Core SWERUS-C3-31-PC (hereafter referred to as 31-PC) was acquired on Leg 2 of the SWERUS-C3 2014 expedition on IB Oden, which departed 21 August from Utqiaġvik (formerly Barrow), Alaska, and ended 3 October in Tromsø, Norway. The core was retrieved from the intersection of the southern Lomonosov Ridge and the Siberian continental margin, bordering the East Siberian and Laptev seas (79.91∘ N, 143.23∘ E; 1120 m water depth) (Fig. 1). It was collected with a piston corer (PC) with an outside/inside diameter of 110/100 mm, rigged with a 1360 kg core head. The unsplit sediment cores were allowed to equilibrate to room temperature (20 ∘C) and subsequently logged shipboard on a Multi-Sensor Core Logger (MSCL). The cores were split and described shipboard and imaged using a digital line-scanning camera.

Bulk density (ρB) and magnetic susceptibility (Bartington loop sensor) were measured at a downcore resolution of 1 cm on the MSCL (Fig. 2). Porosity (ϕ) was calculated from the MSCL measured bulk density using 1ϕ=ρG-ρB(ρG-ρF), where a constant fluid density (ρF) of 1.024 g cm-3 and grain density (ρG) of 2.67±0.01 g cm-3 were applied. The grain density was determined by 11 shore-based measurements made on freeze-dried sediments from 31-PC using a Micromeritics AccuPyc 1340 helium displacement pycnometer.

Sediment grain size (< 2 mm) was measured at a 2 cm downcore resolution using a Malvern Mastersizer 3000 laser diffraction particle size analyzer. Wet samples were immersed in a dispersing agent (< 10 % sodium hexametaphosphate solution) and placed in an ultrasonic bath to ensure full particle disaggregation before analyses.

Bulk total organic carbon (TOC) analyses were performed after freeze-drying and homogenizing 80 samples taken from 31-PC at 10 cm intervals (Fig. 2). A split of ∼ 10 mg of sediment was weighed in silver capsules and acidified with 3 M HCl to remove carbonates. The TOC of the samples was measured using a Carlo Erba NC2500 elemental analyzer in the Department of Geological Sciences, Stockholm University.

Eight bulk sediment samples and one marine mollusk shell were sent to NOSAMS for 14C dating (Table 1). Despite shipboard and shore-based efforts, no further carbonate microfossils were found in 31-PC that could be used for radiocarbon dating, including foraminifera, mollusks, and ostracods. In order to remove carbonates from the bulk sediment samples, the samples received HCl vapor treatment. Results were reported as conventional 14C years (Stuiver et al., 1977) and then converted into 14C calendar years (cal ka BP) using Calib 7.1 (Stuiver et al., 2018). Bulk 14C estimates were calibrated using the IntCal13 radiocarbon calibration curve (Reimer et al., 2013). To conservatively account for the unknown amount of 14C-dead carbon in the bulk sediment samples, we assigned a lower bound on the uncertainty of the bulk dates by arbitrarily subtracting 20 000 years from their median calibrated age (Table 1). We adopted this strategy to achieve highly non-informative priors for the proposed Bayesian age modeling procedure detailed in Sect. 2.4.

By contrast, the mollusk shell 14C estimate (Early Holocene in age) was calibrated against Marine 13 using a regional reservoir correction (ΔR) of 0-400+1000 years to account for unknown changes in the local marine reservoir correction. We deem this choice of ΔR uncertainty to be a conservative estimate. A lower ΔR bound of -400 years implies a marine 14C age that nearly approximates the age of the contemporaneous atmosphere at the time of deposition (i.e., virtually no marine reservoir effect). On the other hand, an upper ΔR bound of +1000 years largely overestimates the established local ΔR value of -30±49 years generally found in the literature (e.g., Bauch et al., 2001). This if further confirmed by precisely dated benthic foraminifera from the Norwegian Sea, which monitor intermediate waters leaving the Nordic Seas and feeding the Arctic Ocean and indicate ΔR values of about 0 years (i.e., close to modern) during the late Younger Dryas stadial and Early Holocene (Muschitiello et al., 2019).

Sediment physical property measurements (bulk density, magnetic susceptibility, grain size, and TOC) performed on 31-PC allow us to firmly establish a lithostratigraphic correlation to a neighboring sediment core (PS2767-4) collected by RV Polarstern during expedition ARK-XI/1 in 1995 (Rachor, 1997). PS2767-4 is underpinned by a published composite chronology that was constructed using cross-correlation and the amalgamation of dating information across a number of regional marine sediment sequences (Figs. 1 and 2) (Stein et al., 2001, 2012).

Core PS2767-4 is located only 24 km south of 31-PC (Fig. 1). This 8.22 m long core is thought to be younger than 60 kyr BP, implying that it contains a complete record of MIS 3 through 1 (Stein et al., 2001, 2012). It has been used to investigate organic matter delivery from the Laptev shelf (Stein et al., 2001; Stein and Fahl, 2000) and more recently for biomarker-based reconstructions of sea ice variability during the last glacial cycle (Stein et al., 2012; Xiao et al., 2015).

The Holocene and postglacial age model for PS2767-4 was dated primarily using 14C radiocarbon dates on three mollusk shells, while the pre-Holocene stratigraphy is based on correlation to other records where oxygen isotopes, magnetostratigraphy, and dinoflagellate biostratigraphy have been used to infer approximate ages (Stein et al., 2001). These older regional age constraints were mapped onto PS2767-4 using a regional correlation between magnetic susceptibility records and sediment lithology (Stein et al., 2001). Stein and Fahl (2012) recognized that in the absence of direct dating, the pre-Holocene stratigraphy of this record remained tentative (Figs. 1 and 2). Examination of the stratigraphic profiles for 31-PC and PS-2767-4 on their independent depth scale reveals an unequivocal coherence across multiple parameters, which argues in favor of a consistent depositional history at the two sites (Fig. 2).

Although the postglacial and Holocene age model for PS-2767-4 is fairly robust, being variably dated by 14C measurements in each of the correlated records, the inferred ages for glacial sediments (MIS 2–4) are considerably less well-constrained (Fig. 2) (Stein et al., 2012). Boundaries between MIS 2, 3, and 4 are demarcated mainly by correlative changes in sediment lithology and organic matter content, and they are dated using a dinocyst-based stratigraphy developed for PS2471-1 (Matthiessen et al., 2001) (Fig. 1). Results from our bulk 14C radiocarbon dates illustrate a clear deviation in the proposed age model for PS-2767-4 beyond ∼ 14 kyr BP. The two near-bottom bulk 14C dates yield an uppermost bound of ∼ 31.8 and ∼ 33.5 kyr BP (Table 1, Fig. 2) for the base of 31-PC. Given that the near-basal depositional age of 31-PC cannot be older than the late MIS 3/early MIS 2, it seems unlikely that sediment from PS2767-4 can date back to 60 kyr BP (i.e., MIS 4), as previously suggested (Stein et al., 2012).

The outcomes of the probabilistic alignment and the resulting age model for 31-PC are presented in Fig. 3. The correlation between ϕ and GISP2 δ18O requires some considerations. Large-scale climate shifts may not be manifested in a similar way in different proxies. In fact, our ϕ record integrates a number of processes, such as sediment grain size, composition, transport, and deposition at the coring site. By contrast, GISP δ18O is responsive to changes in Greenland air temperature. Therefore, the two parameters likely do not scale in a linear fashion, and a perfect match should not be expected. Nonetheless, the degree of correlation between the marine and ice-core data – especially during deglaciation and MIS 2 – is surprisingly well defined and overall compelling.

The results show that 31-PC features relatively linear sedimentation rates with a mean of ∼41±24 (1σ) cm kyr-1, with the exception of the last ∼ 9 kyr where sedimentation rates are slightly lower (∼ 15 cm kyr-1) (Fig. 3). This is in line with the expected decrease in depositional rate resulting from post-glacial transgression of the Siberian shelves (Bauch et al., 2001; Tesi et al., 2016).

The mean 99 % posterior credible interval for our age model is ∼±1.46 kyr (Fig. 3). Mean uncertainties are larger during the Holocene (∼±1.85 kyr). The larger age error reflects on one hand the absence of radiocarbon-based age constraints in the upper 1.4 m and on the other hand the lack of structure in the ϕ and GISP2 δ18O data during the Holocene, which makes the match statistically less robust (Fig. 3).

The implications of these findings are critical for reconstructing and understanding the oceanographic and sea-ice history along the Siberian Arctic shelf. Our new age model also indicates that the entire sedimentary sequence recovered in PS2767-4 was deposited during MIS 2 and the Holocene. Previous studies had speculated that high-frequency variability in biomarkers seen below 450 cm core depth may reflect millennial-scale climate fluctuations during MIS 3 (Fahl and Stein, 2012). The later part of MIS 3 has always been regarded as a time of anomalous warmth in the Arctic, with central Arctic sediments younger than 40–50 kyr usually containing little ice-rafted debris and high numbers of calcareous microfossils (Hanslik et al., 2010; Nørgaard-Pedersen et al., 1998; Poore et al., 1999). However, given the updated chronology, no insights into environmental conditions during MIS 3 can be made from these sedimentary records. What they do provide instead are relatively high-resolution records of environmental conditions during MIS 2 and the deglaciation.

The revised age model reveals substantially faster sedimentation rates during the Last Glacial Maximum (LGM)/MIS 2 and deglaciation compared to the previous chronology (Fig. 4), with a mean sedimentation rates of ∼ 24 cm kyr-1. These are in stark contrast to the LGM hiatus observed in many sediment cores from the western Arctic Ocean, notably on the Mendeleev Ridge, and even a few in the central Arctic on the Lomonosov Ridge (Jakobsson et al., 2014; Poirier et al., 2013; Polyak et al., 2009). This break in sedimentation captured in multiple cores and extending between 13 and 20 14C kyr has been associated with the development of a thick, coherent perennial sea ice cover, possibly even the growth of paleocrystic sea ice or an ice shelf (Polyak et al., 2009). The rapid sedimentation rates found on the southern Lomonosov Ridge off Siberia indicate a profoundly different depositional environment with strong sediment supply, and they are more similar to MIS 2 sedimentation rates reported along the Eurasian continental margin and northern Greenland margins (Jakobsson et al., 2014; Nørgaard-Pedersen et al., 2003).

Direct insights into sea-ice conditions are made by re-evaluating published semi-quantitative biomarker-based sea-ice proxy from PS2767-4 (PBIP25) (Fahl and Stein, 2012). These data were incorporated into spatial biomarker-based reconstructions of Arctic sea ice conditions during the LGM (Xiao et al., 2015) and require a re-assessment given our new chronology. In general, the new chronology indicates for more variable, but more extensive than present day, sea-ice cover during the early LGM/MIS 2 (rather than during the later part of MIS 3). The later part of deglaciation (rather than LGM/MIS 2) featured increasing sea-ice growth with permanent sea ice throughout the year. Decreased sea-ice cover occurred around the transition between deglaciation and the Holocene (rather than during deglaciation) (Fig. 4).

These trends can be further illustrated by cross-plotting the IP25 and brassicasterol data, using the same limits for ice extent defined by Xiao et al. (2015) (Fig. 5). In this scenario, the absence of both IP25 (sea-ice diatom biomarker) and brassicasterol (open-water phytoplankton biomarker) indicates extensive permanent sea-ice or shelf ice conditions, and higher levels of IP25 and brassicasterol indicate more productive ice marginal settings (Müller et al., 2011; Stein et al., 2012). According to our new chronology, there is a clear change from permanent and extended ice cover through most of LGM/MIS 2 and deglaciation towards a more marginal ice-edge setting during the Holocene. Notably, permanent and less scattered sea-ice conditions existed across the deglaciation, which is consistent with the development of thick sea ice postulated for the Younger Dryas stadial (Bradley and England, 2008).