Middle to Late Pleistocene Arctic paleoceanographic changes based on sedimentary records from Mendeleev Ridge and Makarov Basin

Quaternary Science Reviews 228 (2020) 106105

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Middle to Late Pleistocene Arctic paleoceanographic changes based on sedimentary records from Mendeleev Ridge and Makarov Basin € wemark d, Jing Mei a, Wenshen Xiao a, b, *, Leonid Polyak c, Rujian Wang a, b, Ludvig Lo Defang You a, Weiguo Wang e, Li Wu a, Xiaobo Jin a a

State Key Laboratory of Marine Geology, Tongji University, Shanghai, 200092, China University Corporation for Polar Research, Beijing, 100875, China Byrd Polar and Climate Research Center, The Ohio State University, Columbus, OH, 43210, USA d Department of Geosciences, National Taiwan University, P.O. Box 13-318, 106, Taipei, Taiwan e Third Institute of Oceanography, SOA, Xiamen, 361005, China b c

a r t i c l e i n f o

a b s t r a c t

Article history: Received 22 July 2019 Received in revised form 20 November 2019 Accepted 22 November 2019 Available online xxx

Three sediment cores from the Makarov Basin and Mendeleev Ridge off the Siberian margin were analyzed for a number of litho-, bio- and magnetostratigraphic proxies. Correlation to previously studied Arctic records from Mendeleev to Lomonosov ridges enables paleoceanographic reconstructions for the Middle to Late Pleistocene (estimated ~0.8 Ma) on orbital time scales. Cyclic changes in lithology corroborated by multiple stratigraphic constraints were tuned to global climate variations based on the manganese content that increases during interglacials, while calcium and zircon were used as proxies for the Laurentide and Eurasian ice sheet inputs. The combined evaluation of glacial and interglacial proxy variabilities provides a coherent picture of the western Arctic Ocean development in the Pleistocene. We suggest that major cyclic changes were driven by summer processes, such as sea-ice retreat and enhanced biological production vs. suppressed ice melt in warmer and colder summers, respectively. A potential link to low latitudes via atmospheric moisture and heat transport is indicated by similarities in the western Arctic Ocean and East Asian monsoon records. Long-term cooling is signified by biogenic proxies related to sea-ice expansion over the western Arctic Ocean with step changes at MIS 11 and 7 (~0.4 and 0.2e0.3 Ma). This transition also features better fossil preservation indicating less corrosive/ oxidative bottom waters, probably resulting from a decrease in organic export production and/or weakening of bottom water ventilation under perennial sea ice. The latter was possibly enhanced by atmospheric cooling and increased meltwater fluxes with the growth of circum-Arctic ice sheets. © 2019 Elsevier Ltd. All rights reserved.

Keywords: Arctic Ocean Mendeleev Ridge Quaternary stratigraphy Sedimentary proxies Glaciations Interglacials Paleocirculation Orbital cyclicity

1. Introduction Climate warming in recent decades has caused significant sea ice reduction in the Arctic Ocean, both in extent and thickness (e.g., Cavalieri and Parkinson, 2012; Lang et al., 2017). This reduction, along with an increase in Arctic surface air temperature that exceeded northern hemisphere average rates (Arctic amplification: Serreze et al., 2009; Screen and Simmonds, 2010), resulted in pronounced changes in the Arctic atmospheric circulation (Wernli and Papritz, 2018), hydrographic regime (Polyakov et al., 2010),

* Corresponding author. State Key Laboratory of Marine Geology, Tongji University, Shanghai, 200092, China. E-mail address: [email protected] (W. Xiao). https://doi.org/10.1016/j.quascirev.2019.106105 0277-3791/© 2019 Elsevier Ltd. All rights reserved.

and biota (Fountain et al., 2012). There is mounting evidence that the effect of these changes extends far beyond the Arctic, including weakening of North Atlantic deep-water convection (Jahn and Holland, 2013) and mid-latitude weather conditions (Mori et al., 2014; Wu et al., 2016; Screen, 2017). Understanding these profound changes in a broader climatic context requires investigation of the dynamics of the Arctic Ocean system in a long time perspective, from submillennial to glacialinterglacial scales. While past warming periods, such as interglacials, provide the closest paleoclimatic analogs for the modern change, it is the history of glaciations that constitutes a baseline for Quaternary climate evolution. Growth and decay of Arctic ice sheets largely modulated global sea level variations, and affected the global freshwater balance, thermohaline circulation, and carbon cycle (Imbrie and Imbrie, 1980; Miller et al., 2010). Impact on

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the Arctic Ocean was especially pronounced as its geographic and environmental setting during glacial periods was remarkably different from present. Low glacial sea levels exposed vast Arctic continental shelves, which drastically decreased the Arctic Ocean area, and enabled build-up of large marine ice sheets that extended over the deep basins as ice shelves (Polyak et al., 2001; Niessen et al., 2013; Jakobsson et al., 2014, 2016). These glacial build-up and decay events largely controlled longterm Pleistocene sedimentation in the Arctic Ocean. The sediments, therefore, provide valuable paleoclimatic/paleoceanographic records (e.g., Polyak and Jakobsson, 2011). While depositional disruptions are inevitable due to erosion on top of submarine ridges, slope processes, and turbidites in the basins, Arctic Ocean sediment cores provide a more continuous history of circum-Arctic glaciations than sporadic terrestrial records from glaciated areas (e.g., Knies et al., 2001; Spielhagen et al., 2004; Adler et al., 2009; Polyak et al., 2009). Additionally, paleoceanographic data enable coupling the glaciation history with changes in marine environments, such as circulation patterns (e.g., Bischof and Darby, 1997; Phillips and Grantz, 2001; St€ arz et al., 2012). Despite the obvious advantages of studying Arctic Ocean sedimentary records, their investigation is hindered by mostly low sedimentation rates and difficulties in developing reliable age constraints (e.g., Jakobsson et al., 2000; Polyak et al., 2009; Stein et al., 2010). Prior studies have identified a major step change in the western Arctic Ocean depositional history estimated to have occurred at the end of the Mid-Pleistocene Transition (MPT), close to the Early/Middle Pleistocene boundary around ~0.8 Ma (Polyak et al., 2009, 2013; Dipre et al., 2018). This change is expressed in the pronounced rise in IRD deposition, including detrital carbonates indicative of the Laurentide Ice Sheet (LIS) discharge, and onset of perennial sea-ice cover inferred from paleobiological proxies. In addition to this transition, considerable changes have been suggested for the Middle to Late Pleistocene Arctic paleoceanographic record (Cronin et al., 2013, 2017; Polyak et al., 2013; Lazar and Polyak, 2016; Dong et al., 2017). In order to gain further insight into depositional environments in the western Arctic Ocean, their implications for oceanographic changes and glaciation history of the adjacent North American and Eurasian margins, and evaluate the climatic linkages between the Arctic and low latitude processes, we study three sediment cores collected on the CHINARE V and VII expeditions from Makarov Basin and Mendeleev Ridge off the East Siberian margin. Based on a multi-proxy approach, we develop an overall higher-resolution sediment stratigraphy than reported earlier for the study region, enabling reconstruction of Middle to Late Pleistocene paleoenvironments on orbital time scales. Our data corroborate major conclusions from previous studies, and provide new knowledge on Arctic glacial-interglacial changes and their potential controls.

2. Background 2.1. Oceanographic settings The Arctic Ocean is a semi-enclosed basin that connects to the North Pacific Ocean via the shallow (~50 m) Bering Strait, and to the North Atlantic Ocean via the deep Fram Strait (~2600 m) and the Barents Sea shelf (Fig. 1). Shallow circum-Arctic continental shelves comprise about half of the Arctic Ocean's area, making circulation in the Arctic Ocean sensitive to sea level changes (Jakobsson, 2002, Jakobsson et al., 2008). Geologically and morphologically the Arctic Ocean is generally subdivided into Eurasian and Amerasian basins separated by Lomonosov Ridge that extends from the Siberian margin to Greenland and Canada. However, in contrast to its Eurasian counterpart, Amerasian Basin, broadly referred to as the

western Arctic Ocean, comprises a diverse system of basins and ridges. Large scale surface ocean circulation in the Arctic Ocean is characterized by the Transpolar Drift (TPD), that extends from the Siberian shelf towards Fram Strait, and the clockwise Beaufort Gyre (BG) in the western Arctic Ocean driven by anticyclonic atmospheric circulation (Aagaard, 1989; Rigor, 1992) (Fig. 1). The BG stores large volumes of freshwater in the surface layer, resulting in a deepened halocline in the western Arctic Ocean (Proshutinsky et al., 2002, 2009). The boundary between the TPD and BG, controlled by fluctuating atmospheric circulation patterns such as the Arctic Oscillation, varies on inter-annual to decadal time scales (Dickson et al., 2000; Rigor et al., 2002). This variability has also been inferred for longer time periods, such as throughout the Holocene (Darby et al., 2012). On geological time scales, circulation is also influenced by changing bathymetry and freshwater forcing (Prange and Lohmann, 2003). Circum-Arctic rivers drain large areas of Siberia, Northern Europe and North America, with a total annual runoff of about 3200 km3 freshwater equivalent to 10% of the global runoff (Serreze et al., 2006). This runoff delivers ~250 million tons of sediments to the Arctic Ocean (Holmes et al., 2002). Siberian rivers dominate the total runoff (Peterson et al., 2002), while the Mackenzie River contributes more than half of the sediments (Holmes et al., 2002; Macdonald and Gobeil, 2012). Riverine freshwater and sediment discharge show a strong seasonal variability, with highest fluxes in late spring and early summer (Holmes et al., 2002).

2.2. Sediment stratigraphy The year-round ice-covered Arctic Ocean is characterized by low sedimentation rates, low biological productivity, and strong diagenesis of bio- and chemogenic sediment components, which complicates establishment of Arctic sediment chronostratigraphy (e.g., Polyak et al., 2009; Stein et al., 2010; Alexanderson et al., 2014). Discontinuous occurrence of carbonate microfossils in Arctic Ocean sediments hinders the application of oxygen isotope stratigraphy, which is widely used in other oceans. In addition, complex Arctic hydrological process (e.g. freshwater discharge, sea ice melt, brine formation) (Bauch et al., 2011) and possible habitat changes of planktic foraminifera (Bauch et al., 1997; Xiao et al., 2014) further complicate the use of oxygen isotopes as a stratigraphic tool. The application of paleomagnetic data, another stratigraphic tool commonly used for ocean sediments, is also problematic in the Arctic due to unresolved, possibly diagenetic nature of frequent inclination reversals. This issue complicates comparison of Arctic sediment stratigraphy to global geomagnetic polarity changes (Jakobsson et al., 2000; Channell and Xuan, 2009; Xuan and Channell, 2010). While true chronostratigraphic constraints are yet to be developed, recurrent sedimentary beds with distinct lithologies signified by contrasting brownish and grayish/yellowish colors are inferred to be representative of interglacial/major interstadial and glacial/ stadial intervals, respectively. This lithostratigraphic cyclicity provides a basis for climato/cyclostratigraphy conventionally used for Quaternary Arctic Ocean records over the last two decades (e.g., Jakobsson et al., 2000; Polyak et al., 2004; Stein et al., 2010). This approach, enhanced by tuning to global climatic records, has been tested on variations in sedimentary proxies such as bulk density (O'Regan et al., 2008), magnetic grain size (Xuan et al., 2012), or, most commonly, manganese (Mn) content (Jakobsson et al., 2000; €wemark et al., 2008; Wang et al., 2018). Lo Mn measurements, enabled nowadays by continuous X-ray fluorescence (XRF) logging for elemental composition, show

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Fig. 1. Arctic Ocean physiographic and oceanographic settings (Aagaard, 1989; Rigor, 1992; Jakobsson et al., 2008) with locations of reported (red) and referenced (dark gray) sediment cores (Table 1). White shaded areas indicate the maximum extent of Pleistocene glaciations around the Arctic Ocean (Niessen et al., 2013; Jakobsson et al., 2016; Stokes et al., 2016). Dark and light blue dashed lines denote the average summer (September) sea ice extent of 1979e2006 (Parkinson and Cavalieri, 2008) and the 2012 minimum (www. nsidc.org), respectively. Thin black contour on the Siberian and Alaskan shelf denotes 50 m isobath. Large circum-Arctic rivers are indicated by yellow arrows. TPD: Transpolar Drift; BG: Beaufort Gyre. LR: Lomonosov Ridge; MR: Mendeleev Ridge; AR: Alpha Ridge; CB: Canada Basin; MB: Makarov Basin; EB: Eurasian Basin; CP: Chukchi Plateau; NR: Northwind Ridge; CAA: Canadian Arctic Archipelago; BS: Bering Strait; GIS: Greenland Ice Sheet; LIS: Laurentide Ice Sheet; ESIS: East Siberian Ice Sheet; EAIS: Eurasian Ice Sheet. (For interpretation of the references to color in this figure legend, the reader is referred to the Web version of this article.)

explicit variations with high and low values attributed to interglacial/interstadial and glacial/stadial periods, respectively € wemark et al., 2014). (Jakobsson et al., 2000; Polyak et al., 2004; Lo High Mn levels are believed to result from enhanced basin-ward transport of Mn oxy(hydr)oxides from the shelves and their deposition in the central Arctic Ocean during time intervals of high sea level and relatively warm climatic conditions (Macdonald and €wemark et al., 2014; Ye et al., 2019). Regardless of Gobeil, 2012; Lo the paleoclimate interpretation, Mn variability provides a useful stratigraphic correlation tool that shows consistent results across the western Arctic Ocean (Alexanderson et al., 2014; Wang et al., 2018; Schreck et al., 2018). A number of lithologic and biogenic stratigraphic features are used to verify the consistency of cyclic proxy variations between spatially distributed cores. In particular, layers enriched in coarsegrained detrital carbonates, the most distinct of which are known as the pink-white layers since the early studies (Clark et al., 1980), have been attributed to iceberg discharge events from the Laurentide Ice Sheet (LIS). These beds are widely used as Late to Middle Pleistocene stratigraphic markers in the western Arctic Ocean (Polyak et al., 2009; Stein et al., 2010; Bazhenova et al., 2017). Occurrences of several microfossil taxa, mostly foraminiferal and

ostracode species, also provide independent stratigraphic constraints. Their distribution, however, is primarily controlled by environmental factors such as sea ice cover, water depth, and preservation conditions, and is not well calibrated to absolute ages (Poore et al., 1994; Polyak et al., 2004; Cronin et al., 2014; Lazar and Polyak, 2016). A notable exception is a coccolith Emiliania huxleyi that evolved during Marine Isotope Stage (MIS) 8 (Thierstein et al., 1977). This species therefore represents a true biostratigraphic marker for Late to Middle Pleistocene sediments. Consistent with its global stratigraphic range, in Arctic sediment cores E. huxleyi is only found in intervals estimated as MIS 5 and younger, with less certainty in MIS 7 (Gard, 1993; Backman et al., 2004, 2009; Spielhagen et al., 2004). The age model based on stratigraphic approaches mentioned above can be additionally constrained by its relation to the late Quaternary glacial history at the Arctic periphery (Spielhagen et al., 2004; Schreck et al., 2018). On a longer time scale, this age framework is also consistent with the strontium isotope stratigraphy developed recently for the Quaternary to Pliocene record from the western Arctic (Dipre et al., 2018). The resultant cm/kyrscale average sedimentation rates estimated for Quaternary records from the central Arctic agree well with longer-term borehole

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stratigraphy and geophysical data (Backman et al., 2004; O'Regan et al., 2008). In contrast, some other approaches to Arctic Ocean chronostratigraphy, such as attempts to correlate apparent geomagnetic inclination swings to global polarity changes resulted in much lower sedimentation rates on the order of mm/kyr or lower (e.g., Herman, 1970; Clark et al., 1980). Comparably low sedimentation rates were estimated from uranium-series isotope profiles (Ku and Broecker, 1965; Huh et al., 1997; Not and HillaireMarcel, 2010; Hillaire-Marcel et al., 2017). These disparities prompt further testing of varous approaches to Arctic Ocean sediment stratigraphy and comparison of resulting age assessments. To summarize, despite remaining uncertainties, variations in sedimentary parameters suggestive of glacial-interglacial climate controls on Arctic environments make sense in paleoceanographic context. That said, multi-proxy approaches and detailed regional correlations aided by independent stratigraphic markers, where possible, are needed to verify the consistency and fidelity of Arctic sediment stratigraphies. 3. Materials and methods 3.1. Sediment material Gravity cores ARC5-ICE6, ARC5-MA01 and ARC7-E26 (hereafter ICE6, MA01 and E26) were retrieved from Makarov Basin and central to southern segments of Mendeleev Ridge during the 5th and 7th Chinese National Arctic Research Expeditions (CHINARE V and VII) in 2012 and 2016 (Ma, 2013; Li, 2018) (Fig. 1, Table 1). Earlier reported HOTRAX 2005 cores HYL0503-6JPC and 8JPC (Adler et al., 2009; Polyak et al., 2009; Channell and Xuan, 2009; Xuan and Channell, 2010; Cronin et al., 2013) including some additional, previously unpublished data are used for a more comprehensive characterization of Mendeleev Ridge stratigraphy. As sediment cores under study are located at the interaction zone of the BG and TPD, they are expected to record Pleistocene glacial-interglacial changes in circulation and terrigenous inputs from Eurasian and North American ice sheets. Core ICE6 from the Makarov Basin enables correlation of these records with earlier reported cores from central Lomonosov Ridge that characterizes a more persistent TPD influence (cores AO96/12-1 PC and PS2185-6: Jakobsson et al., €wemark et al., 2014). 2000, 2001; Spielhagen et al., 1997, 2004; Lo Core ARC5-BN05 from Canada Basin (Dong et al., 2017) is additionally used for broader regional correlations and paleoceanographic context. 3.2. Laboratory methods Color reflectance measurements were performed at 1-cm

intervals on fresh split-core surfaces using a Minolta CM2002 spectrophotometer. Reflectance data are expressed as CIELAB L*, a* and b* color space indicating lightness, red/green and yellow/blue color indices, respectively (Fig. 2 for L* and a*). Bulk elemental composition was measured at 1-cm intervals on split-core surfaces covered with SPEX CertiPrep 3525 Ultralene foil using the Avaatech non-destructive XRF core scanner at Tongji University. Relative concentrations of the element series from Al to Ba were determined by triple-scanning at 10 kv, 30 kv and 50 kv tube voltages for 30 s for each energy level. Results are given as counts/30s. XRF data for HOTRAX cores 6JPC and 8JPC used for comparison were measured at Byrd Polar and Climate Research Center using an energy-dispersive Innov-X Alpha series handheld analyzer at increments between 2 and 6 cm (Darby et al., 2005; Polyak et al., 2009). XRF data for Mn, Ca and Zr (Fig. 2) were investigated in detail as potential proxies for interglacial (Mn) and glacial (Ca, Zr) conditions (e.g., Jakobsson et al., 2000; Polyak et al., 2009; Wang et al., 2018). In order to reduce the potential influence of matrix effects (e.g. grain size, porosity, water content, organic € ning et al., matter content) on the variability of element content (Bo € wemark et al., 2011; Croudace and Rothwell, 2017), XRF 2007; Lo data for Mn, Ca and Zr are normalized to Ti. We prefer Ti to Al that is commonly used for normalization because Al is enriched in light minerals and shows stronger variation in the Arctic records, likely resulting in a bias by source change. In particular, Zr is more enriched in heavy minerals than Al, similar to Ti, therefor Zr/Ti is less likely to be affected by hydrodynamic processes than Zr/Al €rz et al., 2011). In comparison, Mn and Ca counts in studied (Ma cores show nearly identical variations to those of Mn/Ti and Ca/Ti (same as Mn/Al and Ca/Al), consistent with XRF measurements in €wemark et al., 2014; Dong other Arctic sediment studies (e.g., Lo et al., 2017). This similarity in variation implies that, the inputs of Mn and Ca are largely independent from background lithogenic elements. Sediment subsamples taken at 2-cm intervals were oven-dried at 45
C. About 10 g of dry sediment from each sample was wet sieved through a 63-mm mesh. The >63 mm fraction was then dried and sieved through 150-mm mesh. As the biogenic component (mostly foraminifera) is only minor in Arctic Ocean sediments (Polyak et al., 2004; Cronin et al., 2008), coarse fractions are used to approximate Ice Rafted Debris (IRD) (Fig. 2). An aliquot of the >150 mm residue material was examined under microscope, planktic and benthic foraminifera tests were counted in the aliquot for calculating their total abundance expressed per gram of dry sediment (Fig. 2). The presence of species of stratigraphic significance, such as benthic foraminifera Oridorsalis tener, Bolivina arctica, Bulimina aculeata, and Pullenia bulloides, planktic foraminifera Turborotalita eglida, and ostracod Acetabulastoma

Table 1 Information of studied and referenced cores. Site

Abbr.

Latitude (
N)

Longitude (
E)

Water depth (m)

Area

Reference

ARC5-ICE6 ARC5-MA01 ARC7-E26 HLY0503-6JPC

ICE6 MA01 E26 6JPC

83.628 82.031 79.950 78.294

161.764 178.960 179.697 176.986

2901 2295 1500 800

Makarov Basin Mendeleev Ridge Mendeleev Ridge Mendeleev Ridge

HLY0503-8JPC

8JPC

79.593

172.502

2792

Mendeleev Ridge

AF00-08 AO96/12-1 PC

AF8 96/12

82.087 87.098

179.867 144.773

1530 1003

Mendeleev Ridge Lomonosov Ridge

PS2185-6 ACEX ARC4-BN05

PS2185 ACEX BN05

87.537 87.921 80.484

144.927 139.365 161.465

1052 1209 3156

Lomonosov Ridge Lomonosov Ridge Canada Basin

This study This study This study Darby et al. (2005); Cronin et al. (2014); Lazar and Polyak (2016) Darby et al. (2005); Adler et al. (2009); Polyak et al. (2009) Krylov et al. (2011) Jakobsson et al. (2000), 2001; €wemark et al. (2014) Lo Spielhagen et al. (1997), 2004 O'Regan et al. (2008) Dong et al. (2017)

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Fig. 2. Core photos, color reflectance, XRF element content, coarse sediment fraction (>150 mm), foraminifera abundance, paleomagnetic inclination, and biostratigraphic markers in cores under study. 14C ages (Table 2) are shown next to foraminiferal curves. PW1, PW1.1, PW2 and W3 e detrital carbonate marker layers. Occurrences of foraminiferal/ostracod index species are shown by vertical bars; coccolith Emiliania huxleyi in core MA01 is indicated by star. Yellow shading marks the stratigraphic intervals above the major inclination drop in each core. (For interpretation of the references to color in this figure legend, the reader is referred to the Web version of this article.)

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Table 2 Radiocarbon datings on planktonic foraminifera Neogloboquadrina pachyderma from cores MA01, ICE6 and E26. The calibration program CALIB 7.04 (Stuiver and Reimer, 1993; Reimer et al., 2013) was used to convert the 14C ages to calendar ages with 2 sigma precision. 700 yr and 1400 yr (Hanslik et al., 2010) reservoir corrections are taken for Holocene and deglacial-Last Glacial intervals, respectively. The calibrated ages marked bold were applied for stratigraphy. Core

Depth (cm)

Lab ID

14

1s (±)

Reservoir (yr)

Calib age (yr BP)

Calib age Median probability (yr BP)

ICE6 ICE6 ICE6 ICE6 ICE6

0e2 8e10 18e20 24e26 34e36

UCIAMS183148 UCIAMS183149 UCIAMS183150 UCIAMS183151 UCIAMS183152

5725 11920 41890 42540 43080

15 25 780 850 910

700 1400 1400 1400 1400

5732e5876 12346-12598 42815-45358 43117-46007 43367-46675

5805 12498 44058 44597 45063

MA01 MA01 MA01 MA01 MA01 MA01 MA01 MA01 MA01 MA01 MA01 MA01

0e2 2e4 4e6 6e8 8e10 10e12 12e14 14e16 16e18 18e20 20e22 22e24

UCIT29764 UCIT29765 UCIT29766 UCIT29767 UCIT29768 UCIT29769 UCIT29770 UCIT29771 UCIT29772 UCIT29773 UCIT29774 UCIT29775

5170 9905 12795 28170 33380 37610 39470 40580 40960 41270 38360 42470

20 25 30 160 270 450 630 730 670 700 550 810

700 700 1400 1400 1400 1400 1400 1400 1400 1400 1400 1400

5066e5276 10394-10575 13174-13367 30697-31157 35235-36407 39874-41677 41357-43121 42026-44277 42347-44467 42515-44756 40408-42299 43128-45880

5196 10494 13273 30936 35867 40820 42260 43040 43324 43572 41465 44538

E26 E26 E26 E26

0e2 6e8 22e24 28e30

UCIT36722 UCIT36723 UCIT36724 UCIT36725

5280 11795 38320 45170

20 30 580 1320

700 1400 1400 1400

5275e5423 12074-12462 40310-42307 44859-49736

5325 12274 41423 47101

C age (yr BP)

arcticum and Polycope spp. (e.g., Polyak et al., 2004; Hanslik et al., 2013; Cronin et al., 2013, 2014; Lazar and Polyak, 2016), were registered in both >63 mm and >150 mm fractions. In order to semi-quantitatively examine the occurrence of calcareous nannofossil stratigraphic marker Emiliania huxleyi, a total of 30 smear slides of raw sediment material from high-Mn intervals in the upper sections of all three cores (0e240 cm in ICE6; 0e180 cm in MA01; 0e220 cm in E26) were prepared and examined under a polarized light microscope Leica DM6000B at a magnification of 1250. Nanofossil occurrence was examined in 62 view fields with an area of ~1.24 mm2. Tests of planktic foraminifera Neogloboquadrina pachyderma (Np) were picked from the >150 mm fraction in upper core sediments for Accelerator Mass Spectrometry radiocarbon (AMS 14C) dating at the University of California, Irvine (UCIAMS and UCIT), USA (Fig. 2, Table 2). Conversion to calendar ages was performed using the Calib 7.0.4 program (Stuiver and Reimer, 1993) and the Marine13 calibration dataset (Reimer et al., 2013) with reservoir corrections of 700 years and 1400 years applied to Holocene and older sediments, respectively (Hanslik et al., 2010). Uncertainties with reservoir changes in the Arctic Ocean, especially during glacial intervals, may result in a several hundred years deviation in calibrated calendar ages, yet this uncertainty does not affect age assignments at kyr time scales. Paleomagnetic directions (inclination) and intensities of natural remanent magnetization (NRM) were measured on u-channel samples at 1-cm spacing for core ICE6 and E26, and at 2-cm spacing for MA01 using a 2-G Enterprises 755/760-4k cryogenic magnetometer at the Third Institute of Oceanography, SOA. To characterize post-depositional magnetization, NRM was measured before and after stepwise demagnetization in an alternating field (AF) of 5 mTe80 mT. As polarity directions of Arctic sediments may be subject to self-reversals due to oxidation of magnetic minerals (Channell and Xuan, 2009; Xuan and Channell, 2010), for correlation purposes we used inclination data after a 10 mT demagnetization step that removes weak viscous magnetizations (Fig. 2), similar to the approach of O'Regan et al. (2008).

3.3. Stratigraphic assignments and time-series analysis Radiocarbon dates provide age constraints for the youngest strata (<50 ka) with uncertainties related to insufficiently understood reservoir mixing as indicated above. For older sediments, stratigraphic control is based on cyclic occurrence of climaticallymediated sedimentary proxies, such as Mn content, verified by regional multiproxy correlations to earlier developed Arctic Ocean stratigraphies. These correlations include depositional events, microfossil stratigraphic markers, and paleomagnetic variations as discussed below in section 5.1. Resulting age models for each core were refined by cyclostratigraphic tuning of Mn variations to global paleoclimate and sea level records (Lisiecki and Raymo, 2005; Miller et al., 2011) using AnalySeries software (Paillard et al., 1996) (Fig. 3). Transitions between high- and low-Mn intervals were used as tie points for glacial/interglacial boundaries with an arbitrary 10 kyr error (Suppl. Table 1), as in Pleistocene ACEX stratigraphy estimates (O'Regan et al., 2008). In order to examine trends and periodicities in the time frequency domain, the Matlab® package developed by Grinsted et al. (2004) was adopted to perform continuous wavelet transform on time series of sedimentary proxy variations in core MA01. The default mother wavelet “Morlet” (with u0 ¼ 6) and a significance level of 5% were applied for these analyses. For this evaluation we primarily use Mn as a proxy of interglacial/major interstadial con€ wemark ditions (Jakobsson et al., 2000; O'Regan et al., 2008; Lo et al., 2014). Zr was used as a proxy of glacial depositional environments as it has a more representative downcore distribution than Ca, consistent with the more proximal Siberian provenance inferred for Zr.

4. Results 4.1. Radiocarbon dating Radiocarbon data constrain ages for top-most sediments (0e2 cm) in cores ICE6, MA01 and E26 to ~5e6 ka (Table 2). Similar

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Fig. 3. Stratigraphic correlation of sediment cores from the Lomonosov Ridge to Mendeleev Ridge and Canada Basin by Mn and Ca content (expressed as Mn/Ti and Ca/Ti), and paleomagnetic inclination (see Table 1 for correlation data source). Mn variability is compared to LR04 d18O record (Lisiecki and Raymo, 2005). Major Mn peak intervals are indexed from a to o. Arrows indicate the prominent inclination drop. Marker foraminiferal species B. aculeata, T. quinqueloba, T. eglida, and coccolith E. huxleyi are indicated by vertical bars and stars, respectively. Also shown are Optically Stimulated Luminescence (OSL) ages for core 96/12 (Jakobsson et al., 2003) and Amino Acid Receminization (AAR) ages for core 8JPC (Kaufman et al., 2008).

ages have been reported for surficial sediments from the central Arctic Ocean in multiple studies (e.g., Polyak et al., 2009; Xiao et al., 2014). The lack of younger ages in some cases could be attributed to surface sediment disturbance by coring, however, its consistency suggests a more systematic pattern, such as very low sedimentation rate in the late Holocene. Ages of 12.5e10.5 ka obtained at 8e10 cm in ICE6 and 6e8 cm in E26, and subsurface samples in MA01 (2e4 cm and 4e6 cm), indicate early Holocene to deglacial sedimentation rates of 0.7e1 cm/kyr. A late MIS 3 age of 30.9 ka at 6e8 cm in core MA01 suggests a hiatus or strongly condensed deposition during the Last Glacial Maximum (LGM). Further downcore to 20e30 cm, 14C ages range between 35.8 and 47.1 ka indicating deposition during MIS 3. The closely spaced and occasionally reversed ages across these intervals may result from relatively high sedimentation rates combined with bioturbation characteristic of €wemark et al., 2012) (Table 2). The depositional brown units (e.g., Lo event marker W3 rich in detrital carbonate was dated to ~40.8 cal ka in core MA01 (10e12 cm), similar to previous results from the study region (e.g., Stein et al., 2010), while a somewhat older age of ~47 cal ka was obtained from a correlative layer at 28e30 cm in core E26. We note that calibration of pre-LGM 14C data for Arctic Ocean sediments involves considerable uncertainties, and resulting ages should be treated as approximate estimates. 4.2. Lithostratigraphy Core lithologies are characterized by interbedding of brownish and yellowish-grayish units (Fig. 2), as widely recognized in the Arctic Ocean (e.g., Phillips and Grantz, 1997; Polyak et al., 2004, 2009; Stein et al., 2010). Mn content covaries with these color

cycles, with enriched Mn in darker colored brown layers (Fig. 2). This result is consistent with earlier studies from multiple sites across the central Arctic Ocean (e.g., Polyak et al., 2009; Stein et al., 2010). Coarse fraction contents (>150 mm) range between 0 and 20% in cores ICE6 and E26, and between 0 and 30% in core MA01. Coarse fraction is generally more abundant in the upper part of the stratigraphy, peaking mainly close to lithological boundaries between brown and gray intervals. Most major increases in coarse fraction content from Mendeleev Ridge cores co-occur with Ca peaks, whereas in core ICE6, the coarse fraction is more strongly related to Zr variability (Fig. 2). Overall Zr shows more frequent fluctuations than Ca, with a notable change in core MA01 from a low content below 320 cm to high and variable content in the upper stratigraphy.

4.3. Magnetostratigraphy Paleomagnetic inclination consistently shows positive polarity in the upper part of all three studied cores (Fig. 2). This pattern is consistent in multiple cores investigated across the Arctic Ocean (Fig. 3). A prominent paleomagnetic inclination drop is clearly identified in cores ICE6, MA01 and E26 at 326 cm, 179 cm and 225 cm, respectively. Below this level, all three cores display distinct polarity fluctuations similar to previously studied Arctic Ocean cores (Jakobsson et al., 2000; Spielhagen et al., 2004; O'Regan et al., 2008; Xuan and Channell, 2010). Shallow inclination at core tops could be related to high water content in surface sediment, and/or disturbance by coring process, and is therefore not considered in interpretation.

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4.4. Foraminifera abundances and biostratigraphic marker occurrences Foraminifera occur discontinuously in the upper stratigraphy in all three cores, with their abundances increasing markedly in Mnenriched layers, and drop to nearly complete absence further down-core (Fig. 2). Foraminifera in core ICE6, from a greater water depth and a more northerly location, are overall about twice less abundant than in cores MA01 and E26. Despite their discontinuous occurrence in our records, we identified several foraminifera and ostracod species that are of stratigraphic significance for regional correlations (Ishman et al., 1996; Polyak et al., 2004; Hanslik et al., 2013; Cronin et al., 2013, 2014; Lazar and Polyak, 2016) (Fig. 2). In core E26, that has the best preservation of microfossil assemblages of the studied cores, the index species include benthic foraminifers Oridorsalis tener, Bulimina aculeata, Bolivina arctica, and Pullenia bulloides, ostracods Acetabulastoma arcticum and Polycope spp., and planktic foraminifera Turborotalita quinqueloba. Lesser amounts of marker microfossils were found in core MA01, possibly due to a more northerly core location and/or larger water depth, which may affect productivity and carbonate preservation. In particular, B. aculeata, known to be depth-dependant (Ishman et al., 1996; Polyak et al., 2004) is absent in core MA01. Only O. tener and T. quinqueloba were found in core ICE6. The shorter stratigraphic occurrence of T. quinqueloba in this core may be related to stronger dissolution in the basin. The other Turborotalita species/morphotype T. eglida, reported from core 6JPC raised from the top of Mendeleev Ridge (Cronin et al., 2013), was not found in the studied cores. Among the 30 smear slides from the three studied cores investigated for coccoliths, E. huxleyi was found in two samples (84e86 cm, 118e120 cm) in core MA01 (Fig. 2b). 5. Discussion 5.1. Stratigraphic controls 5.1.1. Radiocarbon ages While the Holocene and MIS 3 intervals are generally captured by 14C ages in cores under study, the downcore age distribution suggests extremely low sedimentation rates or a complete hiatus during the LGM. A similar pattern is widely recognized in multiple cores from various water depths across the central Arctic Ocean, where it may indicate non-deposition due to either a persistent pack-ice cover or an ice shelf capping the basin (Polyak et al., 2004, 2009; Hanslik et al., 2010; Wang et al., 2013, 2018; Cronin et al., €wemark et al., 2016; Chiu et al., 2017). A direct impact of 2014; Lo extended grounded ice could not affect the cores under study that are situated at water depths exceeding the depth range of glacigenic seafloor bedforms (Niessen et al., 2013). To confirm such nondeposition during this period, adequate 14C dating is needed for the discrimination of brown units, corresponding to MIS 1 and 3 in cores from the western Arctic Ocean (e.g., Polyak et al., 2004; Stein et al., 2010). While detailed age constraints are not available for older sediments, depositional hiatuses cannot be excluded for previous glacial maxima as well. 5.1.2. Longer Pleistocene stratigraphy 5.1.2.1. Stratigraphic proxies. Beyond the 14C range, we rely on multi-proxy regional correlations to develop an age framework for the longer stratigraphy recovered. We base the correlation primarily on the downcore distribution of Mn, Ca and paleomagnetic inclination, and the occurrence of biostratigraphic markers (Fig. 3).

Mn variations represent the most continuous and clear cyclicity in Arctic Ocean records due to Mn enrichment in inferred interglacial/interstadial intervals as demonstrated in multiple studies € wemark et al., 2014; Schreck et al., 2018; (Jakobsson et al., 2000; Lo Ye et al., 2019). This approach, however, can be complicated by diagenetic alteration of sedimentary Mn content in relation to €rz et al., 2011; Sundby et al., redox conditions in pore waters (Ma 2015; Meinhardt et al., 2016). While some degree of alteration may be widespread in sediments, major controls on Mn diagenesis are high fluxes of particulate organic carbon and methane seepages, both characteristic of continental margins rather than the central Arctic Ocean (Sundby et al., 2015). Geochemical data also show that even in oxygen-depleted pore waters it is difficult to completely dissolve or precipitate new Mn layers (Sundby et al., 2015; Meinhardt et al., 2016). These patterns are consistent with the distribution of other redox-sensitive elements, such as reactive iron €rz et al., 2012; Ye et al., 2019). and cerium, in Arctic sediments (Ma Within the central Arctic Ocean, diagenetic processes are potentially more significant in the TPD domain due to higher sediment and organic matter exports from the Siberian margin. For example, diagenetic imprints are apparent in core 96/12 at the interval of estimated MIS 5a, signified by high foraminifera abundance but low Mn content (Jakobsson et al., 2000, 2001). In contrast, sediments from the BG domain including parts of Mendeleev Ridge and Makarov Basin away from the Siberian margin do not show significant Mn remobilization in their pore waters (C. €rz, pers. comm., 2019). This pattern supports correlation of MnMa enriched layers across the central Arctic Ocean, especially within the BG domain, as exemplified by a consistent stratigraphic pattern € wemark et al., 2014; Schreck et al., 2018; Wang in multiple cores (Lo et al., 2018). This correlation is further corroborated by cooccurrence of Mn peaks with independent interglacial/interstadial proxies such as microfossils and/or bioturbation structures € wemark et al., 2012; Schreck et al., 2018). (e.g., M€ arz et al., 2011; Lo In our records, Mn highs consistently co-occur with foraminifera peaks in the upper stratigraphy indicating a lack of considerable diagenetic bias on Mn signal at least on glacial-interglacial time scales. In older sediments where foraminifera are absent, the Mn pattern is evaluated by a comparison with glacial proxies, such as Ca, Zr, and coarse fraction content. These proxies are not affected by diagenetic imprints and typically peak in Mn-poor intervals. Other proxies also mark high-Mn intervals, notably 10Be records that peak in high-productivity conditions (Eisenhauer et al., 1994; Aldahan and Possnert, 1998; Spielhagen et al., 2004). In comparison to Mn, bulk Ca can be used as a deglacial provenance proxy for detrital carbonate, mostly dolomites from the Canadian Arctic eroded by the North American ice sheets (Bischof et al., 1996; Polyak et al., 2009; Stein et al., 2010; Bazhenova et al., 2017). While total Ca can potentially bear a mixed signal from biogenic and detrital carbonates, biogenic components are usually minor in central Arctic Ocean sediments (Polyak et al., 2009; Stein et al., 2010; Dong et al., 2017). Furthermore, major Ca peaks are typically decoupled from the foraminiferal distribution, and their occurrence extends beyond the range of calcareous microfossils (e.g., Fig. 2). The uniqueness of the major carbonate (dolomite) provenance area in the Canadian Arctic and the wide dispersal of this material by icebergs via the BG circulation or its glacial counterparts (Bischof and Darby, 1997; St€ arz et al., 2012) make Ca a useful correlation proxy. Paleomagnetic data, in particular the inclination record, provides an independent correlation tool. In some magnetostratigraphic studies major inclination shifts in Arctic records were interpreted as true geomagnetic polarity inversions (e.g., Herman, 1970; Clark et al., 1980; Liu et al., 2019). This inference supports

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the age model with very low, mm/kyr-scale sedimentation rates throughout the Arctic Ocean notwithstanding geographic and bathymetric variability in depositional settings. Paleomagnetic investigations on cores from areas with less sediment-starved environments reveal multiple inclination swings, such as seen in our data (Fig. 2) (e.g., Jakobsson et al., 2000; Spielhagen et al., 2004). This pattern was attributed to short-term geomagnetic excursions, although their nature in the Arctic is not well understood, and their comparison with excursions from other regions (e.g., Laj and Channell, 2007) is not straightforward. Nevertheless, these results agree with higher, mostly cm/kyr-scale sedimentation rate estimates from other geochronological proxies (Nowaczyk and Baumann, 1992; Jakobsson et al., 2001; 2003; Kaufman et al., 2008; Backman et al., 2009), paleoclimatic cyclicity (Polyak et al., 2009; Stein et al., 2010), and longer stratigraphic and geophysical data (e.g., Backman et al., 2004; Dipre et al., 2018). Yet other paleomagnetic studies conclude that the inclination signal in Arctic cores can be imprinted by diagenetic processes likely related to bottom water redox conditions (Channell and Xuan, 2009; Xuan and Channell, 2010), although a recent study challenges this inference (Liu et al., 2019). However, the core stratigraphy in Liu et al. (2019) was based on correlations of paleomagnetic inclinations only, without considering comparisons of other lithologic parameters. Thus, the stratigraphic interpretaton of Liu et al. (2019) needs further evaluation. As a result of these unresolved issues, the application of paleomagnetic inclination variations for Arctic stratigraphic analysis is still limited, but it does represents a useful correlation tool. While correlation of individual paleomagnetic inclination excursions can be biased by changes in lithology and sedimentation rates, the prominent inclination drop has a consistent position in the upper part of the stratigraphy (e.g., Fig. 2). Somewhat less distinct paleomagnetic inclination patterns in 6JPC and 8JPC records in comparison to our data are possibly related to processing treatment of the full inclination data with principal component analysis (Channell and Xuan, 2009; Xuan and Channell, 2010) rather than using inclination measured after a weak (10 mT) demagnetization as in this study. 5.1.2.2. Regional correlation. The regional correlation shown in Fig. 3 connects Mendeleev Ridge cores with central Lomonosov Ridge via Makarov Basin and with Canada Basin. Mn variations, with major high-Mn intervals labeled consecutively from a to r, provide the backbone for this correlation due to their consistent pattern among the studied and referenced cores. Interval a is dated to MIS 1e3 by AMS 14C (Table 2). Interval b in the studied cores includes five Mn peaks consistent with the corresponding interval in referenced cores, especially distinct in cores BN05, 8JPC, and the 10Be record in core PS2185. Core 96/12 exhibits a slightly different pattern possibly due to a diagenetic reduction of Mn at the top of this interval, as suggested by foraminiferal abundance in this core (Jakobsson et al., 2000, 2003). In core 6JPC, the upper stratigraphy including interval b is strongly condensed, which is interpreted as winnowing at the shallow ridge tops (Cronin et al., 2013). Interval b has been dated by optically stimulated luminescence (OSL) in core 96/12 (Jakobsson et al., 2003) and by 14C-calibrated amino acid raceminization (AAR) in the same core and in core 8JPC (Kaufman et al., 2008). Both approaches attributed this interval to MIS 5. In addition, two Ca peaks in the upper and lower parts of interval b can be correlated throughout the studied and referenced cores, where the lower Ca peak corresponds to the Pink-White layer 2 (Clark et al., 1980), tentatively attributed to MIS 5d (Stein et al., 2010). In addition to Ca peaks, interval b contains a number of biostratigraphic correlation markers (Figs. 2 and 3). Benthic

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foraminifera O. tener extends downcore to the middle of this interval, consistent with previous studies (Polyak et al., 2004; Hanslik et al., 2013; Lazar and Polyak, 2016). B. aculeata is restricted to the upper two Mn peaks in interval b at intermediate water depths including cores E26, 6JPC, and 96/12 (Jakobsson et al., 2001; Polyak et al., 2004; Cronin et al., 2014; Wang et al., 2018). A distinct B. aculeata event is commonly attributed to MIS 5a (Backman et al., 2004; Polyak et al., 2004; Cronin et al., 2014), consistent with a suggested correlation of the Lomonosov Ridge records to a welldated core from the Barents Sea slope (Matthiessen et al., 2001; Backman et al., 2004). In addition, the interval with B. aculeata in Baffin Bay has also been dated to MIS 5a (Fillon and Duplessy, 1980). Planktic foraminifera T. quinqueloba is found in two distinct peaks in cores MA01, E26, 8JPC (Figs. 2 and 3), and Lomonosov Ridge cores closer to Greenland, where these T. quinqueloba peaks were attributed to MIS 5a and 5e, (Nørgaard-Pedersen et al., 2007; Adler et al., 2009; O'Regan et al., 2019). In addition to foraminiferal markers, coccolith E. huxleyi was found within the 3rd and 5th Mn peak in interval b, consistent with its stratigraphic distribution in the Lomonosov and Mendeleev Ridge cores (Jakobsson et al., 2000; Spielhagen et al., 2004; Backman et al., 2009) (Fig. 3). As the first appearance of E. huxleyi in the global ocean is dated to ~270 ka in MIS 8 (Thierstein et al., 1977), its occurrence in Arctic Ocean interglacial sediments attests to their age at least not older than MIS 7. Trans-Arctic studies have suggested its first occurrence age as MIS 5 (Gard, 1993; Backman et al., 2009). The prominent paleomagnetic inclination drop found across the central Arctic Ocean cores provides a correlation marker in the high Mn interval c. This step inclination change may be related to a pronounced paleomagnetic excursion (e.g., Jakobsson et al., 2000) or diagenetic alteration of magnetic carriers (Channell and Xuan, 2009; Xuan and Channell, 2010). Regardless of the origin of this marker feature, it is attributed to MIS 7 in recent studies (e.g., O'Regan et al., 2008; Polyak et al., 2009; Stein et al., 2010). The occurrence of P. bulloides and the range of ostracod A. arcticum to the base of interval c in core E26 (Fig. 2) are both in agreement with earlier studies in the western Arctic Ocean proposing the MIS 7 age for these markers (Cronin et al., 2014). A peak of planktic foraminifera species Turborotalita egelida (possibly a morphotype of T. quinqueloba) was identified in multiple cores across the western Arctic Ocean and southern Lomonosov Ridge off Greenland (Herman, 1970; Hanslik, 2011; Cronin et al., 2013, 2014; Wang et al., 2018; O'Regan et al., 2019). The unusual presence of this subpolar foraminifer may indicate a substantial warming in the Arctic Ocean during the interval of its occurrence, consistent with its cyclostratigraphic assignment to MIS 11 (Cronin et al., 2013; O'Regan et al., 2019). While the peak with T. egelida was not found in the three studied cores, it is present in core AF00-08 located nearby MA01 but at a shallower water depth of 1530 m (Fig. 1, Table 1, Krylov et al., 2011). This peak also appears in core JPC6 from the top of the ridge farther south (Cronin et al., 2013). Shallower depths provide better calcite preservation, which may be especially critical for this thin-walled planktic species. The reason for the absence of T. egelida in E26 is less clear and could be related to its relative proximity to the continental margin. Regardless of this uncertainty, by means of Mn and Ca variations, this peak correlates reliably to interval e (Fig. 3). No biostratigraphic correlation markers were identified below interval e, where calcareous microfossil tests are mostly absent probably due to dissolution, being commonly replaced by agglutinated foraminifera (e.g., Poore et al., 1993; Cronin et al., 2008). The most recognizable lithostratigraphic markers in the lower stratigraphy is the first increase in detrital carbonate content (Polyak et al., 2009, 2013; Stein et al., 2010). This event was attributed to

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MIS 16, the time of the Northern Hemisphere glacial expansion, notably the first North American “super glaciation” as indicated by the North Atlantic records (Hodell et al., 2008; Stein et al., 2009; Channell et al., 2012). This marker is registered in cores MA01 and E26 by a distinct Ca peak that correlates to other reference cores from Mendeleev and Lomonosov ridges at the bottom of interval j (Fig. 3). Below this level, our stratigraphy is tentatively based on Mn cycles, with uncertainty in their attribution in the “pre-100-ka world”, when the glacial-interglacial cyclicity was predominantly paced by ~41-ka obliquity cycles (e.g., Clark et al., 2006). We note that our cores do not reach the major lithological boundary related to the glacial-preglacial (Middle to Early Pleistocene) transition. This boundary is characterized by sediment lithology changes from alternating grayish and dark brown layers with dispersed coarse grains in the upper unit, to more homogenous, dark brown sediment without IRD in the lower unit, as shown for cores from the Northwind Ridge (Polyak et al., 2013; Dipre et al., 2018). 5.1.3. Age model An overall robust correlation among the studied and reference cores suggests a generally similar stratigraphic variability in depositional environments across the central Arctic Ocean during the studied time interval. The alignment of this variability, expressed primarily in Mn variations, with global climate LR04 stack (Lisiecki and Raymo, 2005) is in agreement with the position of stratigraphic markers that confirm the correlation and provide additional age estimates (Fig. 3). While verification by more independent chronostratigraphic tools is not yet attainable, this age framework is consistent with multiple Arctic stratigraphic studies and makes reasonable paleoclimatic sense. We therefore believe it is suitable for investigation of proxy variability in the time domain on orbital time scales, keeping in mind the potential age uncertainties. Based on the developed age model, cores ICE6, MA01 and E26 extend back to ~490 ka (MIS 13), ~840 ka (MIS 21), and ~680 ka (MIS 17), with average sedimentation rates of 0.87, 0.64 and 0.56 cm/kyr, respectively (Figs. 4 and 5). Core MA01 thus represents a nearly complete Mid-to Late Pleistocene record with the best average resolution among the published Mendeleev Ridge data, although sedimentation rates were higher during the last ~450e500 ka in core E26. Yet higher rates captured for the same time interval in

Fig. 4. Age-depth profiles for sediment cores under study. The stratigraphically longest core MA01 is shown by a thicker red line. (For interpretation of the references to color in this figure legend, the reader is referred to the Web version of this article.)

core ICE6 are attributed to more sediment inputs from the Siberian margin close to the main pathway of the TPD, compared to Mendeleev Ridge controlled by an interplay between the TPD and BG. We note that while estimated sedimentation rates outline the general depositional history, more details may not be discernible due to potentially uneven sediment deposition, especially within glacial intervals. This feature includes, notably, hiatuses corresponding to glacial maxima and enhanced sedimentation during deglaciations (Adler et al., 2009; Polyak et al., 2009; Schreck et al., 2018).

5.2. Orbitally paced interglacial variations revealed by the Mn record Mn content is consistently high, although variable during interglacials and major interstadials (depending on sedimentation rates), with only minimum background values during glacial intervals (Figs. 3 and 5). This pattern has made records of Mn or its color proxy a primary choice for time-series analysis of Arctic Ocean cores (O'Regan et al., 2008; Adler et al., 2009; Wang et al., 2018). Under modern conditions most of Mn inputs to the Arctic Ocean originate from rivers and coastal erosion, and are then recycled on the shelf under redox environments controlled by seasonally high €rz et al., 2011; Middag et al., 2011; Macdonald and productivity (Ma € wemark et al., 2014; Ye et al., 2019). Particulate Mn Gobeil, 2012; Lo oxy(hydr)oxides are then transported to the Arctic interior by currents and sea ice, and are subsequently deposited in the deepsea sediment under oxic conditions. In contrast, during glacial times, low sea-level combined with growth of circum-Arctic ice sheets drastically reduced riverine inputs and export via shallow continental shelves, and thus the supply of Mn to the central Arctic Ocean. Based on these patterns, glacial-interglacial Mn cyclicity in deep-sea sediments is primarily explained by variations in Mn inputs associated with climate and sea level changes. The “on and off” mode of Mn content on interglacial-glacial time scales (Figs. 5 and 6) indicates a non-linear response of Mn deposition to climate/sea-level fluctuations. Abrupt Mn changes correspond to global sea levels around 50~-60 m. This depth interval encloses the predominant portion of the Siberian continental margin (Fig. 1; Jakobsson, 2002), which is most sensitive to sealevel changes, and thus Mn recycling on the shelves and its export to the central Arctic Ocean. Furthermore, this depth level may have a critical effect on the Arctic Ocean circulation via opening and closure of Bering Strait, as well as control on the Barents Sea branch of Atlantic water inflow. Along with sea level, climatic variability may also affect Mn distribution, including changes in the hydrological cycle, biological production on the shelves, and sea-ice and circulation conditions. These factors control Mn inputs from rivers, recycling on the shelves, and transport to the central Arctic Ocean, respectively (Middag et al., 2011; Macdonald and Gobeil, 2012; Ye et al., 2019). Climatic control may account for apparent discrepancies between Mn and sea-level records, such as high Mn contents in some interstadials with only modest sea-level increase, notably in MIS 3 (Figs. 5 and 6). Both marine (Cronin et al., 2012) and terrestrial data (Andreev et al., 2011; Wetterich et al., 2014) indicate a significant effect of this interstadial on Arctic environments. The strong expression of MIS 3 climatic amelioration in the Arctic may be related to high seasonality modeled for northern high latitudes (Van Meerbeeck et al., 2009), which should have intensified Mn remobilization and transportation in summer. With the uncertainties outlined above, we relate Mn variations in core MA01 at interglacial-glacial time scales to global stack records and orbital data (Fig. 6). The Mn wavelet spectrum power at

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Fig. 5. Proxy variations in sediment cores under study, as in Fig. 2, plotted vs. age and compared to the LR04 d18O record (Lisiecki and Raymo, 2005).

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Fig. 6. MA01 Mn/Ti and Zr/Ti time series and paleoclimatic comparisons: (a) orbital precession and summer insolation at 65
N (Laskar et al., 2004), (b) Chinese speleothem d18O record (Cheng et al., 2016) (left) and East Asian Summer (left) and Winter (right) Monsoon proxy records derived from Northern China loess grain size (Hao et al., 2012), (c) MA01 records of Mn/Ti, Ca/Ti and Zr/Ti; (d) global records of sea level (Miller et al., 2011) and LR04 d18O (Lisiecki and Raymo, 2005); 50-m sea level is marked. (e) Wavelet power spectra of Mn/Ti and Zr/Ti. Solid black contours indicate areas with power above 95% confidence level. Translucent areas on either side of the wavelet power spectrum indicate the “cone of confidence” where edge effects become important (Grinsted et al., 2004). White dashed lines represent the major orbital periods of eccentricity (100 kyr), obliquity (41 kyr) and precession (19e23 kyr).

the eccentricity band (~100 kyr) shows an increase towards the Late Pleistocene, most prominent during the last three glacial cycles (MIS 1e9). At the obliquity band (~41 kyr), high power is most evident during MIS 13e15. While at the precession band (19e23 kyr) most interglacials exhibit high power after MIS 16. Overall, strong resultant eccentricity and precessional signals are consistent with earlier Arctic time series generated on Mn, color reflectance, bulk density, rock magnetic properties, or microfossil abundance in the upper stratigraphy (O'Regan et al., 2008; Adler et al., 2009; Marzen et al., 2016; Wang et al., 2018). At the same time, our results show a more comprehensive picture than more condensed (Wang et al., 2018), stratigraphically shorter (Adler et al., 2009) or composite stacked records (Marzen et al., 2016). For example, compared to a condensed Mn record from Alpha Ridge with a precessional cyclicity seen only between ~150 and 350 ka (Wang et al., 2018), the MA01 record shows a strong signal throughout most of the Late to Middle Pleistocene. The strong precessional signal is typical for low latitude as a response to seasonal insolation contrast (Berger et al., 1984). Interglacial Mn peaks in core MA01 correspond to precession

minimum and Northern Hemisphere summer insolation maximum intervals (Laskar et al., 2004), thus closely resembling variability of the Asian Summer Monsoon record dominated by low latitude climate processes (Hao et al., 2012; Cheng et al., 2016) (Fig. 6). This pattern suggests a persistent precessional impact of low latitude processes on high Arctic environment in comparison to a commonly inferred obliquity control on high latitude climate during the Pleistocene (Huybers, 2006). A connection with low latitudes may be realized via heat and moisture transport from lower latitudes and the Pacific and/or Atlantic Ocean that affects Arctic ice melt and river discharge (e.g., Ye et al., 2004), as well as the modulation of sea ice concentration via atmospheric circulation (Grunseich and Wang, 2016), and thus Mn delivery. While transport and precipitation may occur mostly in winters as snowfall, the highest discharge rates characterize the spring freshet (late MayeJune) (Holmes et al., 2002). Further, mobilization and delivery of Mn from permafrost melting and coastal erosion, as well as its subsequent recycling mediated by biological processes on the shelf, operate in summer. These processes are related to higher temperatures and retreating sea ice

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controlled by boreal summer insolation (Tuenter et al., 2005) as well as atmospheric transport from extra-Arctic oceanic regions (Crasemann et al., 2017). These connections may explain the strong precessional signal in the Mn record during interglacials. We note that a possibility of the influence of Arctic processes on low-latitude circulation, such as monsoons, via atmospheric teleconnections as suggested by some climate modeling studies (e.g., Guo et al., 2014; He et al., 2017), cannot be discarded either. More investigation is needed to comprehend the involved linkages and their causality. 5.3. Proxy history of glaciation dynamics and related circulation changes 5.3.1. IRD and Ca inputs Three proxies used to characterize the glacial/deglacial depositional conditions are coarse sediments, Ca and Zr. A major step change in the western Arctic Ocean depositional history attributed to MIS 16 at the end of the Mid-Pleistocene Transition (Polyak et al., 2009, 2013; Dipre et al., 2018) is expressed in the studied cores as increased IRD deposition (Fig. 5). This step change is also marked by a pronounced Ca peak indicative of LIS discharge (Figs. 5 and 7). Further increase in content and variation of coarse sediment fraction occurs around MIS 12 (Fig. 5). This timing corresponds to a climatic transition known as the Mid-Brunhes Event (MBE) that is signified by an increase in amplitude of glacial-interglacial cycles in high latitudes (Jansen et al., 1986; EPICA, 2004; Candy and McClymont, 2013), possibly including the Arctic Ocean (Cronin et al., 2017). This IRD pattern indicates enhanced circum-Arctic ice sheet dynamics, which is further exemplified by an increase in the content of Ca and Zr during glacial/deglacial intervals around the same time (Figs. 5 and 7). As these mineral elements are considered as provenance proxies for Laurentide and Eurasian glacial discharge, respectively, their records may be used for tracking the relative magnitude of these ice sheets. Detrital carbonate deposition during glacial intervals of the LIS primarily occurred in cores under study in MIS 16, 12, 10, 8, and the last deglaciation (Figs. 5 and 7). This pattern appears to be consistent with the North Atlantic records that show detrital carbonate deposition mostly during glacials of MIS 16, 12, 10 and 8, as well as the millennial scale Heinrich Events in the last glacial cycle (Hodell et al., 2008; Obrochta et al., 2014) (Fig. 7). In addition, Ca peaks in the Arctic cores are pronounced in the stadials of interglacial intervals since MIS 13, especially consistent in MIS 5 and 3 (Fig. 7). This pattern indicates considerable differences in LIS dynamics on the Arctic and Atlantic sides. In particular, a prominent, widespread Arctic carbonate deposition event known as PW2 layer in the lower part of MIS 5 is not identified in the North Atlantic record, but may be synchronous with Eurasian ice sheet growth in the Arctic during MIS 5d (Spielhagen et al., 2004; Svendsen et al., 2004). MIS 5d was a period of minimum summer insolation, which may have facilitated the growth of Arctic ice sheets on both North American and Eurasian continents (Lachniet et al., 2014). It must be noted that, despite the demonstrated usefulness of Ca peaks as lithostratigraphic markers in western Arctic Ocean sediments, a detailed correlation suggests that amplitudes of Ca peaks at the same stratigraphic intervals may vary between cores (Figs. 3 and 7). This variance is not surprising as the expression of glacial discharge events in sediments depends on ocean circulation and depositional patterns, which vary both geographically and temporally. For example, the PW1 detrital carbonate layer is expressed as a prominent Ca peak slightly below the paleomagnetic inclination drop in MIS 7 in cores BN05 and 6JPC (Figs. 3 and 7, and Suppl. Fig. 1). In core MA01, however, the highest amplitude Ca peak (PW1.1 in Fig. 2b) occurs above the inclination drop, correlative to the second Ca peak in core 6JPC. This pattern indicates

Fig. 7. Ca/Ti and Zr/Ti variations in cores BN05, E26, MA01, ICE6, and 96/12, compared to Ca/Sr record from the North Atlantic site U1308 (Hodell et al., 2008), and LR04 d18O record (Lisiecki and Raymo, 2005).

multiple iceberg discharge pulses in an interval that may look like a single depositional event in stratigraphically condensed records. 5.3.2. Zr variation Distribution of Zr in the studied cores shows overall more variability than Ca. This pattern is expected from a larger choice of sources for Zr-carrying heavy minerals, which hydrodynamically behaves like coarse-grained siliciclastic material exemplified by quartz (e.g., Vogt, 1997). In some cases, such as in core BN05 (Fig. 7), Zr and Ca show mostly synchronous variations suggesting a common control of the BG on their deposition. However, overall

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sediments from areas affected by the TPD show higher contents of quartz along with various Siberian source indicators, suggesting an affinity of Zr with these sources (Schoster, 2005; Haley et al., 2008; Krylov et al., 2008; Immonen et al., 2014). The most pronounced increase in both total coarse sediment and Zr content occurs in cores along the main TPD pathway (Makarov Basin and central Lomonosov Ridge) and close to the Siberian margin in later Pleistocene glaciations starting at MIS 6 (Figs. 5 and 7). This pattern is interpreted in multiple studies as an expansion of Eurasian ice sheets into the Arctic (Spielhagen et al., 1997, 2004; Jakobsson et al., 2001; O'Regan et al., 2008). Inputs of siliciclastic IRD during these intervals extend far beyond regular TPD-affected areas into the western Arctic Ocean, concurrent with lower occurrence of the LIS provenance markers such as detrital carbonates (Dong et al., 2017; Wang et al., 2018). The opposing Zr and Ca variations suggest an asynchronous development of North American and Eurasian ice sheets, as also inferred in a number of data and modeling studies (e.g., Liakka et al., 2016; Rohling et al., 2017). For example, high Zr and low Ca values in MIS 6 are consistent with evidence for relatively low LIS and high EAIS volumes (Obrochta et al., 2014; Colleoni et al., 2016; Rohling et al., 2017). In addition to direct glacial inputs, preferential Eurasian or North American ice-sheet growth is likely to have affected Arctic Ocean circulation, and thus delivery of sediment from the North American vs. the Siberian side. In particular, a shrinking BG and a TPD expansion towards the western Arctic Ocean inferred for some €rz et al., 2012) would have glacial intervals (Bischof et al., 1996; Sta enhanced sediment transportation from the Siberian Shelf. A prominent detrital carbonate peak in MIS 16 in cores from central Lomonosov Ridge (Figs. 3 and 7) suggests an extension of a BG-type circulation to this area, currently controlled by TPD. In contrast, the following glacial stages MIS 12, 10, and 8 are marked by detrital carbonate occurrence only in the western Arctic Ocean, indicating a growing influence of the Eurasian glaciation inputs by way of TPD. This notion is consistent with a lagged increase in Zr/Ti in core MA01 compared to that in core E26 located closer to the Siberian margin (Figs. 5 and 7). Nevertheless, during some glacial events, such as a stadial in the lower part of MIS 5, detrital carbonates are found in both western Arctic and central Lomonosov Ridge cores (e.g., Spielhagen et al., 1997, 2004, Fig. 3; Suppl. Fig. 1). This broad geographic distribution of LIS provenance markers indicates a complex interaction of North American and Eurasian glaciations and the attendant circulation systems. Despite high variability in glacial sedimentation rates, such as potential hiatuses during glacial maxima and enhanced deposition during deglaciations, the Zr/Ti time series in MA01 appears to show an interpretable paleoclimatic picture (Fig. 6). Zr/Ti shows a strong increase since MIS 13, but no clear cyclicity prior to MIS 7. This pattern is consistent with a long-term growth history of the Eurasian Ice Sheet that has significantly extended since MIS 6 (e.g., Spielhagen et al., 2004). A lack of distinct cyclicities in the older stratigraphy may also result from low sedimentation rates during early glaciations. In the overall higher-resolution sections after MIS 7, the Zr record shows high power at the precession band similar to Mn/Ti, but not limited to interglacials. Multiple Zr peaks discernible in the most expanded glacial intervals provide evidence of multiple glacial discharge events. In particular, at least three distinct Zr peaks in MIS 6 (Figs. 5 and 6) indicate a considerable variability within this glacial epoch. This pattern may reconcile seemingly conflicting evidence of extensive ice sheet/ice shelf across the central Arctic Ocean (Jakobsson et al., 2014, 2016), and relatively open water conditions at the Siberian margin during MIS 6 (Stein et al., 2017). An overall high Zr content during glacial intervals since MIS 12

tentatively correlates to the East Asian Winter Monsoon (EAWM) maxima identified in Northern China loess records (Hao et al., 2012) (Fig. 6), thus providing a possible link to low-latitude impacts. Intensified EAWM is interpreted to indicate enlarged Northern Hemisphere ice sheets (Hao et al., 2012), showing multiple ice advances and retreats during glacial intervals. Similarly, the glacial development pattern for the European Ice Sheet (western part of the EAIS) was related to a monsoonal forced moisture transportation in the eastern North Atlantic during the past ~250 ka (Kaboth et al., 2018). In Siberia, maximal precipitation occurs during winter and spring as snow/ice originating from Pacific and especially North Atlantic moisture sources (Ye et al., 2004; Aizen et al., 2005). These processes can potentially lead to development of ice sheets during cold climatic intervals (e.g., Stauch and Gualtieri, 2008). However, changes in snow/ice accumulation are unlikely to have a strong control on ice sheet variability due to relatively small variations of winter insolation at high latitudes. In contrast, low summer insolation at precession maximum intervals can efficiently prevent ice melt, and thus facilitate ice sheet growth (Lachniet et al., 2014). 5.4. Sea-ice and circulation changes inferred from biogenic proxies The transition to enhanced circum-Arctic ice sheet dynamics in the Middle Pleistocene, co-eval to the MBE and expressed in IRD and provenance markers, is accompanied by changes in biogenic proxies indicative of hydrographic and sea-ice conditions. In particular, changes in foraminiferal and ostracode assemblages, most pronounced at estimated MIS 11 and 7 (ca. 0.4 and 0.2e0.3 Ma), have been related to expansion of sea-ice over the western Arctic Ocean (Cronin et al., 2013, 2017; Polyak et al., 2013; Lazar and Polyak, 2016; Dong et al., 2017). The latter step change is especially conspicuous, being marked by a strong increase in calcareous microfossil abundances and several independent proxies such as an inclination drop and intensification of Mn-related color cyclicity (e.g., Jakobsson et al., 2000; Adler et al., 2009). In earlier studies, the correlative stratigraphic level was identified as a major lithological boundary (Herman, 1970). Cores MA01 and E26 show an overall strong increase in the numbers of calcareous microfossils, mostly foraminifers, at the level of MIS 7 with very low abundances to complete absence in underlying sediments (Fig. 5bec and 8b). This result is consistent among multiple central Arctic Ocean sites (Adler et al., 2009; Polyak et al., 2013; Lazar and Polyak, 2016). In core ICE6 (Fig. 5a) as well as prior data from the Makarov Basin (Krylov et al., 2011), foraminiferal presence is limited to an even shorter interval from upper MIS 5. As suggested in several studies, the rise in abundances of calcareous microfossils at younger stratigraphic intervals likely indicates stronger biogenic carbonate dissolution in the older strata (Cronin et al., 2008; Lazar and Polyak, 2016). Enhancement of biogenic carbonate preservation co-occurs with indicators of sea ice growth, such as foraminifers related to perennial vs. seasonal sea ice, e.g. Epistominella arctica vs. E. exigua (Fig. 8c; Lazar and Polyak, 2016), and ostracod Acetabulostoma arcticumand (Fig. 8d; Cronin et al., 2017). This co-occurrence suggests the likelihood of a causal link. A reduced sea-ice cover before the MIS 7 transitional event would favor a prolonged productivity season, and thus more organic carbon exported to the seafloor, with more corrosive interstitial waters resulting from degradation of this organic matter. A more restricted foraminiferal occurrence in Makarov Basin records may result from both lower productivity under more extensive sea-ice cover and/or stronger dissolution at greater water depths. In addition to organic matter fluxes, its decomposition is enhanced by well oxygenated bottom/interstitial water. A notion of

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relation to stronger deep-sea convection (Dong et al., 2017). This hydrographic setting would be favorable for improving the deep Arctic Ocean ventilation by an efficient sinking of well-oxygenated water. The cause of Arctic sea-ice expansion and related hydrographic changes in MIS 7 are yet to be comprehended. Growth of the circum-Arctic ice sheets could be one possibility as the accompanying atmospheric cooling and increased meltwater fluxes would facilitate sea-ice development. This inference is corroborated by the intensification of coarse sediment and Zr variability in cores more strongly affected by the TPD. This pattern is related to inputs from Eurasian ice sheets (Figs. 5 and 7), coeval with a rise in glacigenic IRD in interglacial sediments from Canada Basin, and is potentially indicative of a lingering LIS since MIS 7 (Dong et al., 2017). Another factor could be a strengthening of the Beaufort Gyre that would lead to more sea ice filling the western Arctic Ocean (Proshutinsky et al., 2009). 6. Summary and conclusions

Fig. 8. Stratigraphic records for cores MA01 and E26 compared to proxy records of sea ice conditions in the western Arctic Ocean exemplifying long-term middle Pleistocene changes: paleomagnetic inclination and foraminiferal abundance in MA01 and E26 (a, b), 6JPC foraminiferal (c; Lazar and Polyak, 2016) and stacked western Arctic Ocean ostracod (d; Cronin et al., 2017) sea ice proxy records, and the LR04 d18O record (e; Lisiecki and Raymo, 2005). Yellow shadings highlight major step changes. MBE: MidBrunhes Event. (For interpretation of the references to color in this figure legend, the reader is referred to the Web version of this article.)

a better ventilated deep Arctic Ocean before MIS 7 is consistent not only with stronger carbonate dissolution, but also with a change in bioturbation pattern from Chondrites-Skolithos to Planolites-Chondrites assemblages reported for cores from Northwind Ridge (Phillips and Grantz, 1997) and frequent paleomagnetic inclination reversals consistently recorded in Arctic sediments (e.g., Spielhagen et al., 2004; O'Regan et al., 2008; Polyak et al., 2009), including cores analyzed in this study (Fig. 8a). These inclination reversals were interpreted as transformation of magnetic minerals, presumably controlled by seafloor oxidation (Channell and Xuan, 2009; Xuan and Channell, 2010). We note that the inferred decrease in Arctic bottom water ventilation from the Middle to Late Pleistocene, in particular in Canada Basin, is opposite to the trend reported for the North Atlantic Ocean (Kawagata et al., 2005), which implies disconnected deep water histories between these oceanic basins. Under modern conditions, deep Arctic basins are filled with dense waters partially derived from sea-ice brine formation on the shelves (e.g., Rudels et al., 2000; Rudels, 2015). In the older part of the Pleistocene, less sea ice during interglacials may have weakened surface water stratification and enhanced air-sea gas exchange as well as brine production. This inference is consistent with sedimentologic data from Canada Basin, interpreted as an increased delivery of shelf sediments in pre-MIS 7 interglacials, in

Three sediment cores collected from Makarov Basin and Mendeleev Ridge in the Arctic Ocean off the Siberian margin by the CHINARE V and VII expeditions show a consistent stratigraphic variability of sedimentary proxies. Litho/biostratigraphic proxy data from these cores correlated to earlier investigated sediment records from the central Arctic Ocean enable paleoceanographic reconstructions for the Middle to Late Pleistocene (estimated ~0.8 Ma) epoch. Cyclic glacial-interglacial changes in lithostratigraphy exemplified by elemental Mn content were tuned to global climate variability. Precessional signal is identified in Mn variations during interglacials, with Mn peaks corresponding to Northern Hemisphere summer insolation maxima. This pattern is likely related to summer processes such as enhanced inputs from the coasts, biological production on the shelf, and efficient shelf-basin interactions. Resemblance of this variability to the Asian Summer Monsoon record suggests a potential link to low latitudes via atmospheric moisture and heat transport and prompts further investigation of the involved linkages. Glacial input indicators, such as coarse sediment fraction, Ca and Zr contents show intensified signals since MIS 16 that is consistent with the onset of LIS discharge into the North Atlantic. Further increases in these proxies during glacial intervals reflect enhancement of circum-Arctic glaciations since MIS 12, co-eval with intensification of glacial-interglacial cyclicity known as the MidBrunhes Event. Comparisons between Ca and Zr distribution patterns suggest the possibility of asymmetric development of the North American and Eurasian ice sheets. Multiple Zr peaks seen in the most expanded glacial intervals, such as MIS 6, provide evidence of multiple glacial events. We infer that ice sheet advances were facilitated by decreases in Northern Hemisphere summer insolation. In addition to changes in detrital inputs, transition to enhanced Arctic ice sheet dynamics is reflected in biogenic proxies, such as foraminiferal and ostracode assemblages. These changes, most pronounced at estimated MIS 11 and 7 (ca. 0.4 and 0.2e0.3 Ma), are likely related to expansion of sea ice over the western Arctic Ocean. The step change at MIS 7 is marked by a strong increase in calcareous microfossil abundances and by several independent lithological and paleomagnetic proxies. We infer that this transition was controlled by less corrosive/oxidative bottom waters resulting from a decrease in organic material exported to the seafloor with sea-ice expansion. Another factor could be weaker ventilation of bottom waters under perennial sea ice due to lower oxygen saturation of surface waters and reduced density-driven convection. Sea-ice

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development and related hydrographic changes were possibly controlled by growth of circum-Arctic ice sheets accompanied by atmospheric cooling and increased meltwater fluxes. We note that while sedimentary proxy variabilities in our data make reasonable paleoceanographic sense, their accurate relation to global paleoclimatic history requires more rigorous chronostratigraphic studies. The areas of particular attention may include the nature of paleomagnetic sedimentary properties, biostratigraphic events, radiogenic (e.g., uranium-series) tracers, and integration of sediment-core and geophysical records. Declaration of competing interest The authors declare that they have no conflicts of interests to this work. Acknowledgements This work was funded by the Chinese National Natural Science Foundation of China (No. 41776187 and 41030859), the Chinese project of Arctic Ocean marine geology investigation (CHINARE 2012-2017-03-02). LP's contribution was supported by the US National Science Foundation awards ARC-1304755, and the Tongji SKL fund MGK1801. LL's contribution was financially supported by “The Featured Areas Research Center Program” from the Ministry of Education in Taiwan, and by the Ministry of Science and Technology (MOST grant 106-2116-M-002-021). This work is a contribution to the 5th and 7th Chinese National Arctic Research Expeditions (CHINARE-2012 and -2016) conducted by R/V Xuelong. We thank Linsen Dong for sharing data from core BN05. Chuang Xuan is acknowledged for providing additional paleomagnetic data. Also acknowledged is the core repository of the Polar Research Institute of China for providing the sediment samples. We thank two anonymous reviewers for detailed comments that helped to improve the manuscript. Appendix A. Supplementary data Supplementary data to this article can be found online at https://doi.org/10.1016/j.quascirev.2019.106105. References Aagaard, K., 1989. A synthesis of the Arctic Ocean circulation. Rapp. P.-V. Reun. Cons. Int. Explor. Mer. 188, 11e22. Adler, R.E., Polyak, L., Ortiz, J.D., et al., 2009. Sediment record from the w Arctic ocean with an improved late quaternary age resolution: HOTRAX core HLY0503-8JPC, Mendeleev Ridge. Glob. Planet. Chang. 68, 18e29. Aizen, V.B., Aizen, E., Fujita, K., Nikitin, S.A., Kreutz, K.J., Takeuchi, L.N., 2005. Stableisotope time series and precipitation origin from firn-core and snow samples, Altai glaciers, Siberia. J. Glaciol. 51 (175), 637e654. Aldahan, A., Possnert, G., 1998. A high-resolution 10Be profile from deep sea sediment covering the last 70 ka: indication for globally synchronized environmental events. Quat. Geochronol. 17, 1023e1032.  Jakobsson, M., lfsson, O., Alexanderson, H., Backman, J., Cronin, T.M., Funder, S., Ingo € wemark, L., Mangerud, J., Ma €rz, C., Mo € ller, P., O'Regan, M., Landvik, J.Y., Lo Spielhagen, R.F., 2014. An Arctic perspective on dating Mid-Late Pleistocene environmental history. Quat. Sci. Rev. 92, 9e31. Andreev, A., Schirrmeister, L., Tarasov, P.E., Ganopolski, A., Brovkin, V., Siegert, C., Wetterich, S., Hubberten, H.-W., 2011. Vegetation and climate history in the Laptev sea region (arctic Siberia) during late quaternary inferred from pollen records. Quat. Sci. Rev. 30, 2182e2199. Backman, J., Jakobsson, M., Løvlie, R., Polyak, L., Febo, L.A., 2004. Is the central Arctic Ocean a sediment starved basin? Quat. Sci. Rev. 23, 1435e1454. Backman, J., Fornaciari, E., Rio, D., 2009. Biochronology and paleoceanography of late Pleistocene and Holocene calcareous nannofossil abundances across the Arctic Basin. Mar. Micropaleontol. 72, 86e98. Bauch, D., Carstens, J., Wefer, G., 1997. Oxygen isotope composition of living Neogloboquadrina pachyderma (sin.) in the Arctic Ocean. Earth Planet. Sci. Lett. 146, 47e58. € lemann, J., Andersen, N., Dobrotina, E., Nikulina, A., Kassens, H., 2011. Bauch, D., Ho

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