Constraining the dating of late Quaternary marine sediment records from the Scotia Sea (Southern Ocean)

Quaternary Geochronology 31 (2016) 97e118

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Quaternary Geochronology journal homepage: www.elsevier.com/locate/quageo

Research paper

Constraining the dating of late Quaternary marine sediment records from the Scotia Sea (Southern Ocean) Wenshen Xiao a, b, *, Thomas Frederichs c, Rainer Gersonde a, Gerhard Kuhn a, Oliver Esper a, Xu Zhang a a b c

Alfred-Wegener-Institut Helmholtz-Zentrum für Polar- und Meersforschung, Bremerhaven D-27568, Germany State Key Laboratory of Marine Geology, Tongji University, Shanghai 200092, China €t Bremen, D-28334, Germany Fachbereich Geowissenschaften, Universita

a r t i c l e i n f o

a b s t r a c t

Article history: Received 26 February 2015 Received in revised form 3 November 2015 Accepted 5 November 2015 Available online 11 November 2015

In Southern Ocean sediments south of the Antarctic Polar Front, the scarcity of calcareous microfossils hampers the development of sediment chronologies based on radiocarbon dating and oxygen isotope stratigraphy established from carbonate. In this study, radiometric dating, magnetic susceptibility (MS), biogenic opal content, diatom abundance fluctuation, and paleomagnetic information were investigated on a northesouth transect of central Scotia Sea sediment cores to verify their reliability as stratigraphic tools in the study area. Radiocarbon dating on organic carbon humic acid fraction can be used to establish the stratigraphy of upper core sections, but regional comparison and correlation are needed to verify a possible bias by fossil carbon contamination. For the long-term stratigraphy, MS, which can be correlated to the Antarctic ice core dust/climate signal, represents the most valuable parameter. Fine-grained single domain magnetite, probably of biogenic origin, makes a significant contribution to the interglacial MS signal, while major contributions from detrital material affect the glacial MS record. The core from the southern Scotia Sea contains significant proportions of biogenic magnetite also in glacial sediments, suggesting depositional environments different from those of the northern Scotia Sea. Our data suggest low contributions of high-coercive minerals to the overall magnetic intensity of glacial and interglacial Scotia Sea sediments, which excludes dust as a main source of the magnetic signal. Opal content can be used to distinguish between cold and warm intervals for the past 300 thousand years. Abundance fluctuation patterns of diatom species Fragilariopsis kerguelensis and Eucampia antarctica are useful stratigraphic tools for periods back to Marine Isotope Stage (MIS) 6. The Mono Lake geomagnetic excursion is identified in Scotia Sea sediments for the first time. Possible correlations of ash layers are suggested between Scotia Sea sediments and East Antarctic ice cores. They have potential to serve as additional age markers for further studies in this area. © 2015 Elsevier B.V. All rights reserved.

Keywords: Scotia sea Magnetic susceptibility AMS 14C Geomagnetic paleointensity Diatom stratigraphy Tephrochronology

1. Introduction Stratigraphic age assignment is a crucial prerequisite of paleoceanographic studies to place the paleoceanographic history in a time frame at the highest possible accuracy, and as a baseline for

* Corresponding author. Alfred-Wegener-Institut Helmholtz-Zentrum für Polarund Meersforschung, Bremerhaven D-27568, Germany. E-mail address: [email protected] (W. Xiao). http://dx.doi.org/10.1016/j.quageo.2015.11.003 1871-1014/© 2015 Elsevier B.V. All rights reserved.

precise correlation of marine records among each other and with climate records obtained from land and ice cores. In Southern Ocean sedimentary records south of the Antarctic Polar Front (APF), the application of the well-established oxygen isotope stratigraphy (e.g., Lisiecki and Raymo, 2005) and radiocarbon dating based on calcareous microfossils is hampered by the lack of continuous planktonic and benthic foraminifera records for most of the sedimentary basins. Alternatively, radiocarbon chronology has been accomplished by dating the bulk acid insoluble organic carbon (AIO) (e.g., Domack et al., 2001; Pugh et al., 2009; Hillenbrand et al., 2010; Collins et al., 2012a) or the humic acid and residue fractions of organic carbon (e.g., Bianchi and Gersonde, 2004; Denis et al.,

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2009). However, application of radiocarbon chronology is restricted to the past 40e50 ka (ka: 1000 years) and can be complicated in Southern Ocean sediments by temporal and spatial changes of the marine carbon reservoir effect (Sikes et al., 2000; van Beek et al., 2002; Skinner et al., 2010). Contamination with 14C-inactive fossil carbon, and with modern 14C during processing opal-rich sediments (Zheng et al., 2002) may further disturb down-core radiocarbon profiles. Alternative methods have been developed successfully to generate reliable chronologies for Southern Ocean late Quaternary sediment cores. This includes down-core variability patterns of magnetic susceptibility (MS) having the potential to delineate glacial/interglacial intervals (e.g., Bareille et al., 1994; Diekmann et al., 2000). This is specifically useful for the Scotia Sea, showing close resemblance of MS variations with Antarctic ice core dust records and allowing the transfer of ice-core age models to marine sediment records (Pugh et al., 2009; Weber et al., 2012). The increased occurrence of ferromagnetic minerals in ash layers resulting in distinct MS-peaks can be used for exact ash layer correlation at regional scale. However, volcanic shards are difficult to separate from the surrounding sediment and thus prevent independent 40Ar/39Ar dating (Moreton, 1999). Another stratigraphic method relies on the occurrence and abundance patterns of siliceous microfossils widespread and well preserved in Southern Ocean sediments. Relative abundance fluctuations of the radiolarian species Cycladophora davisiana (e.g. Morley and Hays, 1979; Brathauer et al., 2001) and diatom species Eucampia antarctica (Burckle and Cooke, 1983; Burckle and Burak, 1988; Collins et al., 2012a) display increased relative abundances in glacial intervals and have been tentatively correlated with oxygen isotope records in late Quaternary Southern Ocean sediments. Diatom or radiolarian biostratigraphic marker species represent another useful tool for age assignments of Southern Ocean sediments (e.g., Gersonde et al., 1990a,b; Zielinski and Gersonde, 2002). Geomagnetic paleointensity analysis presents another approach to generate independent age assignments in Quaternary records and allow for inter-hemispheric correlation (Laj et al., 2004). Several studies have been carried out in the Antarctic Peninsula region for the Holocene-Last Glacial Maximum (LGM) interval and back to 270 ka (Sagnotti et al., 2001; Brachfeld et al., 2003; Macri et al., 2006; Hillenbrand et al., 2010); and in Pata-Pronovost et al., 2013). A recently derived gonia back to 50 ka (Lise palaeomagnetic record in the Scotia Sea extends back to the last glacial and includes the record of the global palaeomagnetic Laschamp excursion (Collins et al., 2012a), recognized as well from Northwest Weddell Sea sediments (Grünig, 1991). These records prove this method as a useful stratigraphic tool for Southern Ocean sediments. Although different methods have been applied to establish chronologies for Scotia Sea sediments, there remains a lack of detailed inter-comparison to test and validate their consistency and reliability. We present new data obtained from continuous Scotia Sea sediment sequences, mainly composed of biosiliceous microfossil and terrigenous materials deposited at different sedimentation rates covering the last 300 ka. We combined radiometric measurements (210Pb, AMS 14C), records of MS, biogenic opal content, diatom abundance fluctuations, and geomagnetic paleointensity, to test the applicability of the different methods and to distill a robust pattern of stratigraphic signal variability to millennial scale resolution. The obtained pattern was calibrated with comparable signals from well-dated sediments and ice core records. Prominent ash layers were also used for regional correlation. Possible correlations of ash layers in the Scotia Sea sediments and Antarctic ice cores are proposed.

2. Regional settings The Scotia Sea receives water from the Antarctic Circumpolar Current (ACC) via the Drake Passage and the Weddell Sea (Maldonado et al., 2003) (Fig. 1). The southern boundary of the ACC straddles across the Scotia Sea linked to the extent of the Antarctic winter sea ice (WSI) field. The associated Southern ACC Front (SACCF) is closely tied to bottom topography, similar to the APF and the Subantarctic Front (SAF) in the northwestern Scotia Sea (Orsi et al., 1995; Sokolov and Rintoul, 2009; Hogg and Munday, 2014). Blooming and export of phytoplankton starts in early spring as a result of increased water column stability induced by ice-melting processes (Fryxell and Kendrick, 1988). Phytoplankton distribution patterns are strongly affected by the frontal system, ice edge and the grazing pressure (Treguer et al., 1991; Socal et al., 1997). The interplay of bottom topography, water mass circulation pattern and biological export production exert strong control on the deposition of sediments. The atmospheric circulation in the Scotia Sea is dominated by the Southern Hemisphere Westerlies covering a wide latitudinal range of 40e70
S and being strongest at 45e50
S (Toggweiler and Russell, 2008), transporting dust as well as volcanic ash particles eastwards. The central Scotia Sea is isolated from major sources of sediment supply from continental margins (Maldonado et al., 2003), but receives terrigenous sediment by ice rafting, ocean currents and eolian input (Diekmann et al., 2000; O'Cofaigh et al., 2001; Weber et al., 2014). 3. Materials and methods 3.1. Materials Sediment cores PS67/197-1, PS67/205-2, PS67/206-1, PS67/2191 and PS67/224-1 were collected from the Scotia Sea during R/V Polarstern cruise ANT-XXII/4 in 2005 (Schenke and Zenk, 2006); two other Polarstern cores, PS2319-1 and PS2515-3, were recovered during cruises ANT-X/5 (1992) (Gersonde, 1993) and ANT-XI/2 (1994) (Gersonde, 1995), respectively (Table 1; Fig. 1). The cores are located south of the APF on a NortheSouth transect across the central Scotia Sea. This study focuses on piston cores PS67/197-1 and PS67/219-1, both of which consist of olive to olive gray diatom ooze, olive gray to gray diatomaceous mud and diatom-bearing mud. Grains of IRD and drop stones are scattered in the southerly located Core PS67/219-1, ranging from a few millimeters to 2e3 cm in size, while less abundant in Core PS67/197-1. In Core PS67/219-1, a prominent volcanic ash layer occurs at 58e63 cm core depth, comparable to ash layers encountered in Core PS67/206-1 at 1464e1470 cm, in Core PS67/205-2 at 279e284 cm, and less pronounced in Core PS67/197-1 at 80 cm and in Core PS2319-1 at 37 cm. Dispersed ashes were also observed in other sections of the cores by microscopical analysis (Table 2). Based on core-to-core correlation (Fig. 2), the identified ash layers were numbered AS (Ash Scotia Sea) 1 through 9 as a baseline for the establishment of a regional tephrochronology. 3.2. Methods 3.2.1. 210Pb excess measurements In order to determine modern deposition, 210Pb activities were measured on uppermost sediment samples taken from piston cores PS67/197-1 and PS67/206-1. Approximately 10 g of freeze-dried sediments at 3e5 cm intervals were packed into a plastic lid, pressed and sealed in radon proof foil for 3 weeks to establish radioactive equilibrium between 226Ra and 222Rn before

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Fig. 1. Core locations (Table 1), oceanographic settings and sea ice distribution in the Scotia Sea. Cores PS67/197-1 and PS67/219-1 central to this study are highlighted. SAF: Subantarctic Front; APF: Antarctic Polar Front; SACCF: Southern ACC Front; WSI: modern winter sea ice; SSI: modern summer sea ice. Locations of oceanic fronts are from Orsi et al., 1995; sea ice distribution is from Comiso, 2003. Pathways of Weddell Sea water into the Scotia Sea are modified from Heywood et al. (2002).

Table 1 Locations of Scotia Sea sediment cores. Station

Latitude


PS67/197-1 PS67/205-2 PS67/206-1 PS67/219-1 PS67/224-1 PS2319-1 PS2515-3

55 56
57
57
57
59
53


Longitude 0




8.24 S 42.110 S 24.670 S 13.220 S 56.570 S 47.30 S 32.70 S

44 43
43
42
44
42
45


0

6.28 W 21.450 W 27.50 W 28.020 W 11.790 W 410 W 17.50 W

measurement. Concentrations of 210Pb were determined by lowlevel low-background gamma spectroscopy using a coaxial HPGe detector (Canberra Industries) at the Institute of Environmental Physics, University of Bremen. With a half-life of 22.3 years, the 210 Pb excess (210Pbex) chronology is generally valid for the recent 150 years (7 half-life cycles), and most reliable for the recent 100 years (5 half-life cycles) (Appleby, 2001). Dating of sediments by this method beyond 100 years is based more on extrapolation of a large dataset (exponential fit) than the actual individual data.

Water depth (m)

Recovery (m)

Reference

3837 3790 3206 3619 2868 4323 3467

23.31 19.27 23.62 20.71 22.01 11.54 13.07

this study this study this study this study this study Diekmann et al., 2000 Diekmann et al., 2000

3.2.2. AMS 14C dating Due to the absence or low content in foraminiferal carbonate in the studied sediment cores, 12 bulk sediment samples from Core PS67/197-1 and 16 from Core PS67/219-1 were selected for AMS 14C dating of both the humic acid and the residue fractions of organic material. This is complemented by 4 datings of Core PS67/206-1, one below the prominent ash layer AS 1. Sample preparation and dating were performed by the Poznan Radiocarbon Laboratory (Poland). For separation of the soluble humic acid fraction from the residual fraction (humines), samples were stepwise treated with

Table 2 Core depth of ash layers (Fig. 4) found in Scotia Sea cores (unit: cm).

AS AS AS AS AS AS AS AS AS

1 2 3 4 5 6 7 8 9

PS67/197-1

PS67/205-2

PS67/206-1

PS67/219-1

PS67/224-1

80 543 664 894 1580 2050

279e284 515 555 655 785 1645

1464e1470

58e63 540 623 786 998 1175 1648 1891 2003

145 185 235 495 750 1395 1625 1915

PS2319-1

PS2515-3

37 236 290 350 485 660 1082

156 260 433 658 900

100

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Fig. 2. Scotia Sea coreecore correlation based on variations of magnetic susceptibility (k). Ash layers used for correlation are labeled AS (Scotia Sea Ash) (Table 2). Horizons with single drop stones or drop stone layers are labeled D. Shaded areas mark the glacial intervals attributed to Marine Isotope Stage (MIS) 2, MIS 4, MIS 6 and MIS 8, all characterized by increased k.

chemicals. CO2 was extracted by combustion and then converted to graphite for dating (Czernik and Goslar, 2001). The 14C concentration was measured with a “Compact Carbon AMS” spectrometer (Goslar et al., 2004). The conventional radiocarbon ages were calibrated to calendar years, using the program CALIB 7.0.2 with calibration dataset Marine 13 (Stuiver and Reimer, 1993; Reimer et al., 2013) (Table 3). Although marine radiocarbon reservoir/ ventilation ages are variable through time as a result of climate related changes in ocean circulation and ventilation (Sikes et al., 2000; Skinner et al., 2010), we preferred to apply a constant reservoir age of 1300 years (DR ¼ 900 years) to all samples. Such reservoir age was proposed by previous studies including AMS 14C dating for the wider area (e.g., Domack et al., 2001; Pugh et al., 2009) considering radiometric dating of modern organisms (Berkman and Forman, 1996). 3.2.3. Magnetic susceptibility For cores PS67/206-1 and PS67/219-1, high-resolution measurements of volume-specific magnetic susceptibility (k) were performed at 1-cm intervals on split half core sections using a GEOTEK multisensor core logger at AWI-Bremerhaven utilizing a Bartington MS2F spot sensor. The other cores were measured on board R/V Polarstern prior to core splitting with an MS2C coil sensor. In addition, k was also determined for discrete samples of cores PS67/197-1 and PS67/219-1 at 5-cm intervals prepared for geomagnetic paleointensity analysis at the University of Bremen (see below). k is presented as dimensionless SI units. 3.2.4. Biogenic opal content Biogenic opal content was measured at 10-cm intervals on Core PS67/219-1 according to Müller and Schneider (1993), and afterwards corrected for the salt content of dried samples calculated from pore water content after Kuhn (2013). Because the sediments

are represented by a two component system with high (lithogenic, 2.7 g/cm3) and low (opal, 2.2 g/cm3) grain densities (Weber, 1998), a linear correlation between opal and grain density (GD) was found for Core PS67/219-1: opal ¼ 387.45143.48*GD (R2 ¼ 0.89) (Suppl. Fig. 1). This relationship was applied to results of GD measurements on powdered samples at AWI-Bremerhaven in order to calculate the opal content in Core PS67/197-1 at 10-cm intervals. 3.2.5. Light microscopic diatom analyses For quantitative diatom analysis, subsamples were taken at 10cm intervals from cores PS67/197-1 and PS67/219-1, and freezedried for diatom slides preparation. The sample treatment and slide preparation followed the standard procedure developed at AWI-Bremerhaven (Gersonde and Zielinski, 2000). The slides were examined with a Zeiss light microscope at a magnification of 1000  . The counting method is according to Schrader and Gersonde (1978). For taxonomy we followed Zielinski and Gersonde (1997, 2002). More than 400 specimens were counted for each sample for statistical reliability. Relative abundances (%) of each species and total diatom abundances (valves/gram of dry sediment) were calculated. 3.2.6. Paleomagnetic measurements Samples for paleomagnetic investigations were taken at 5-cm intervals for cores PS67/197-1 and PS67/219-1 with 2.2 cm  2.2 cm  1.8 cm plastic cubes. The discrete samples were analyzed at the paleomagnetic laboratory at the Department of Geosciences, University of Bremen. Palaeomagnetic directions and intensities of natural remanent magnetization (NRM), anhysteretic remanent magnetization (ARM) generated in a peak alternating field of 100 mT and a biasing DC field of 40 mT as well as isothermal remanent magnetization (IRM) generated in a DC field of 100 mT

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Table 3 Radiocarbon datings from cores PS67/197-1, PS67/206-1 and PS67/219-1. The calibration program CALIB 7.0.2 (Stuiver and Reimer, 1993; Reimer et al., 2013) was used to convert the 14C ages to calendar ages with 2 sigma precision, applying a reservoir correction of 1300 years (DR ¼ 900 years). TC: trigger core; KOL: piston core. Carbon sources for dating are: H- humic acid fraction; R-residue fraction. Core

Depth (cm)

Lab ID

C Source

14

C age (yrs BP)

Error (yrs)

Lower cal range (cal. yrs BP)

Upper cal range (cal. yrs BP)

Median probability (cal. yrs BP)

S67/197-1 TC

6e9 6e9 96e99 96e99 0e4 0e4 80e83 80e83 200e203 200e203 270e273 270e273 329e332 329e332 470e473 470e473 542e545 542e545 663e666 663e666 718e721 718e721 780e783 780e783 6e9 6e9 0e5 0e5 1470e1473 1470e1473 2358e2362 2358e2362 1e5 1e5 82e86 82e86 1e5 1e5 63e66 63e66 91e94 91e94 131e134 131e134 190e193 190e193 240e243 240e243 275e278 275e278 321e324 321e324 381e384 381e384 471e474 471e474 541e544 541e544 624e627 624e627 651e654 651e654 721e724 721e724

Poz-32666 Poz-32810 Poz-32674 Poz-32812 Poz-32667 Poz-32738 Poz-32672 Poz-32813 Poz-32668 Poz-32739 Poz-32669 Poz-32740 Poz-32670 Poz-32742 Poz-32671 Poz-32849 Poz-32693 Poz-32743 Poz-32694 Poz-32744 Poz-32798 Poz-32811 Poz-32799 Poz-32850 Poz-31350 Poz-31351 Poz-31290 Poz-31291 Poz-31333 Poz-31335 Poz-31347 Poz-31349 Poz-32802 Poz-32851 Poz-32805 Poz-32872 Poz-32807 Poz-32873 Poz-32866 Poz-32816 Poz-32806 Poz-32817 Poz-32801 Poz-32852 Poz-32803 Poz-32888 Poz-32804 Poz-32814 Poz-32853 Poz-32855 Poz-32856 Poz-32857 Poz-32859 Poz-32860 Poz-32861 Poz-32889 Poz-32862 Poz-32863 Poz-32864 Poz-32865 Poz-32867 Poz-32869 Poz-32870 Poz-32871

H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R H R

4930 17,390 6650 9010 4470 7260 10,310 18,230 13,590 16,290 13350 16,720 8630 15,010 16020 17,700 17370 20,580 16950 21,100 18270 21,660 21400 20,500 2435 6660 2360 3445 8195 10,300 11990 15,190 2660 4160 7060 7780 5730 6520 9140 12,570 9880 10970 11,770 13,490 10,750 11,480 12,590 14740 13,500 10,700 15,620 17,970 15,420 18,020 16,650 22,390 17,250 20,550 20,930 23,420 21,510 21,550 22,610 25,370

40 90 50 80 40 40 70 90 130 80 110 80 60 90 160 140 130 120 130 200 150 130 160 190 35 50 35 35 35 70 120 90 35 40 50 50 50 50 80 60 70 60 60 90 60 60 70 80 110 70 100 140 120 150 130 180 250 360 360 270 210 170 230 290

3909 19,137 6017 8391 3360 6715 10,123 20,132 13,893 17969 13,652 18,506 8041 16,250 17,522 19,449 19008 22,894 18,629 23,332 20,065 24,099 23,751 22617 978 6025 920 2119 7665 10,108 12,268 16,477 1244 2952 6462 7285 4963 5905 8553 12,991 9495 10,872 12,120 13,821 10,616 11,461 12,993 15,896 13,814 10545 17,110 19,725 16,764 19,770 18,321 24,985 18,723 22,422 22,738 25,876 23,787 23,914 25,142 27,661

4178 19,655 6274 8848 3564 6941 10,463 20,654 14,914 18427 14,190 18,848 8330 16,888 18308 20,144 19711 23,555 19,210 24,321 20,848 24,946 24,545 23596 1182 6281 1106 2310 7843 10,459 12,853 17,125 1383 3205 6722 7488 5277 6160 8985 13,306 9878 11,208 12,579 14,464 11,008 11,510 13,332 16,434 14,640 10967 17,719 20,481 17,526 20,552 18,869 25,819 19,868 24006 24,454 27,062 24,958 24,897 26,022 28,683

4047 19,408 6172 8586 3462 6828 10,259 20,411 14,320 18210 13,931 18,685 8201 16,545 17,906 19,787 19378 23,220 18,899 23,824 20,458 24,462 24,162 23128 1094 6184 1005 2220 7755 10,246 12,608 16,820 1307 3084 6593 7393 5128 6020 8769 13,163 9653 11,072 12,391 14,087 10,790 11,863 13,179 16,166 14,117 10722 17,436 20,111 17,170 20,172 18,609 25,424 19,253 23,190 23,621 26,380 24,297 24,333 25,618 28,136

PS67/197-1 KOL

PS67/206-1 TC PS67/206-1 KOL

PS67/219-1 TC

PS67/219-1 KOL

102

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were measured on a cryogenic magnetometer (model 2G Enterprises 755HR). NRM was measured on each sample before it was subjected to a systematic demagnetization treatment involving 16 steps for each sample starting with an alternating field (AF) of 5 mT to a maximum field of 100 mT. A detailed vector analysis was applied to the results of NRM demagnetization (Kirschvink, 1980) in order to determine the characteristic remanent magnetization (ChRM). Samples showing no systematic demagnetization pattern were rejected. In addition, 3 samples from each core were selected to measure their S-ratio and hard IRM (HIRM). These samples were taken from intervals of low and high k, representing interglacial and glacial intervals. IRM acquisition curves in DC fields up to 700 mT were generated using the in-line pulse magnetizer of the 2G Enterprises cryogenic rock magnetometer. For higher magnetic fields up to 2.7 T, a 2G Enterprises Model 660 external pulse magnetizer was used. After applying a magnetic field of 2.7 T, the remanence (IRM300mT) imparted in a reverse field of 300 mT was determined. In turn, HIRM was calculated as the difference between the remanence acquired in a magnetic field of 2700 mT and that acquired in a magnetic field of 300 mT. S-ratio (S0.3T) was calculated as S0.3T ¼ (1-IRM0.3T/IRM2.7T)/2 (Bloemendal et al., 1992). To characterize magnetic mineral assemblages, First-Order Reversal Curves (FORCs) (Pike et al., 1999; Roberts et al., 2000) were conducted on 2 samples from Core PS67/197-1 and 6 from Core PS67/219-1, using an alternating gradient magnetometer (Princeton Measurement Corporation, Micromag Model 2900). FORCs were measured with coercivity fields (Hc) between 0 and 120 mT, and local interaction fields (Hu) between 40 and 40 mT. We applied an averaging time of 1s and a field increment of 2 mT up to a saturation field of 1T. FORC diagrams were analyzed using the software of Harrison and Feinberg (2008) with a smoothing factor of 6. In a FORC diagram, spreading of the distribution along the vertical Hu axis corresponds to magnetostatic interactions for single-domain (SD) or internal demagnetizing fields for multidomain (MD) grains. Usually, closed structures parallel to the horizontal Hc axis with pronounced peaks are characteristic of SD grains. Contours parallel to the Hu axis indicate increasing particle sizes. 3.2.7. Correlation of sedimentary records Detailed inter-core correlation can be established by correlating the fluctuation patterns of k, biogenic opal content, specific diatom species composition, and the occurrence of ash layers. The strong inter-core correlation indicates consistent sedimentation pattern for this area. 4. Results 4.1.

210

Pbex determination

The supported 210Pb activity in Core PS67/197-1 is variable. Pbex was not determined in the sediment surface but at 7 cm core depth, possibly due to bioturbation of the sediments. Presumably the core top sediments are not recent, as also suggested by old AMS 14C ages (Table 3; Fig. 3). In contrast, the activity concentrations in Core PS67/206-1 are relatively stable. The clear presence of the 210Pbex signal indicates modern deposition at the core top. Assuming that the 210Pbex covers a time interval of about 100 years (top 20 cm), the 210Pbex dating would yield a sedimentation rate of about 2 m/ka by linear fit, similar to the core average sedimentation rate calculated by radiocarbon ages from this core. 210

4.2. Radiocarbon dating Radiocarbon ages derived from the soluble humic acid fraction are generally younger than ages obtained from the insoluble residue (Table 3; Fig 4). The humic acid fraction radiocarbon dates obtained from the core tops at Site PS67/206 (~2400 yrs BP) are similar to the trigger core top at the nearby Site PS67/219 (2660 yrs BP). The surface residue 14C ages of both cores range from 3445 to 6660 yrs BP. This indicates strong and variable contamination by fossil carbon preserved in the residue fraction. The age difference between the two fractions ranges in the southern cores PS67/ 206-1 and PS67/219-1 between 1085 and 4225 years (average 2654 years) and 40e5740 years (average 2245 years), respectively, and in the northern Core PS67/197-1 between 1680 and 12,460 years (average 4126 years). Exceptions are present in two samples where the humic acid fraction ages exceed those of the insoluble residue. They are glacial-age sample from Core PS67/197-1 (780e783 cm) by 900 years and late termination-age sample from Core PS67/219-1 (275e278 cm) by 2800 years. The surface residue ages are always more than one thousand years older than the humic acid ages and vary at different locations. In the southern Core PS67/219-1, the core top age difference between the two fractions is less than 1000 year and increases to a maximum at about 9 ka humic acid age; the age differences remain between 1000 and 2000 year until about 15 ka humic acid age and deviate significantly until 22 ka humic acid age. For the northern Core PS67/197-1, significantly old surface humic acid age (older than 4 ka) was obtained in both trigger and piston cores. There is also a remarkable rise of age differences (up to 8000 years) between the two fractions at 10 ka humic acid age. The ages start to deviate at about 16 ka humic acid age, large differences remain until about 21 ka humic acid age. In both cores, at 21e22 ka humic acid age, the ages dated on humic acid and residue fractions show the smallest deviation. In general, the down-core pattern of the younger humic acid ages shows a more constrained age-depth relationship, while that of the insoluble residue displays stronger scatter (Table 3; Fig. 4). The older ages and the less stable age-depth relationship of the insoluble residue fraction are interpreted to indicate stronger and variable contamination by reworked old carbon. As the humic acid ages represent better constrained age-depth relationship and regional consistency, we assumed the humic acid ages better represent the true 14C ages and thus were used for 14C chronology. This is the same case for Bouvet Island area in the Southern Ocean Atlantic sector, where the humic acid ages are comparable to those dated on foraminifers (Bianchi and Gersonde, 2004). The variation of reservoir effect in the Holocene and Last Glacial is not well understood in this area, and the contamination by relic carbon cannot be quantified down-core at the moment. We applied a constant correction of 1300 years as normally applied in this area for the calculation to calendar ages (Table 3). As suggested by 210 Pbex from Core PS67/206-1, this may introduce about 1000 years error to the core top sediments, while the error for the deeper sediment remains unresolved. 4.3. Magnetic susceptibility age model The k records from the Scotia Sea show pronounced patterns of low and high values (Fig. 2), corresponding to units of olive gray diatom ooze and gray diatomaceous to diatom bearing mud, respectively. Several sharp peaks coincide with ash layers and drop stones observed visually or under the microscope. Drop stones were especially recognized in the southern Core PS67/219-1, indicating the southern origin of IRD reaching the core site which is

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Fig. 3. Total 210Pb and 214Pb activity and 210Pb excess (210Pbex) of uppermost sediments from piston cores PS67/197-1 and PS67/206-1. Dashed lines mark zero excess of 210Pb. Bars indicate the analytical errors of 210Pbex. Arrows indicate core top 14C ages from humic acid fraction (Table 3).

located in the “iceberg alley” crossed by icebergs calving from Antarctica (Anderson and Andrews, 1999). Across the northesouth transect of the Scotia Sea, all cores show consistent patterns of k variability, allowing us to develop a regional correlation over the area based on these variations, showing higher values in the northern cores than in the southern cores. The k signals of our cores show remarkable similarities to variations of dust concentrations and climate variability recorded in Antarctic ice cores (Fig. 5). Following previous studies (Pugh et al., 2009; Weber et al., 2012), we correlate the k to the Antarctic ice core dust and temperature records. k of Core PS67/197-1 and of the upper 15 m of Core PS67/219-1 were correlated to the EPICA Dronning Maud Land (EDML) ice core atmospheric non-sea-salt

103

Ca2þ (nssCa2þ) concentration record, as an indicator of dust emission from Patagonia (Fischer et al., 2007); and to the d18O record (EPICA members, 2006), which represents high resolution climate changes linked to the Southern Ocean Atlantic sector. The lower part of Core PS67/219-1 was correlated to the EPICA Dome C (EDC) ice core dust concentration record (Lambert et al., 2008) and to the deuterium record (Jouzel et al., 2007) for climate variations on larger time scales. High k values correspond to glacial/stadial intervals with high dust deposition as recorded in Antarctic ice cores. Low k values reflect interglacial/interstadial intervals with low dust deposition. The AnalySeries software (Paillard et al., 1996) was applied for graphic correlation. Correlation of k to ice core records shows that Core PS67/197-1 extends back to about 86 ka and Core PS67/219-1 to about 300 ka. Core PS67/197-1 shows much higher resolution than the upper 13 m of Core PS67/219-1, representing the same age interval (Fig. 5). The distinct linear correlation between k of Core PS67/197-1 and the EDML nssCa2þ, based on 14 tie points, is reflected by the high correlation coefficient (r) of 0.89. 29 tie points were used to correlate k of Core PS67/219-1 to EDML/EDC dust proxies, resulting in a correlation coefficient of 0.83. The correlation weakens during peak interglacials which are characterized by low signals for both k and ice core dust proxies. Due to their pronounced correlation to Antarctic ice core records, age models inferred from k were used as initial points for core stratigraphies (MS chronology). To estimate the age uncertainties of these age models with respect to the selected tie points, 500 replications were conducted by nonparametric bootstrap re-sampling (Mudelsee, 2014). The bootstrap method generates artificial resamples of the time series, repeats for each resample the estimation and calculates the error bars from the distribution of the replications. The 1-sigma standard deviations of the estimates were considered as probable age uncertainties. For Core PS67/197-1, the age uncertainties are within 1 ka. For Core PS67/219-1, the age uncertainties are lower than 1 ka for the upper core section representing Marine Isotope Stage (MIS) 1e3. Larger uncertainties of up to 2e6 ka occur in the lower section especially for MIS 6e7 sections. The larger uncertainties at this interval are due to the re-sampling process of bootstrap, which could bias the interpolated results at intervals of few tie points. The age-depth model derived from the linear interpolation is well within the uncertainty range.

Fig. 4. Radiocarbon datings of cores PS67/197-1, PS67/219-1 and PS67/206-1. Bars indicate age errors (Table 3). Ash layer intervals in each core are labeled AS (Table 2).

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4.4. Biogenic opal content The biogenic opal contents of cores PS67/197-1 and PS67/219-1 vary from 7 to 46%, and 4e64%, respectively. They show strong variations clearly anti-correlated to the respective k, peaking at interglacials/interstadials, while low content characterizes glacial/ stadial intervals (Fig. 6). This also reflects the dilution effect on the overall MS signal by biogenic opal deposition. 4.5. Diatom stratigraphy The initial biostratigraphies of cores PS2319-1 and PS2515-3 were established by relative abundance fluctuations of radiolarian species C. davisiana, with high values during glacial times (Diekmann et al., 2000) (Suppl. Fig. 2). PS2319-1 extends back to about 200 ka and PS2515-3 to about 150 ka. These age models agree with the MIS stages as defined by MS chronology. They help to constrain the basic stratigraphic frames of other cores by regional core correlation. In this study, we used relative abundances of diatom species Fragilariopsis kerguelensis and Eucampia antarctica, and diatom concentration for diatom fluctuation stratigraphy (Fig. 6). In Core PS67/197-1, relative abundances of F. kerguelensis range from 30 to 75%; strong fluctuations of the species covary with the opal content and are anti-correlated to the k signal. This identifies F. kerguelensis as the main contributor to the circum-Antarctic opal flux, representing the open ocean sedimentation regime north of the sea ice extent (Zielinski and Gersonde, 1997; Abelmann et al., 2006). Eucampia antarctica abundances range from 0 to 30%, peak values occur at intervals with low F. kerguelensis abundance. The diatom concentration fluctuates between 106 and 2  107 valves/gram, generally covarying with opal content. In Core PS67/219-1, abundance of F. kerguelensis ranges from 10 to 60%, with fluctuations comparable to the MS record for the upper 12 m, but demonstrating different patterns in the lower part. Eucampia antarctica abundance ranges between 0 and 45%, with temporal maxima in MIS 2, 4 and 6. Two stratigraphic marker species occur in the lower part of this core. The last occurrence (LO, 1%) of Rouxia leventerae is at 1420 cm core depth, with an estimated age of 137.7 ± 4.1 ka according to the MS chronology. This age matches previous investigations of its last occurrence datum (LOD) at ~130e140 ka in the eastern South Atlantic (Zielinski et al., 2002). At core depth 1670 cm, the LO of another biostratigraphic marker Hemidiscus karstenii is at about 189.2 ± 6 ka. This is also in agreement with the literature (~190e200 ka, Zielinski and Gersonde, 2002) and roughly marks the boundary of MIS 6 and MIS 7. The diatom concentration ranges between 8  106 and 1.8  107 valves/ gram. Eucampia antarctica has been regarded as a biostratigraphic marker for the late Quaternary Southern Ocean (Burckle and Cooke, 1983; Burckle and Burak, 1988). This species is most abundant in neritic-like environments, and linked to sea ice conditions (Burckle, 1984). The E. antarctica relative abundance peaks were correlated to glacial intervals with maximum sea ice extent. A detailed E. antarctica stratigraphy can be established for Core PS67/197-1 for MIS 1e4 and well in line with the top 12 m of Core PS67/219-1. The lower part (older than MIS 4) of the diatom stratigraphy is based on Core PS67/219-1 with relatively lower resolution. The MIS 1e6 E. antarctica stratigraphy is in agreement with the C. davisiana stratigraphies in cores PS2319-1 and PS2515-3 (Diekmann et al., 2000). However, continuous low occurrence of E. antarctica in the intervals older than MIS 6 in Core PS67/219-1 weakens its stratigraphic significance. This has been also recognized for the Kerguelen Plateau region in the

Southern Ocean Indian sector (Kaczmarska et al., 1993), where the E. antarctica maximum is poorly developed in MIS 8. Furthermore, this interval does not show clear patterns for F. kerguelensis (Fig. 6). The variations of diatom species (and also of k) of Core PS67/ 197-1 can be well correlated to those of the upper 12 m of Core PS67/219-1 (Fig. 6). The higher occurrence of E. antarctica, and the lower occurrence of F. kerguelensis in the southern core compared to the northern core suggest more open ocean conditions in the northern Scotia Sea, and more neritic conditions thus sea ice influence in the south (Zielinski et al., 1997). Diatom concentrations of both cores generally covary with opal content and k signal (Fig. 6). In our records, despite the indistinguishable E. antarctica in the lower part of Core PS67/219-1, MIS 7 and 8 can be recognized by diatom concentration/opal content signals. The concentration spikes (Collins et al., 2012a), due to the dominance of Chaetoceros resting spores, were not found in our records. Beside this unusual spike, our records suggest identification of cold/warm intervals by diatom concentration in a plausible way. 4.6. Paleomagnetic directional data and relative paleointensity (RPI) Complete removal of possible viscous components or magnetic overprints was achieved generally at low demagnetizing fields (<20 mT) (Suppl. Fig. 3). The obtained demagnetization data were analyzed by principal component analysis (PCA). The maximum angular deviation (MAD) values for ChRM of both cores are below 4
, indicating a well-defined magnetization component (Stoner and St-Onge, 2007). Inclination of the ChRM of Core PS67/197-1 generally oscillates from 88
to 11
with a mean inclination of 70.4
(a95 ¼ 1.7
), matching the expected inclination of a geomagnetic axial dipole (GAD) of 70.8
at the core site (Suppl. Fig. 4). The inclination shallowing trend of Core PS67/197-1 from core top to the bottom is probably attributed to the coring process and compaction of sediments towards the core base (Arason and Levi, 1990). A clear inclination reversal and a major excursion occur at 1014 cm (11.3
) and 1299 cm (2.9
), respectively. The mean relative declination of ChRM of the core is 106.3
, but strong scatter exists between 400 and 1000 cm core depth (Suppl. Fig. 4). Since there is no clear indication for a change in their magnetic properties, the large declination scatter at this interval probably reflects changes in physical properties of the bulk sediments rather than of the magnetic fraction. Nevertheless, for a geocentric axial dipole at our high-latitude core sites, the vertical component of the magnetic remanence vector (inclination) is of almost threefold the magnitude of the horizontal and thus should weigh more than the horizontal component (declination). In Core PS67/219-1, the inclination of ChRM generally fluctuates between 86
to 22
with a mean value of 63.1
(a95 ¼ 1.3
), suggesting that the core suffers from inclination shallowing considering an expected inclination of 72.2
of the GAD at the site latitude. Besides sediment compaction, high water content and bioturbation effects (Egli and Zhao, 2015) at the time of deposition, especially during warm intervals, may also contribute to the stronger shallowing of ChRM in the sediment of this core (Suppl. Fig. 4). No geomagnetic reversal was found in this core. Its (relative) ChRM declination shows much more uniform values than that of Core PS67/197-1, with an averaged value of 137.2
. The shallow inclination at the core top of both cores may be ascribed to the reorientation of magnetic particles during or after core recovery (Hillenbrand et al., 2010). The intensities of NRM, ARM100mT, IRM100mT of Core PS67/197-1 generally range from 8 to 101 mA/m, 30e250 mA/m, and 1e16 A/m,

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Fig. 5. Magnetic susceptibility records (k) of cores PS67/197-1 and PS67/219-1 correlated to Antarctic ice core dust and temperature records, ash layers and main tie points (dashed lines) are shown. EDML nssCa2þ from Fischer et al. (2007), d18O from EPICA members (2006); EDC dust concentration from Lambert et al. (2008), deuterium record from Jouzel et al. (2007). Intervals of Vostok, Dome F and EDC ash layers are labeled V, DF and EDC, respectively (Basile et al., 2001; Kohno et al., 2004; Narcisi et al., 2005).

respectively (Suppl. Fig. 4). Magnetic grain sizes were estimated by the ratio of anhysteretic susceptibility (kARM) to volume specific susceptibility (k), kARM/k (King et al., 1983) and ARM/IRM100mT (ARM and IRM imparted at fields of 100 mT, respectively, (Maher, 1988)). kARM/k and ARM/IRM100mT show the same fluctuation patterns and range from 3 to 21, and 0.013 to 0.050, respectively

(Suppl. Fig. 4). A remarkable increase in intensity of all concentration dependent parameters is found for core depths 400e700 cm (MIS 2), and 880e1000 cm (30.8e33.7 ka). This increase is accompanied with a coarsening of the magnetic particles as indicated by decreasing kARM/k and ARM/IRM100mT ratios. This coincides with the relevant declination scatter in this interval,

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Fig. 6. Comparison of magnetic susceptibility, opal concentration, diatom valve concentration and abundance pattern of prominent diatom species as recorded in cores PS67/197-1 and PS67/219-1. The age of last occurrence (LO, 1%) of the diatom biostratigraphic species R. leventerae and H. karstenii are according to the MS chronology. Opal content of Core PS67/197-1 is calculated from grain density (GD) using the opal-GD relationship established for Core PS67/219-1 (Suppl. Fig. 1).

indicating environmental changes affecting (not only) magnetic particles. The core interval from 300 to about 2050 cm (MIS 2 to 4) is characterized by overall lower kARM/k and ARM/IRM100mT ratios than the units above and below this section. Nevertheless, ARM100mT and IRM100mT vary by less than one order of magnitude. Correlation coefficients between k and ARM100mT, and k and

IRM100mT are 0.45 and 0.95, respectively. Being a measure of stability (coercivity) of the magnetic remanence, the median destructive field of NRM (MDFNRM) ranges from 5 to 40 mT (mean 24 mT). The MDFARM ranges between 20 and 33 mT (mean 27 mT) (Suppl. Fig. 4), indicating that ARM100mT is carried by a low-coercive mineral like magnetite (Sagnotti et al., 2001). In summary, the

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magnetic mineral properties of Core PS67/197-1 meet the requirements for sediments suitable for paleointensity investigations (Tauxe, 1993). Mean magnetic grain sizes are smaller than 1 mm (Fig. 7). Larger grains characterize the glacial interval of MIS 2e4 while interglacial intervals are characterized by grains smaller than 0.1 mm, particularly in MIS 1 and 5. The narrow ridge along the Hc axis in the FORC diagram (129 cm) suggest a major fraction of noninteracting stable SD (rather than superparamagnetic) particles in the interglacial (MIS 1) sediments (Egli et al., 2010; Channell et al., 2013). In contrast, the high intensities along the Hu axis combined with low Hc values of the glacial sample (1244 cm, MIS 3) suggest the presence of pseudo single-domain (PSD) or multi-domain (MD) detrital magnetite (Fig. 7). In Core PS67/219-1, intensities of NRM, ARM100mT, IRM100mT range from 2 to 104 mA/m, 9e300 mA/m, and 0.3e18 A/m, respectively (Suppl. Fig. 4). The kARM/k and ARM/IRM100mT ratios range from 2 to 45, and 0.001 to 0.027, respectively. MDFNRM and MDFARM range from 4 to 35 mT (mean 24 mT), and 11e36 mT (mean 26 mT) (Suppl. Fig. 4), respectively, similar to Core PS67/ 197-1. The variabilities of the concentration-dependent magnetic parameters of this core, expressed as standard deviations, are higher than those of Core PS67/197-1. The correlation coefficients (r) between k and ARM100mT, and k and IRM100mT are 0.79 and 0.93, respectively. Magnetic grain sizes in this core (Fig. 7) are generally smaller than in Core PS67/197-1 for MIS 1e5 and MIS 7, almost all the samples show mean grain sizes smaller than 0.1 mm. Larger grains occur in glacial intervals of MIS 6 and 8. FORC diagrams suggest a dominance of SD particles in the interglacial sample (1323 cm, MIS 5), and increased contributions of PSD/MD particles in the glacial sample (558 cm, MIS 2). Relative paleointensity (RPI) of both cores were calculated using the so-called ‘slope-method’ or pseudo Thellier method (Tauxe et al., 1995; Channell et al., 2002). RPI was computed as the slope of the regression line of NRM intensities plotted versus the intensities of ARM100mT and IRM100mT for AF demagnetization levels 25e50 mT (RPIARM ¼ NRM25e50mT/ARM25e50mT and RPIIRM ¼ NRM25e50mT/IRM25e50mT). The results were standardized as zero mean with standard deviation 1 (Suppl. Fig. 4; Fig. 8). For correction of an eventual bias to RPI caused by relevant magnetic grain size variations, we applied a correction factor following -Pronovost previous studies (Brachfeld and Banerjee, 2000; Lise et al., 2013): RPI' ¼ RPI * MDFNRM-CM/MDFNRM, where MDFNRM-CM is the center of mass of MDFNRM. The resulting RPI’ARM and RPI’IRM of both cores show fluctuation patterns similar to the original data, but show excursions as significant lower RPI0 at core depths of 1984 cm in Core PS67/197-1; and 938 cm, 1658 cm, 1898 cm, 1963 cm, 2043 cm in Core PS67/219-1. Except for 1658 and 1898 cm in Core PS67/219-1 close to AS7 and AS8 at 1648 and 1891 cm (Table 2), respectively, all other intervals with suggested geomagnetic excursions are located away from ash layer intervals. Thus we consider also these RPI0 minima as being not related to the presence of ash. In addition, the paleomagnetic parameters and demagnetization diagrams from samples at or next to the ash layers do not show significant differences to those from adjacent samples. Thus RPI0 anomalies at 1658 and 1898 cm may reflect actual changes in geomagnetic field. Significantly low RPI0 values overlap with low MDFNRM values (Suppl. Fig. 4), reflecting the overall weakening of the geomagnetic intensity during excursion periods. For five out of six representative samples (Table 4), S-ratios vary in a narrow range between 0.97 and 0.98 indicating ratios of low- (e.g. magnetite) to high-coercive (e.g. hematite) minerals of about 20%e80% according to Bloemendal et al. (1992). Only the Holocene sample from Core PS67/219-1 (158 cm)

107

demonstrates a lower S-ratio of 0.95 indicating slightly lower low-to high-coercive mineral concentration of about 15%e85%. HIRM amounts to about 7% of IRM2.7T for the Holocene sample of Core PS67/219-1, while for the other 5 samples HIRM is only about 3% of IRM2.7T. Core PS67/197-1 shows almost no differences in low- and high-coercive mineral composition between glacial and interglacial periods, and for PS67/219-1 such variability is relatively weak. 5. Discussion 5.1. Reworked old carbon and radiocarbon calibration In both cores PS67/197-1 and PS67/219-1, the residue ages generally show much older and stronger down-core variation than the humic acid ages. This clearly points to strong and variable fossil carbon contamination in the residue fraction. Ages obtained from bulk acid-insoluble organic (AIO) material (same as the residue fraction in our dating) have been used to establish radiocarbon chronology around Antarctica in previous studies (e.g., Domack et al., 2001; Pudsey et al., 2006; Pugh et al., 2009; Hillenbrand et al., 2010; Collins et al., 2012a). However, most of the Antarctic surface sediments show significantly old ages ranging from less than 2 ka to more than 10 ka if dated on bulk AIO, even some recognized as modern deposition by 210Pb dating (Andrews et al., 1999; Pudsey et al., 2006). In addition, down-core profiles of AIO ages are often not in stratigraphic order (e.g., Pugh et al., 2009; Collins et al., 2012a), as the case of residue ages in our cores. Reworking of old carbon is generally considered as the reason for the older surface ages. The application of AIO ages for the age model requires the following prerequisites (Licht et al., 1998): 1. relatively high (low) influx of biogenic (terrigenous) material; 2. little reworked old carbon. In regions where reworking of old carbon is unavoidable, a local contamination offset of the fossil carbon age is applied, by assuming that the contamination is constant through time (e.g., Hillenbrand et al., 2010). However, this assumption may not hold true in the Antarctic Peninsula region and the Scotia Sea. The potential sources of fossil carbon are the admixture of old sediments by bottom currents, and the input of terrigenous old carbon from adjacent land. More terrigenous detritus was supplied into the area by wind and current transport as well as ice rafting during the glacial than the interglacial, in response to lowered sea level and extended sea ice (e.g., Clark and Mix, 2002; Gersonde et al., 2005). In the Holocene, the relic carbon for the northern site may be mainly transported by wind or wind driven currents; the southern site may also be affected by a second source from Antarctica via ice transportation and deep flow from the Weddell Sea. During the last glacial and the following termination 1, ice transportation and deep flow from the south may play an increasing important role because sea ice also expanded to the northern site (Gersonde et al., 2005; Collins et al., 2012b); moreover, the northern core may also receive increased terrigenous input from southern Patagonia due to expanded ice fields and coastal erosion at low sea level stands (Diekmann et al., 2000). Besides admixture of old sediments, the persistence of old dissolved organic carbon (DOC) in the water column (Hansell, 2013; Lechtenfeld et al., 2014) may also contribute to the older than expected 14C ages in the sediments. However, such influence may be not significant because in the Southeast Atlantic, the humic acid fraction show similar ages to foraminifera (Bianchi and Gersonde, 2004). The variations of age differences of the humic acid and residue fractions show similarities in cores PS67/197-1 and PS67/219-1. At about 9e10 kyr BP and 15e21 kyr BP humic acid ages, temporal

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Fig. 7. Plots of anhysteretic susceptibility (kARM) against magnetic susceptibility (k) for cores PS67/197-1 and PS67/219-1 with the calibration of magnetic grain size from King et al. (1983). Samples from different marine isotope stages are denoted by different colors (a, b). First-order reversal curves (FORCs) of samples representative of interglacial (c, d) and glacial intervals (e, f) from both cores. FORC diagrams constructed using a smoothing factor of 6. Hysteresis parameters Hc, Mr/Ms and Hcr/Hc of each sample are indicated.

maximum of age differences between the two fractions occur in both cores, indicating enhanced reworked carbon supply to this region during termination 1. This is accompanied by the major melting of ice sheets in (Western) Antarctica and Patagonia which  Cofaigh brought terrigenous detritus to the marine sediments (O et al., 2014; Hillenbrand et al., 2014). The 2.3e2.4 kyr BP surface humic acid ages seem to be a regional phenomenon in the southern cores PS67/206-1 and PS67/219-1. It is about 1000 years larger than the normally assumed reservoir correction of 1300 ± 100 years in the study area (Gordon and Harkness, 1992; Berkman and Forman, 1996). This leads to the assumption that fossil carbon may also exist in the humic acid fraction. Compared to the Bouvet Island area (Bianchi and Gersonde, 2004), sediments from the Scotia Sea have higher potential receiving fossil carbon from terrigenous detritus and reworked sediments. This is shown by almost pure diatom ooze of Holocene sediment around Bouvet Island, where humic acid and residue ages are quite similar in the Holocene, and diverse towards the LGM (Bianchi and Gersonde, 2004). Due to the better consistency of humic acid ages, we propose that both humic acid and residue ages can be affected by fossil carbon, but humic acid ages may be less influenced, and thus were preferentially used for radiocarbon chronology. The age reversals occurring in our records may be also affected by the temporal changes of carbon reservoir during the termination, which is closely related to ocean ventilation (Sikes et al., 2000; Butzin et al., 2005; Skinner et al., 2010). But the extremely young humic acid age at 329e332 cm in Core PS67/1971 remains unexplained. Our data point out the weakness in accuracy of radiocarbon chronology in Scotia Sea sediments, and the necessity of regional correlations for validation of the radiocarbon ages.

5.2. Correlation of Scotia Sea magnetic susceptibility signals to Antarctic ice core dust/climate records Similar to other Scotia Sea records (Pugh et al., 2009; Weber et al., 2012), the strong correlation between k of the studied cores and Antarctic dust/climate records suggests a close link between these two archives. The k signals show clear glacial/interglacial cycles for the Scotia Sea with high values in glacial periods and low values in interglacial times. This indicates a strong supply of magnetic particles during glacial times. Patagonian dust has been considered as a main contributor to Scotia Sea MS. The rapid transport of dust from Patagonia to the South Atlantic and to Antarctica (Li et al., 2010) is assumed to be the reason for the strong coupling between the physical properties of Southern Ocean sediments, e.g. MS, and Antarctic dust records (Weber et al., 2012). The mean MDFARM of our records are close to the value of dust (~30 mT) (Egli, 2004). However, the MDFARM is rather an expression of relative proportions of magnetic components than the indication of the occurrence of a specific component. The S-ratios suggest about 80e85% of high-coercive minerals (e.g. hematite, as indication of dust) (Bloemendal et al., 1992) in both interglacial and glacial sediments, however they account for only 3e7% of the total IRM2.7T (Table 4). This suggests a dominant control of magnetic intensity (and susceptibility) by low coercive mineral (e.g. magnetite). Strong correlations were found between Antarctic dust records and MS and lithogenic/detrital flux in different Southern Ocean basins (Bareille et al., 1994; Pugh et al., 2009; Mazaud et al., 2010; Martinez-Garcia et al., 2014; Lamy et al., 2014), showing a general glacial increase of MS but pointing to different sources. In the subAntarctic Indian Ocean, the glacial high MS was previously attributed to an increased input from volcanic sources by erosion (Bareille et al., 1994). A recent study in the same region suggests biogenic magnetite to be a significant contributor to MS, which

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Fig. 8. Magnetic susceptibility (k), inclination of ChRM and RPI records of cores PS67/197-1 and PS67/219-1 according to their MS chronology, compared to reference RPI records, -Pronovost et al., 2013), SAPIS (Stoner et al., 2002), SINT-800 (Guyodo and Valet, 1999), and ODP 1089 (Stoner et al., 2003). Ages of global geomagnetic including the PASADO (Lise excursions are from Lund et al. (2006a). The estimated ages of excursions in our cores are indicated.

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Table 4 IRM acquisition measurements and S-ratios of samples from cores PS67/197-1 and PS67/219-1. Core

Depth (cm)

MIS

IRM300mT [mA/m]

IRM2.7T [mA/m]

IRM300mT [mA/m]

HIRM [mA/m]

S-ratio

PS67/197-1

154 509 1279 158 588 843

1 2 3 1 2 3

3864.3 9987.5 6089.4 507.1 11249.0 8163.9

3994.0 10335.0 6319.5 547.0 11577.0 8446.9

3836.8 9713.5 6062.6 488.0 11064.0 8020.8

129.7 347.5 230.1 39.9 328.0 283.0

0.980 0.970 0.980 0.946 0.978 0.975

PS67/219-1

increases along with the input of terrigenous components during glacials (Yamazaki and Ikehara, 2012). In our cores, FORC diagrams suggest a dominance of fine-grained magnetite of probable biogenic origin in interglacial sediments, whereas glacial sediments are characterized by increased contributions of low-coercive coarser PSD/MD magnetite of detrital origin. This is also indicated by high/low ARM100mT/IRM100mT (and kARM/k) values during interglacials/glacials (Suppl. Fig. 4), suggesting small magnetic grain sizes and/or high proportion of biogenic magnetites during interglacial, and vice versa (Verosub and Roberts, 1995; Mazaud et al., 2007; Yamazaki, 2008). Also hysteresis ratios (Mr/Ms, Hcr/Hc) show the interglacial sediments closer to the SD range, and glacial sediments closer to the PSD/MD range in the Day plot (Dunlop, 2002). The low coercivities also suggest that PSD/MD detrital magnetite dominates over the increase of eolian maghemite during glacials. In the southern Core PS67/219-1, mean magnetic grain sizes are smaller than 0.1 mm even during glacial stages MIS 2e4 (Fig. 7). Such fine magnetic particles also suggest dust may not be the main contributor to k in the core. In contrast to the northern core, the FORC diagram suggests significant contributions of biogenic magnetite in this more sea ice influenced southern core during glacial periods (Fig. 7). This may result from a higher dilution by coarser particles of detrital origin in the northern Scotia Sea. Or it may also indicate a more favorable environment for the production and preservation of biogenic magnetite at the southern site, which rely on the balance between sufficient concentrations of organic carbon and reactive iron to stimulate the growth of magnetotactic bacteria, and organic carbon insufficient to establish reducing diagenetic conditions which may lead to magnetite dissolution (Channell et al., 2013). ARM is principally linked to small magnetic grains including non-interacting SD biogenic magnetite, while IRM is sensitive to a wider range of magnetic grain sizes including interacting SD and PSD/MD detrital magnetite (Verosub and Roberts, 1995; Mazaud et al., 2007; Yamazaki, 2008). The higher correlation coefficients between k and IRM100mT (>0.9) compared to that of k and ARM100mT for both cores (especially for the northern core) also suggest that overall MS of our sediments is mainly controlled by components of detrital origin. Unlike in the sub-Antarctic area with high ocean productivity in glacial times (Anderson et al., 2014; Martinez-Garcia et al., 2014), the strongly extended sea ice cover (Gersonde et al., 2005; Collins et al., 2012b) may prohibit bioproductivity in the glacial Scotia Sea (Sprenk et al., 2013), similar to other Southern Ocean areas south of the APF (Anderson et al., 2009). Although more investigations are needed to quantify its contribution, biogenic magnetite seems not to be the principal contributor to the high glacial MS in Scotia Sea sediments, as suggested by the FORC diagrams (Fig. 7). Besides eolian and biogenic origin, other components contributing to MS may be transported by ocean currents. Glacial high susceptibility is in agreement with the increase of terrigenous

accumulation (Diekmann et al., 2003; Hemming et al., 2007). The northern Scotia Sea sediments point to a Patagonian supply as well as an influx from the southeast Pacific through the Drake Passage; while in the southern Scotia Sea, sediment suspension from the Weddell Sea is considered as the main terrigenous source (Diekmann et al., 2000). During glacials, more terrigenous supply can be expected by nepheloid suspension at low sea level stands. Extended areas exposed to weathering and longer coastlines favored current erosion (Diekmann et al., 2000, 2003). This material is carried by the ACC and is dispersed in the Southern Ocean (Hemming et al., 2007). In areas distant from terrigenous source region, dust may play a more important role for detrital input (Anderson et al., 2014), but this is not the case for the Scotia Sea (Diekmann et al., 2000). The supply of detrital material varies systematically with climate, resulting in the high correlation coefficient for sediment susceptibility and Antarctic dust/climate records but do not originate from ACC intensity variations in the Scotia Sea. Sortable silt records suggest slower glacial than Holocene flow in the southern Scotia Sea due to increased sea-ice coverage, whereas just slightly stronger flow is suggested for the northern part (McCave et al., 2014). Stronger currents in the northern compared to the southern Scotia Sea may also be responsible for overall smaller magnetic grain sizes in the southern core. Comparison of ARM and IRM signals with k suggests different sources for both two cores. ARM and IRM, representing only remanence-carrying minerals, demonstrate almost equal intensities for both cores (Suppl. Fig. 4). Whereas k, which is affected by bulk sediment composition, is higher for Core PS67/197-1 than for Core PS67/219-1. This may indicate that, compared to the southern site, the northern site receives additional material that contains only minor amounts of remanence carrying minerals but mostly paramagnetic (probable clay) minerals. 5.3. Opal content for stratigraphy Biogenic opal content variations demonstrating a clear anticorrelation to k also show great potential to distinguish cold and warm periods south of the APF. Sediments show high interglacial content and low glacial content, in relation to extension/retreat of sea ice and northward/southward displacement of the opal belt (Pudsey and Howe, 1998; Diekmann, 2007; Anderson et al., 2009). During interglacial times, reduced terrigenous input and increased biogenic opal reduce the overall susceptibility to minimum values, and vice versa. Weber et al. (2012) abandoned biogenic silica (BSi) content as an efficient proxy for tuning stratigraphy. As they found anti-phased or shifted relationship at some intervals between the Antarctic ice core record and the wet bulk density derived BSi fluctuations, and they correlated the BSi signal to the North Atlantic climate. However, BSi mainly consisting of diatoms directly reflects conditions at the ocean surface, and is driven by climate variability over the region. By re-evaluation of their age

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model, Weber et al. (2012) erroneously set the “ground-truth” data point for the basic stratigraphic framework (e.g., the tie points for the MIS 1e2 and MIS 3e4 boundaries), which introduced further errors in detailed correlation. This also results in the unusual opal peak in MIS 4 in their record (Sprenk et al., 2013), which was not found in our cores (Fig. 6). Moreover, the wet bulk density calculated BSi content may yield larger errors compared to direct opal measurement. Opal contents of our cores also show variations at intervals without strong oscillations of MS signal. The MIS 1e2 boundary of our sediments are verified by AMS 14C dating, and clearly show high opal content not only at the early Holocene but also at about 14e15 ka, corresponding to Antarctic Isotope Maximum 1 (AIM 1) of the deglacial warming interval (EPICA members, 2006), and mirroring the results of other Southern Ocean records (Anderson et al., 2009). The preservation and deposition of biogenic opal are strongly affected by the sea ice extension, which represents the southern boundary of the circum-polar opal belt (Geibert et al., 2005; Esper and Gersonde, 2014). Although it does not allow detailed correlations of MS to Antarctic ice core ssNaþ signal, which represents the general areal sea ice coverage of the € thlisberger et al., 2010), sea ice extension at a Southern Ocean (Ro specific site may modulate the overall MS by reducing the opal preservation and deposition, which contributes inversely to the MS signal. Our records show that opal content variations have potential providing additional age control at intervals of no strong variation of MS (e.g. glacial-interglacial transition) and paleomagnetic information. 5.4. Paleomagnetic intensity, paleoenvironment implications, and stratigraphic correlations RPI records based on ARM and IRM normalization of both cores show similar but not identical variations for each core (Suppl. Fig. 4; Fig. 8). RPIIRM of both cores can be well correlated based on MS chronology or diatom species variations. The largest differences between RPIARM and RPIIRM in both cores occur at termination 1 to the Holocene interval. This may result from different compositions of the respective magnetic mineral assemblages. The decreasing values of RPIARM in Core PS67/197-1 after the LGM could be caused by a relatively increased supply of small magnetic particles (including in-situ production of biogenic magnetite) to the northern Scotia Sea rather than by a true decrease of the intensity of the Earth's magnetic field. On the contrary, the increasing RPIARM values in Core PS67/219e1 may result from a reduced supply of small magnetic particles. Such variations may suggest a reduced/enhanced Holocene circumpolar current in the northern/southern Scotia Sea along with retreating sea ice fields, respectively, in agreement with sortable silt data (McCave et al., 2014). In summary, RPIIRM seems to be more suitable than RPIARM for regional correlations of geomagnetic records in this area. The MS chronologies suggest that, Core PS67/197-1 extends back to about 86 ka, and Core PS67/219-1 to about 300 ka. Several geomagnetic excursions occurred during this time, such as the Mono Lake (33 ± 1 ka), Laschamp (41 ± 1 ka), NorwegianeGreenland (61 ± 2 ka), Blake (123 ± 3 ka), Iceland Basin (~190 ka), Pringle Falls (~220 ka), CR0/8a (~260 ka), and 9a (~290e310 ka) excursions (Lund et al., 2006a). The Laschamp excursion has been previously identified in the Scotia Sea (Collins et al., 2012a; Weber et al., 2012), in the Northwestern Weddell Sea (Grünig, 1991), and several other Southern hemisphere re-Pronovost et al., 2013). In cords (e.g., Channell et al., 2000; Lise contrast, the Mono Lake excursion was hardly observed in highlatitude Southern Ocean sediments, except for some high-

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sedimentation rate sites (Lund et al., 2006b). The ages based on MS chronology of the clear inclination reversal and the major excursion at 1014 cm and 1299 cm in Core PS67/197-1 are 33.9 ± 0.3 and 41 ± 0.4 ka, respectively, perfectly match and thus were related to the Mono Lake and Laschamp excursions (Fig. 8; Suppl. Fig. 3; Suppl. Fig. 5). This is the first clear identification of the Mono Lake excursion in the Scotia Sea. In addition, further excursions can be recognized by other RPI minima. Grain size corrected RPI records (RPI0 ) can enhance intervals of low relative paleointensity, i.e. geomagnetic excursions compared to the original RPI data. A prominent RPI0 minimum at 1984 cm, with an estimated age of 64.2 ± 0.7 ka, is possibly related to the NorwegianeGreenland Sea excursion, a ‘possible’ geomagnetic excursion (Roberts, 2008). In contrast, no geomagnetic excursions were observed in Core PS67/219-1. The lack of clear excursions in this core may result from its lower sedimentation rates, bioturbation effects and acquisition of a magnetic overprint with normal polarity after deposition (Roberts and Winklhofer, 2004; Roberts et al., 2013) as in records from the western continental rise of the Antarctic Peninsula (Macri et al., 2006). The grain size corrected RPI records show minimum values at 938 cm, 1658 cm, 1898 cm, 1963 cm, 2043 cm, with estimated ages of 41.6 ± 0.6 ka, 185.8 ± 5.7 ka, 225.3 ± 5.6 ka, 256.9 ± 2.3 ka, 291.9 ± 0.8 ka, respectively. At these instances, the demagnetization diagrams suggest low vector intensity and higher demagnetization fields to extract the ChRM (Suppl. Fig. 3). These RPI minimums are likely related to the Laschamp, Iceland Basin, Pringle Falls, CR0/8a, and 9a excursions, respectively, and thus can be used to verify the MS chronology also for intervals of low variability. The Mono Lake, NorwegianeGreenland Sea and Blake geomagnetic excursions are not well recorded in this core. For the corresponding core intervals, the demagnetization diagrams suggest usual behavior as for other intervals, indicating almost identical magnetic properties. This is probably due to the relatively short duration of the excursions (e.g. Mono Lake) and to the low sedimentation rate for the lower part of the core. The relatively weak correlation of the lower core section to RPI reference curves might be related to the significantly lower sedimentation rates, as correlation between individual RPI records are systematically degraded with decreased sedimentation rates and increased age errors (McMillan and Constable, 2006). A minimum sedimentation rate of 10 cm/ka was suggested as prerequisite for reliable RPI recording in order to avoid smoothing effects on the post-depositional remanent magnetization lock-in process (Roberts and Winklhofer, 2004). In the upper part (MIS 1e4) of Core PS67/219-1, sedimentation rates vary between 6 and 37 cm/ka with a mean of 16 cm/ka. For deeper strata, sedimentation rates range from 1 to 12 cm/ka with a mean of only 4 cm/ka. The low sedimentation rate in the lower section may cause an incomplete recording of the variations of the Earth's magnetic field. Based on the MS chronology, the RPI records were compared to reference paleomagnetic curves (Fig. 8), including the South Atlantic paleointensity stack for the past 80 ka (SAPIS) (Stoner et al., 2002); the PASADO record from the South America for the past 51 ka; the composited paleointensity stack SINT-800 spanning the past 800 ka (Guyodo and Valet, 1999); and the highresolution RPI record from ODP Site 1089 as regional reference curve, which spans the past ~580 ka and stems from the Atlantic Southern Ocean (Stoner et al., 2003). In general, the RPI records of both cores show fairly good correlation to the reference curves at short time scales, but the long-term trend differs in amplitude. This is also recognized in other Scotia Sea records and was attributed to the high opal content of the sediments (Weber et al., 2012). This is especially the case in the lower section of Core PS67/

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219-1. Changes in magnetic mineralogy during glacial-interglacial cycles may alter the general features of the RPI records as well. Interglacial MIS 5 and 7 mainly consist of biogenic magnetite, while glacial MIS 6 and 8 have greater amounts of detrital material. ARM/IRM100mT ratios in Core PS67/197-1 vary systematically with k, showing high values during interglacial/interstadial intervals and low values during glacial/stadial intervals (Suppl. Fig. 4). This indicates a relatively higher input of fine magnetic particles (Mazaud et al., 2007), or a higher proportion of biogenic magnetite during warm compared to cold climates. A high correlation coefficient (r ¼ 0.78) is computed between ARM/IRM100mT and both the EDML d18O and the EDC deuterium records. Thus it suggests a consistent pattern of magnetic particle deposition driven by climate at the core site during this time interval. At a longer time scale, ARM/IRM100mT variations in the southern Core PS67/219-1 show similar patterns, but the correlation to Antarctic climate weakens (r ¼ 0.39). This is probably due to lower sedimentation rates and thus incomplete recording of the magnetic signal in the southern core, and subject to a depositional environment different from that of the northern site. 5.5. Age model comparison and validation The individual chronologies based on different parameters show generally good consistency (Fig. 9). Yet discrepancies reflect the weakness of each parameter. The core top Holocene to termination 1 section lacks detailed variability in k and opal content and paleomagnetic signal that can be correlated to reference curves for precise age assignments. The absolute age control of this section relies on AMS 14C dating. Due to the uncertainty caused by fossil carbon contamination, a regional correlation is needed for the validation of each 14C date. The radiocarbon chronology covers a relatively short period of time, thus another stratigraphic approach is needed for older sequences. Fluctuations in diatom species abundances can be useful as they are sensitive to variability in surface ocean conditions that can be correlated across large areas of the ocean. However, the poor development of certain stratigraphic species hampers its application for the old intervals. Although MS shows low values and little variation during interglacials, it represents the best tool for the highly variable glacial chronology, as it shows close resemblance to Antarctic ice core dust/climate records allowing the transfer of ice core age models to marine sediments. This also accounts for the anti-correlation of k to opal content variability in our records. Paleomagnetic chronology may provide support by additional age markers in terms of geomagnetic excursions and to a lesser extent as RPI fluctuations. However, incomplete recording of the paleomagnetic signal due to low sedimentation rates and/or changes in magnetic mineralogy can bias these signals. In addition, the lock-in process has potential shifting the paleomagnetic signal downwards that increases the age uncertainties (Roberts and Winklhofer, 2004; Roterts et al., 2013). Such effect should not be significant in our records, because the proposed geomagnetic excursions show similar ages to those given by literature. All in all, the paleomagnetic signal can be used to verify and refine the MS stratigraphy. In our cores, the 14C ages of deeper strata are significantly younger than those from age models based on MS chronology (Fig. 9). The 14C ages of the deepest dating interval in both cores are 4e5 ka younger than those based on the MS chronology. Such young glacial 14C ages were also found in other Scotia Sea records of similar MS based ages. They are attributed to absorption of modern carbon to the active opal surface of diatom valves during sample preparation for 14C dating (Pugh et al., 2009). However, the glacial

sediments are typically characteristic of low opal content. In our case, this can be caused by the removal of old labile carbon during alkali extraction for preparing humic acid fraction samples for dating. For Core PS67/197-1 there is good agreement between the individual age models based on different parameters throughout the full time coverage, implying that the developed age models are consistent and thus reliable. Therefore these parameters are suitable for the establishment of a stratigraphy in the studied area. In contrast, larger discrepancies between individual age models were found for the lower section of Core PS67/219-1. This may be mainly due to low sedimentation rates that resulted in incomplete recording of (at least) parts of the signals (e.g. for the paleomagnetic signal) and also due to the limitation of available parameters (e.g. diatom species variation). Integrated age models for cores PS67/197-1 and PS67/219-1 were established by correlation of k to Antarctic ice core records, supplemented by AMS 14C humic acid ages for the upper core section where datings from both cores show similar ages for intervals of identical variations of k or diatom abundance fluctuations. Detailed inter-core correlations can be achieved also by opal content and diatom species fluctuations for intervals of low k variability. An additional age control point is transferred using the prominent Holocene ash layer (AS 1) as age marker, dated by AMS 14 C in Core PS67/206-1 (as discussed below). The age uncertainties of the upper core sections are constrained by radiocarbon age calibration; whereas for the lower core sections, they are given by bootstrap re-sampling estimations (Fig. 9). For both cores, ages for geomagnetic excursions based on MS chronology differ less than 1 ka from those given in literature (Lund et al., 2006a; Roberts, 2008) for MIS 3. Larger deviations of 2e5 ka occur in the lower section of Core PS67/219-1. This can be explained by lower sedimentation rates for older intervals (e.g. MIS 8) as compared to younger sections (e.g. MIS 2e3). Additionally, weak oscillations of k during warm intervals (e.g. MIS 5 and 7) contribute to the uncertainties (Fig. 8). However, we abandoned the geomagnetic excursions as age control points for our integrated core stratigraphies. Because the geomagnetic excursion ages were determined at different locations, they might vary by a few thousand years, based on their individual dating strategies and sample resolution (e.g. Lund et al., 2006a; Channell et al., 2009). Nevertheless, the determined ages for the geomagnetic excursions in the studied cores, as well as the LOD of diatom stratigraphic marker species, are well in the range of the uncertainties of the age model based on the correlation of k to Antarctic dust records. Based on the integrated age models presented here (Fig. 9; Suppl. Table 1), time series of different proxies for the past 300 ka, as well as age markers like ash layers, are available and can be used for correlation or dating purposes of other Scotia Sea sediments (Fig. 10). 5.6. Scotia Sea ash layer ages and correlation to Antarctic ice cores The occurrence of volcanic ash is a common phenomenon in cores from the Antarctic Peninsula region (e.g., Moreton, 1999; Hillenbrand et al., 2008), due to the wide distribution of active volcanoes in this area. Ashes are intercalated in marine sediments while also preserved in the Antarctic ice. Based on the integrated age models of cores PS67/197-1 and PS67/219-1, ages of the Scotia Sea ashes are calculated (Table 5). We tried to establish correlations between the marine ash layers and the volcanic signals recorded in Antarctic ice cores. The prominent Holocene ash layer AS1, well preserved in cores PS67/206-1, PS67/219-1 and PS67/205-2, is also found in other Scotia Sea cores (Moreton and Smellie, 1998; Moreton, 1999). This ash layer is AMS 14C dated at 7.8 ka in high-sedimentation rate

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Fig. 9. Age-depth models of cores PS67/197-1 and PS67/219-1 including radiocarbon chronology, correlation of magnetic susceptibility to Antarctic ice core records and diatom stratigraphy. Last occurrence (LO: 1%) of diatom species and postulated geomagnetic excursions are indicated by arrows with estimated ages inferred from MS chronology. The final age models are based on radiocarbon datings on humic acid fraction from the upper core sections and MS chronology, marked by black solid lines connecting each age control point indicated by filled symbols. The age uncertainties for the upper core sections are estimated by radiocarbon calibration (yellow bars), and bootstrap re-sampling estimation for the lower core sections (green and blue dashed line represent 1-sigma standard deviation of the estimation). Sedimentation rates are calculated based on the final age models.

Core PS67/206-1 (Table 3). However, in cores PS67/197-1 and PS67/219-1 it shows more than 1 ka older ages than in Core PS67/ 206-1. These older ages may be attributed to lower sedimentation rates at the core sites and possible bioturbation effect. Thus, the ash age from Core PS67/206-1 is assumed to be more accurate and is therefore transferred to other cores as age marker. AS1 cannot be correlated to published East Antarctic ice core volcanic signals (Severi et al., 2007), suggesting a minor eruption of the source volcano, probably on Deception Island (Moreton, 1999). The ages of AS 2 and AS 3 are 19.4 ± 0.3 and 24.9 ± 0.6 ka, and 19.3 ± 0.6 and 23.6 ± 0.9 ka, for cores PS67/197-1 and PS67/219-1, respectively. These ages are younger than previous datings (21.6 and 26.4 ka, respectively, Moreton, 1999) from other Scotia Sea cores. The AIO ages from Moreton (1999) may be affected by fossil carbon contamination as discussed before. In contrast, our datings show similar ages as those of two ash layers in the Dome Fuji ice core (19 ka and 24 ka) (Kohno et al., 2004). The South Shetland Islands are regarded as their source region. The two ashes in our cores are likely the same ones preserved in the Dome Fuji ice core (Fig. 10), and suggest a reasonable reservoir correction of ~~1300 yrs for the LGM at the core sites. Ash ages for the lower core sections were calibrated by the MS chronology. AS 4 is best preserved in Core PS67/197-1. Our age models suggest ages of 31.1 ± 0.4 ka for Core PS67/197-1, and 31.2 ± 0.7 ka for Core PS67/ 219-1, which is close to the age of an ash layer in the Vostok ice core (32.2 ka). The South Sandwich Islands have been suggested as the source (Basile et al., 2001). AS 5 is widely recorded in Scotia Sea cores. Geochemical analysis also points to Deception Island origin (Moreton, 1999). Our age models suggest an age of 48.1 ± 0.7 ka and 48.8 ± 0.3 ka in cores PS67/197-1 and PS67/2191, respectively. No visible ash layers at this time interval have yet been reported in the reference Antarctic ice cores (Fig. 10; Basile et al., 2001; Kohno et al., 2004; Narcisi et al., 2005). The age by Moreton (1999) is too young (35.4 ka) based on 14C dating, possibly because this interval is close to the limit of 14C dating. Low 14C content in the sediment is easily biased towards young ages by removing old labile carbon during alkali extraction. Ash AS 6 occurs at a depth of 2050 cm in Core PS67/197-1 and

1175 cm in Core PS67/219-1, giving ages of 68.5 ± 0.5 ka and 69.6 ± 3.2 ka, respectively. The ages are close to those of an ash in Dome Fuji at 68 ka or 70 ka (Kohno et al., 2004) (Fig. 10). AS 7 has an age of 180.4 ± 4.5 ka in Core PS67/219-1, similar to an ash layer in Vostok ice core (179.3 ka). The South Sandwich Islands were considered as origin (Basile et al., 2001). Another ash layer occurs in Dome Fuji ice core at 177 ka (Kohno et al., 2004). AS 8 yields an age of 223.7 ± 5.4 ka in Core PS67/219-1, similar to the Dome Fuji ash layer at 226 ka (Kohno et al., 2004). AS 9 occurs at 274.4 ± 1.5 ka in Core PS67/219-1, but no ice core ash layer information is available yet in the relevant time interval (Narcisi et al., 2005, 2010). Concrete correlation of ash layers between different cores and archives relies on geochemical evidences. In this study, the geochemical composition of the tephra is not available, thus the correlation remains tentative. However, tephra in marine sediments are still difficult to correlate geochemically (Hillenbrand et al., 2008) because of at least 3 complications: 1. Chemical composition of tephra changes through time even from the same source; 2. Similarities of tephra particles from different volcanic sources; 3. Alteration by submarine weathering of sediments. Geochemical identification of tephra is also difficult between the tephra preserved in marine sediments and in ice cores, because the tephra in ice cores are sorted by wind and only fine grained particles can reach the ice core sites. Only very strong geochemical fingerprints may show convincing evidence of correlation. Nevertheless, our study suggests possible correlations between the Scotia Sea sediments and Antarctic ice cores, and provides additional age markers for the Scotia Sea records.

6. Conclusions A detailed Scotia Sea sediment stratigraphy for the past 300 ka is established based on radiometric dating, correlation of magnetic susceptibility to Antarctic ice core signals, biogenic opal content variability, diatom stratigraphy and is further evaluated by paleomagnetic information. This multi-parameter approach improves the fidelity of the presented age models.

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Fig. 10. Stratigraphic parameters of cores PS67/197-1 and PS67/219-1 compared with dust records from EDML (Fischer et al., 2007) and EDC (Lambert et al., 2008), respectively, and ODP 1089 RPI record (Stoner et al., 2003). Antarctic ice core ash (Basile et al., 2001; Kohno et al., 2004; Narcisi et al., 2005) are labeled as in Fig. 5, possible correlations to Scotia Sea ash (AS) are indicated.

The generally good consistency between individual datings based on different parameters indicates the validity of these parameters to be applied as dating tools for Scotia Sea sediments. AMS 14 C chronology is useful for the upper core section but is complicated by the uncertainty of fossil carbon contamination. Regional comparison and correlations are required to determine reliable ages. Magnetic susceptibility represents the best long-term stratigraphic approach for Scotia Sea sediments. Low-coercive minerals dominate the magnetic inventory in both glacial and interglacial

sediments. MS is distinctly affected by fine-grained (SD) particles, probably biogenic magnetite, in particular in interglacial sediments. The southern core shows a significant content of finegrained (SD), probably biogenic, magnetite also during glacials, suggesting a depositional environment different from that at the northern site. Paleomagnetic information supports the MS chronology, and suggests the first identification of the Mono Lake geomagnetic excursion in Scotia Sea sediments. The abundance fluctuations of F. kerguelensis and E. antarctica are useful stratigraphic tools for the past six marine isotope stages

W. Xiao et al. / Quaternary Geochronology 31 (2016) 97e118 Table 5 Ages of Scotia Sea ash layers in each core. AMS 14C derived ages are marked by stars, and the other ages are based on MS chronology. Age uncertainties are based on 14C calibration and bootstrap re-sampling estimation. Scotia ash

AS AS AS AS AS AS AS AS AS

1 2 3 4 5 6 7 8 9

Age (ka) PS67/197-1

PS67/219-1

PS67/206-1

10.3 ± 0.2* 19.4 ± 0.3* 24.9 ± 0.6 31 ± 0.4 48.1 ± 0.7 68.5 ± 0.5

8.8 ± 0.2* 19.3 ± 0.6* 23.6 ± 0.9* 31.2 ± 0.7 48.8 ± 0.3 69.6 ± 3.2 180.4 ± 4.5 223.7 ± 5.4 274.4 ± 1.5

7.8 ± 0.09*

(MIS) in the Scotia Sea, but the fluctuations weaken during MIS 7 and 8. Diatom concentration and biogenic opal content can be used as alternative parameters to distinguish cold and warm intervals and for regional core correlation. Based on the presented age models, possible correlations between several ash layers found in Scotia Sea sediments and those found in East Antarctic ice cores are suggested and provide additional age markers for further studies in this area. Acknowledgments This work was co-supported by the Alfred Wegener Institute PACES program (Polar Regions and Coasts in the Changing Earth System), Deutsche Forschungsgemeinschaft DFG Research Center FTZ-15 MARUM-OC3 Project, DFG project KU 683/9, and the HOLOCLIP Project, a joint research project of ESF PolarCLIMATE program, funded by national contributions from Italy, France, Germany, Spain, Netherlands, Belgium and the United Kingdom. We thank the members of R/V Polarstern cruises ANT-X/5, ANT-XI/ 2, and ANT-XXII/4, for retrieving the study materials. The constructive comments from an anonymous reviewer and the editor were very helpful in improving the manuscript. Appendix A. Supplementary data Supplementary data related to this article can be found at http:// dx.doi.org/10.1016/j.quageo.2015.11.003. References 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. http://dx.doi.org/10.1029/2005PA001199. PA1013. Anderson, J.B., Andrews, J.T., 1999. Radiocarbon constraints on icesheet advance and retreat in the Weddell Sea, Antarctica. Geology 27, 179e182. 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, 1143e1148. Anderson, R.F., Barker, S., Fleisher, M., Gersonde, R., Goldstein, S.L., Kuhn, G., Mortyn, P.G., Pahnke, K., Sachs, J.P., 2014. Biological response to millennial variability of dust and nutrient supply in the Subantarctic South Atlantic Ocean. Phil. Trans. R. Soc. A 372, 20130054. Andrews, J.T., Domack, E.W., Cunningham, W.L., Leventer, A., Licht, K.J., Jull, A.J.T., DeMaster, D.J., Jennings, A.E., 1999. Problems and possible solutions concerning radiocarbon dating of surface marine sediments, Ross Sea, Antarctica. Quat. Res. 52, 206e216. Appleby, P., 2001. Chronostratigraphic techniques in recent sediments. In: Last, W.M., Smol, J.P. (Eds.), Tracking Environmental Change Using Lake

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