtropical Front (Barker et al., 2010) (Fig. 7d), and by radiolarian-derived SSSTs with similar amplitude (ODP1089/ PS2821-1, Cortese et al., 2007) (Fig. 6a). The SSSTs drop to close to glacial values in our southern Bouvet Island area cores, suggesting intense cooling in areas with large variations of cold water/sea ice extent. However, the less abundant sea ice species and better preservation indicate overall warmer conditions than the LG. The timing is broadly consistent with a cooling phase in Patagonia and the Andes (Rodbell et al., 2009; Hein et al., 2010; Jomelli et al., 2014), New Zealand (Hajdas et al., 2006; Putnam et al., 2010), and marine records from Southern Ocean Pacific and Indian sectors (Alloway et al., 2007; Barrows et al., 2007; Bostock et al., 2013). This interval generally corresponds to the Antarctic Cold Reversal (ACR) recorded in Antarctic ice cores, which is tightly coupled with the Bølling-Allerød (B-A) warm interval in the Northern Hemisphere (Pedro et al., 2011). The north warming/south cooling is regarded to be associated with the rapid resumption of NADW formation together with the shifting SWW (Knorr and Lohmann, 2003, 2004; Anderson et al., 2009), which removes heat from the Southern Ocean. In the high latitude Southern Ocean, this bipolar seesaw cooling process overwhelms the cross-equatorial atmospheric heat flux from the Northern Hemisphere (Pedro et al., 2015). The ACR cooling was amplified by expansion and thickening of sea ice in the south (Pedro et al., 2015). This is supported primarily by the strong increase of WSIC in our Scotia Sea cores, and to a lesser extent by the increase in WSIC and diatom sea ice indicators in cores from the Bouvet Island area and the Western Indian sector. Sea ice expansion during the ACR can be related to significant ice melt from the Antarctica, which contributed to the global sea level

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rise of approximately 1.7e4.3 m during the Malt-Water Pulse (MWP) 1a at this time interval (Golledge et al., 2014). The deglacial melting and thinning of the WAIS continuously introduced meltwater in the adjacent ocean. This is reflected by Scotia Sea iceberg rafted debris (IBRD) indicating major Antarctic iceberg discharge (Weber et al., 2014), and is also confirmed by the concurrent high Chaetoceros RS abundance in our Scotia Sea records. Although the dust input was strongly reduced during this period (Fischer et al., 2007; Lambert et al., 2008), the increases of Chaetoceros RS in our records may reflect fertilization of iron released by sea ice and iceberg during the melting season, in combination with enhanced surface water stratification induced by meltwater input (Crosta et al., 1997). The pronounced Chaetoceros RS increase is not recorded in the Bouvet Island area and Western Indian sector cores, suggesting a diminished influence of ice-delivered iron and melt water to the open ocean South Atlantic. The overall Southern Ocean cooling during the ACR allowed cold water and sea ice export to the Bouvet Island area and Western Indian Sector. As a major area for Antarctic iceberg discharge along the “iceberg alley” (Gladstone et al., 2001; Silva et al., 2006), the South Atlantic experienced stronger cooling during the ACR compared to other Southern Ocean sectors (Pedro et al., 2015). The meltwater discharge also strongly reduced the AABW formation, which further amplified the cooling in the South Atlantic (Golledge et al., 2014). The difference in amplitude of warming/cooling in the South Atlantic changed the SST latitudinal gradient, and was accompanied by the migration of the SWW and associated upwelling regions. Deglacial maximum opal flux (coinciding with the maximum warming) mirrors increased opal production and silicon supply, linked to upwelling driven by the poleward-shifted SWW (Marchitto et al., 2007; Anderson et al., 2009; Rose et al., 2010). Our records are in agreement with the results of Anderson et al. (2009) that the temporary warming (opal deposition) maximum intruded to 53.2
S in the Bouvet Island area, along with the southward migration of SWW and upwelling zone. In addition, our records suggest that during the ACR, all the related features (e.g., region of strong upwelling and maximum opal flux) may have shifted northward to ca. 50
S (ODP1093) (Figs. 2; 5). Deglacial abrupt atmospheric CO2 rise is coeval with the H1 and Younger Dryas (YD) cold phase in the Northern Hemisphere (Lourantou et al., 2010). It also coincides with the abrupt SSST increase and sea ice retreat in our records (Figs. 4; 7a, k). This indicates the South Atlantic south of the Polar Front as a potential important source region of deglacial CO2 release. During the ACR, the atmospheric CO2 remained stable (Monnin et al., 2001), in agreement with the return to almost full WSI cover at the Scotia Sea sites, and slight increase in WSIC in the Bouvet Island area (Fig. 4a). The second step of CO2 rise towards the AHO (Termination Ib) is also accompanied by a strong decrease in WSIC in the Scotia Sea, as well as a further retreat of WSI in the Bouvet Island area and the Western Indian sector. This may suggest the source region of CO2 release during the second warming phase shifted/expanded further to the south. 6.3. Holocene 6.3.1. Antarctic early Holocene optimum After the ACR, Termination Ib led the warming towards the AHO, exceeding modern SSSTs by 1e2
C in the Bouvet Island area and at the Western Indian sector site, and being slightly lower than modern values in the Scotia Sea (Fig. 3). The AHO in our records is generally confined at 11e9 ka, a timing which is synchronous with marine records in the Pacific sector of the Southern Ocean (Lamy and de Pol-Holz, 2013; Pahnke and Sachs, 2006; Bostock et al., 2013); the tropical Andes (Thompson et al., 1998, 2000), and most

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Fig. 8. Modern (a) and time slice reconstructions of past (bef) SSSTs and sea ice conditions, and schematic illustration of climate development (gel) during the past 30 ka. The reconstructed SSSTs and WSIC are indicated in brackets. ANT: Antarctica; F & R: Filchner and Ronne Shelf Ice; SWW: Southern Westerly Wind; AC: Agulhas Current; AABW: Antarctic Bottom Water; CDW: Circumpolar Deep Water; AAIW: Antarctic Intermediate Water; NADW: North Atlantic Deep Water. Sea ice edge is defined as 15% sea ice

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Antarctic ice core records (Masson et al., 2000; EPICA members, 2006; Stenni et al., 2011; Mulvaney et al., 2012). In contrast, the AHO timing is heterogeneous in some Antarctic coastal marine and ice core records, depending on local deglaciation histories (Cremer et al., 2007; Wagner and Melles, 2007; Denis et al., 2009; Stenni et al., 2011). The advances of some Patagonian glaciers during this period are ascribed to increased precipitation (Ackert et al., 2008; Rodbell et al., 2009). During this period, the sea ice field retreated south of the Bouvet Island area and the central Scotia Sea, as indicated by WSIC reduced to <15% at sites PS1649-2 and PS67/2191, together with the strong decrease of cold water/sea ice diatoms (Fig. 4). This time interval marks the warmest conditions and strongest sea ice reduction in the Southern Ocean since the last glacial. Termination Ib corresponds to a weakened AMOC during the YD (Fig. 7b), which kept the heat in the Southern Ocean by virtue of the bipolar seesaw mechanism. After the YD, the re-intensification of AMOC is supposed to trigger cooling in the southern high latitudes (Masson et al., 2000). However, our records show high SSSTs were maintained till ca. 10 ka, which is in agreement with what observed in Antarctic ice core records (Stenni et al., 2011). This has been attributed to the maximum annual mean and summer insolation at southern high latitudes during the early Holocene (Shevenell et al., 2011; Etourneau et al., 2013). Moreover, the reduced sea ice field would have reduced AABW formation by brine rejection (glacial mode, Watson and Naveira Garabato, 2006). The cavities beneath ice shelves were not fully developed (Hillenbrand et al., 2014), which further weakened AABW production. Consequently, the enhanced production of NADW at that time (McManus et al., 2004) (Fig. 7b) would have resulted in this water mass mixing with CDW and intruding further south into the Southern Ocean than under modern conditions. This would introduce relatively warm deep water in the high latitudes Southern Ocean, eventually upwelling to the surface, and promoting sea ice/ice shelf melt. In turn, the retreat of sea ice would further enable enhanced warm water transport from the Pacific and Indian Oceans, and southward propagation of warm surface water. The southward intrusion of CDW is recorded as subsurface warming in the Antarctic Peninsula region (Shevenell et al., 2011; Etourneau et al., 2013) (Fig. 7j). The upwelled warm CDW promoted ice shelf melting (Etourneau et al., 2013), and enhanced glacier discharge, as indicated by increased sediment magnetic susceptibility and decreased diatom d18O (Domack et al., 2001; Pike et al., 2013). This is also reflected by the high abundance of IBRD in the Scotia Sea (Weber et al., 2014), and the increased occurrence of Chaetoceros RS in Palmer Deep (Shevenell et al., 2011) and our Scotia Sea records (Fig. 5). The influence of glacier discharge did not extend to the Bouvet Island area and further east, and this process may thus be responsible for the less pronounced AHO warming in the Scotia Sea (especially the southern/central Scotia Sea, Fig. 3i) than in the Bouvet Island area and the Western Indian sector. Our data show warm SSSTs, but small difference (~1
C) between 50 and 55
S in the Bouvet Island area during the AHO. This is in agreement with the poleward-shifted, but weak SWW at this latitudinal band (Mayr et al., 2007; Lamy et al., 2010; Moreno et al., 2012), resulting in increased accumulation rate (more moisture) in Antarctic ice cores (Ruth et al., 2007), and small variation of atmospheric CO2 concentration (Monnin et al., 2001) (Fig. 7k). Previous studies suggested that the enhanced upwelling and opal

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deposition zone shifted southwards at least to 53.2
S in the Bouvet Island area (TN057-13) between 10 and 13 ka (Anderson et al., 2009). Our data further indicate that this pattern expanded to at least 55
S (PS1649-2). Our dating results also show that the early Holocene maximum opal flux is confined between 11 and 10 ka, at least in the southern cores, corresponding to the maximum sedimentation rates and pure diatom ooze intervals in the sediments of Bouvet Island area cores. The good preservation of weakly silicified diatom species, and minor contribution of terrigenous material in the early Holocene sediments from all the studied cores exclude lateral transport or sediment focusing as the reason for such high sedimentation rates. The discrepancy in the timing of the maximum opal flux compared to Core TN057-13 (Anderson et al., 2009) may be due to the age model of the latter core not being well constrained in the 14e10 ka interval. 6.3.2. Mid-late Holocene cooling The mid-late Holocene cooling, starting at 8e9 ka in our records, is characterized by SSSTs drops of 1.3e2.4
C to ACR levels south of the modern PF, along with increases in diatom cold water and sea ice species. WSIC increased to ca. 20% in the Bouvet Island area at 53
S and at the Western Indian sector site. The cooling was also recognized in other marine records from the South Atlantic (Hodell et al., 2001; Divine et al., 2010), Antarctic Peninsula (Shevenell et al., 2011; Etourneau et al., 2013), and South Pacific (Pahnke and Sachs, 2006), and was defined as “neoglaciation” in the Southern Ocean. General sea ice expansion around Antarctica after the AHO was documented in several ice cores by the ssNaþ flux (Fischer et al., 2007; Schüpbach et al., 2013). Our marine records provide unequivocal evidence of sea ice occurrence in the Bouvet Island area, presently located in the POOZ. It marks a strong cold water reexpansion of the Weddell Gyre after the AHO. The colder than modern mid-late Holocene conditions transitioned into warming and sea ice retreat over last few decades, as recorded by in situ observations and satellite monitoring. Model simulations suggest strong cooling in the South Atlantic between 9 and 7 ka, as a response to the deglaciation of the Laurentide Ice Sheet, which cooled the NADW and upwelled and cooled the Southern Ocean (Renssen et al., 2010). Such interpretation may explain the initial cooling, but not the persistent midlate Holocene cold conditions in our records south of the modern PF. Ice core records from the Antarctic plateau suggest a second warming after the initial cooling during the mid-Holocene (EPICA members, 2006; Jouzel et al., 2007), reflecting mid-latitude ocean conditions (Delaygue et al., 2000; Sodemann and Stohl, 2009). In contrast, similar to our marine records south of the modern PF, a James Ross Island ice core record (JRI) from the Antarctic Peninsula shows persistent colder than AHO conditions during the mid-late Holocene, and its temperature evolution is strongly connected to the grounding line retreat of the ice shelf (Mulvaney et al., 2012). Continuous mid-late Holocene cooling was also observed in marine records in the Ross Sea sector (Steig et al., 1998; Das and Alley, 2008). This cooling has been attributed to the decrease of summer duration, which covaries with austral spring insolation, rather than maximum summer insolation in the southern high latitudes (Fig. 7m) (Huybers and Denton, 2008; Shevenell et al., 2011). The deglaciation of WAIS proceeded with continuous thinning and retreating grounding line during the Holocene (Hillenbrand et al.,  Cofaigh et al., 2014). This allowed 2014; Anderson et al., 2014; O

concentration. Modern SSSTs are from World Ocean Atlas 2013 (Locarnini et al., 2013). Positions of Modern Winter and Summer Sea Ice edges (MWSI and MSSI) and PF (MPF) are from Comiso (2003) and Orsi et al. (1995), respectively. Yellow solid and dashed lines denote reconstructed WSI and SSI edges, respectively. Core location and references refer to Fig. 1 and Table 1. Dashed line in (i) indicates strongly reduced NADW formation during H1 and YD (McManus et al., 2004), and in (j) AABW formation during ACR (Golledge et al., 2014). (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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cold surface water generation under the shelf ice, with this cold water being transported to the open ocean by the Weddell Gyre, and contributing to maintain the cold conditions in the high latitude Atlantic Southern Ocean. The developing shelf-ice cavities also allowed for enhanced AABW formation, which may have reduced CDW upwelling in the shelf area, as reflected by a subsurface temperature decrease in the Palmer Deep (Shevenell et al., 2011; Etourneau et al., 2013) (Fig. 7j). The strong mid-late Holocene cooling was not persistent at ODP1093 and further north, and the cooling was less pronounced in the Western Indian sector than in the Bouvet Island area (Fig. 6a; c). This suggests the influence of cold-water expansion from the Weddell Gyre diminished towards the north/northeast. During this time interval, the SSST difference in the Bouvet Island area is small (<0.5
C) between 55 (PS1649-2) and 53.2
S (ODP1094), but increases (~2.5
C) towards the modern PF (50
S, ODP1093). This is accompanied by decreased opal deposition and weakened SWW over TN057-13, close to ODP1094 (Anderson et al., 2009). This scenario is consistent with Southern Patagonia records (Douglass et al., 2005; Lamy et al., 2010) suggesting a northward shift of the maximum wind stress zone after the AHO. In contrast, southward migration of the SWW was inferred from the position shifts of the confluence of the subtropical Brazil Current and the subantarctic Malvinas Current (Brazil-Malvinas Confluence, BMC) (Voigt et al., 2015). Under modern conditions, the BMC is located at ca. 40
S aligning with the STF, and north of the core of the SWW centered at 50e55
S (Schneider et al., 2003) (Fig. 1). Thus, the southward shift of BMC is likely related to weakened wind stress of the northern limb of the SWW. The northward shift of the steep SSST gradient zone towards the modern PF indicated by our data suggests a northward migration of the SWW, at least its southern boundary. In spite of the WSI extending to the Bouvet Island area (Fig. 4a), Chaetoceros RS remained low in abundance in the mid-late Holocene sediments from the Bouvet Island area and the Western Indian sector cores (Fig. 5). This reflects the strongly reduced dust input to this region during the Holocene (Fischer et al., 2007). In contrast, high abundance of Chaetoceros RS in the Scotia Sea may be caused by the ice discharge from Antarctica (Weber et al., 2014). The spatial differences in Chaetoceros RS abundance evolution pattern between the Bouvet Island area and the Scotia Sea cores suggest different sources for bioavailable iron in these two regions. 7. Summary and conclusions By compilation of high resolution diatom composition records from the Antarctic Zone of Atlantic and Western Indian sectors of the Southern Ocean, we determined spatial and temporal variability in SSST and sea ice conditions of the past 30 ka, which also has implications for surface/deep water circulation, the overlying atmospheric circulation and carbon cycle (Fig. 8). The LG South Atlantic is by 1e3
C cooler than present in the area south of the modern PF, a cooling amplitude smaller than Subantarctic and Subtropical Zone. WSI extended to the modern Polar Front. Along with the SWW and the upwelling zone, the maximum opal deposition also shifted to north of 50
S in the Bouvet Island area. Our data suggest that the South Atlantic Antarctic Zone became a carbon sink during the LG. The deglacial evolution in the studied area supports the bipolar see-saw mechanism. The covariance of SSST increase/sea ice reduction and the increase of atmospheric CO2 points to the South Atlantic Antarctic Zone as an important source region of deglacial CO2 degassing. More persistent cold conditions to the south are attributed to ice discharge from Antarctica throughout the deglaciation.

The early Holocene optimum is characterized by SSSTs exceeding modern values, and a WSI edge which probably located south of its modern position. The maximum opal deposition zone shifted southward at least to 55
S in the Bouvet Island area, indicating a southward migration of the SWW. The mid-late Holocene cooling is characterized by SSST drops and WSI expansion to the Bouvet Island area, probably related to cold water generation under the Western Antarctic Ice Shelf. The cooling at this time suggests a northward shift of the SWW, at least its southern boundary. Acknowledgement This work was co-supported by the “Past4Future” Climate change e learning from the past climate, the Alfred Wegener Institute PACES program (Polar Regions and Coasts in the Changing Earth System), and MARUM (The Ocean in the Earth System). We acknowledge U. Bock and R. Cordelair for their technical assistance. We are grateful to G. Cortese and an anonymous reviewer for their constructive and detailed comments which helped to improve the paper. Appendix A. Supplementary data Supplementary data related to this article can be found at http:// dx.doi.org/10.1016/j.quascirev.2016.01.023. References Abelmann, A., Gersonde, R., 1991. Biosiliceous particle flux in the Southern Ocean. Mar. Chem. 35, 503e536. Abelmann, A., Gersonde, R., Cortese, G., Kuhn, G., Smetacek, V., 2006. Extensive phytoplankton blooms in the Atlantic sector of the glacial Southern Ocean. Paleoceanography 21, PA1013. http://dx.doi.org/10.1029/2005PA001199. Abelmann, A., Gersonde, R., Knorr, G., Zhang, X., Chapligin, B., Maier, E., Esper, O., Friedrichsen, H., Lohmann, G., Meyer, H., Tiedemann, R., 2015. The seasonal seaice zone in the glacial Southern Ocean as a carbon sink. Nat. Commun. 6, 8136. http://dx.doi.org/10.1038/ncomms9136. Ackert Jr., R.P., Becker, R.A., Singer, B.S., Kurz, M.D., Caffee, M.W., Mickelson, D.M., 2008. Patagonian glacier response during the late glacial-Holocene transition. Science 321, 392e395. Ahn, J., Brook, E.J., 2008. Atmospheric CO2 and climate on millennial time scales during the Last Glacial period. Science 322, 83e85. Allen, C.S., Pike, J., Pudsey, C.J., Leventer, A., 2005. Submillennial variations in ocean conditions during deglaciation based on diatom assemblages from the southwest Atlantic. Paleoceanography 20, PA2012. http://dx.doi.org/10.1029/ 2004PA001055. Allen, C.S., Pike, J., Pudsey, C.J., 2011. Last glacialeinterglacial sea-ice cover in the SW Atlantic and its potential role in global deglaciation. Quat. Sci. Rev. 30, 2446e2458. Alloway, B.V., Lowe, D.J., Barrell, D.J.A., Newnham, R.M., Almond, P.C., Augustinus, P.C., Bertler, N.A.N., Carter, L., Litchfield, N.J., McGlone, M.S., Shulmeister, J., Vandergoes, M.J., Williams, P.W., NZ-INTIMATE members, 2007. Towards a climate event stratigraphy for New Zealand over the past 30000 years (NZ-INTIMATE project). J. Quat. Sci. 22, 9e35. Anderson, R.F., Ali, S., Bradtmiller, L.I., Nielsen, S.H.H., Fleisher, M.Q., Anderson, B.E., Burckle, L.H., 2009. Wind-driven upwelling in the Southern Ocean and the deglacial rise in atmospheric CO2. Science 323, 1443e1448. Anderson, J.B., Conway, H., Bart, P.J., Witus, A.E., Greenwood, S.L., McKay, R.M., Hall, B.L., Ackert, R.P., Licht, K., Jakobsson, M., Stone, J.O., 2014. Ross Sea paleoice sheet drainage and deglacial history during and since the LGM. Quat. Sci. Rev. 100, 31e54. Bard, E., 1988. Correction of accelerator mass spectrometry 14C ages measured in planktic foraminifera: paleoceanographic implications. Paleoceanography 3, 635e645. Bard, E., Rickaby, R.E.M., 2009. Migration of the subtropical front as a modulator of glacial climate. Nature 460, 380e383. Barker, S., Diz, P., Vautravers, M.J., Pike, J., Knorr, G., Hall, I.R., Broecker, W.S., 2009. Interhemispheric Atlantic seesaw response during the last deglaciation. Nature 457, 1097e1102. Barker, S., Knorr, G., Vautravers, M.J., Diz, P., Skinner, L.C., 2010. Extreme deepening of the Atlantic overturning circulation during deglaciation. Nat. Geosci. 3, 567e571. Barrows, T.T., Juggins, S., De Deckker, P., Calvo, E., Pelejero, C., 2007. Long-term sea surface temperature and climate change in the AustralianeNew Zealand region. Paleoceanography 22, PA2215. http://dx.doi.org/10.1029/2006PA001328. Bathmann, U., Schulz-Baldes, M., Fahrbach, E., Smetacek, V., Hubberten, H.-W., et al.,

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