During the SWERUS-C3 expedition in 2014 on the Swedish icebreaker Oden, an eight-meter-long piston core (31-PC) was obtained at 1120 m depth on the Southern Lomonosov Ridge (79.91° N, 143.23° E, Fig. 1). Due to disturbance of the surface sediments during the piston coring processes, surface sediment, representing the top of the core was used from multicore 31-MUC, sampled at the same location. The chronology of core 31-PC was established through a combination of bulk radiocarbon dating and stratigraphic correlation with Greenland ice-core records, resulting in an age-model covering the last 27 kyr (Muschitiello et al., 2020; Martens et al., 2020). The age-model was developed using a Markov Chain Monte Carlo algorithm to probabilistically align variations in sediment porosity with δ18O records from Greenland ice cores, with eight bulk radiocarbon ages providing maximum age constraints. The age-model development is described in detail in Muschitiello et al. (2020). Sedimentation rates in 31-PC were on average 41 ± 24 cm kyr⁻¹, apart from the past 9 kyr where they were lower (15 cm kyr⁻¹; Muschitiello et al., 2020). The sampling location on the intersection between the Siberian continental margin and the Southern Lomonosov Ridge enabled sea ice reconstructions encompassing parts of the East Siberian- and Laptev Sea. Additionally, the sampling location enabled quantification of fluvially transported inland PF-C from the Lena River throughout the last deglaciation, including the Holocene. Notably, the core site was adjacent to ICD PF on the continental shelf, which formed during low sea levels and exposed deposits that are now submerged as subsea permafrost. This makes the core site an important receptor for terrOC from past shelf flooding and coastal erosion of the continental margin.

The biomarker analyses were carried out following the method initiated by Belt et al. (2012), with minor modifications. Briefly, freeze-dried sediment samples were homogenised and ground to fine powders, followed by additions of 20 µL of 7-hexylnonadecane (0.00465 µg µL⁻¹), 9-octylheptadec-8-ene (0.00535 µg µL⁻¹), and androstanol (0.45 µg µL⁻¹), respectively, as internal standards to correct for analyte losses during the laboratory procedure. Lipid extraction was performed on the sediments with saponification using KOH (5 %, MeOH / H₂O: 9/1, v/v; KOH ≥ 85 % purity; MeOH ≥ 99.9 % purity), heating for 1 h at 70 °C. The saponified compounds were then extracted by repeated liquid-liquid extractions with hexane after vortexing and centrifugating the samples. Condensed extracts were eluted into a non-polar (using 6 mL n-hexane) and a polar fraction (using 6 mL DCM / MeOH, 1:1, v/v) through open-column chromatography with 1 % H₂O deactivated silica gel (63–200 µm; Supelco) as the stationary phase.

Before instrumental analyses, the polar fraction was derivatized with N, O-Bis (trimethylsilyl) trifluoroacetamide (BSTFA) at 70 °C for 1 h. Both polar and non-polar fractions were analysed using gas chromatography-mass spectrometry (GC-MS) (GC-MS 7820-A, Agilent Technologies USA). Non-polar compounds were analysed in splitless mode using single ion monitoring mode (SIM), starting at a temperature of 40 °C, ramping up 10 °C min⁻¹ up to a maximum temperature of 300 °C for 10 min. To minimize potential carryover, a one-minute post-run at 305 °C was conducted for each sample. The polar fraction was analysed in splitless mode using total ion current mode (TIC), with a starting temperature of 60 °C held for one minute, followed by a 10 °C per minute temperature ramp to a maximum temperature of 310 °C, maintained for 16 min.

Identification of IP₂₅ was based on the retention times of the compound in reference sediments provided by Professor Simon Belt (University of Plymouth). To ensure accurate identification of IP₂₅, the reference sediment and subsamples of 31-PC were first analyzed in full scan mode before SIM analysis. The target sterol compounds were identified based on the mass fragmentation pattern, and monitored through selective ion chromatograms, m/z 348, 470, 472, 486, 500 (androstanol, brassicasterol, campesterol, β-sitosterol, and dinosterol). Quantifications of IP₂₅ and sterols were based on the differences in instrumental mass spectral responses of the target compounds and the internal standards, following the method of Belt et al., (2012). The response factors were determined using external calibration curves of HBI reference sediment, and external standards of the target sterols (Sigma-Aldrich). Additionally, the PIP₂₅-indices (Eqs. 1 and 2) were calculated according to Müller et al. (2011). 1PIP25=IP25IP25+P×c2c=meanIP25mean[P] where P is a phytoplankton biomarker (e.g., brassicasterol or dinosterol), and c is the concentration balance factor to account for concentration differences in IP₂₅ and phytoplankton biomarkers (Müller et al., 2011). Biomarker fluxes were determined by multiplying analyte concentrations with the dry bulk density and sedimentation rate in each slice of the sediment core following the approach of Martens et al. (2020). The limit of detection (LOD) for each biomarker was experimentally determined for each biomarker through external calibration curves of reference sediments (IP₂₅) and sterols (external standards) according to Eq. 3. 3LOD=3.3×SDxy/b where SD_xy represents the residual standard deviation of the calibration curve and b is the slope of the calibration curve.

The stable carbon isotope ratio (δ13C-OC) in the surface sediment of 31-MUC was determined using an elemental analyser coupled to an isotopic ratio mass spectrometer (EA-irMS) following the procedure of Martens et al. (2020). The radiocarbon (Δ14C) content of OC in 31-MUC was analysed through accelerator mass spectrometry at the Tandem Laboratory, Uppsala University. The other carbon isotopic composition and radiocarbon ages of OC in 31-PC were already published and re-used from Martens et al. (2020). However, new pre-depositional ages of the OC were employed in this study, now with the ages further corrected for atmospheric Δ14C changes, according to Intcal20 (Reimer et al., 2020). The radiocarbon of OC at the time of deposition (Δ14C_initial) was calculated through Eq. (4) (Schefuß et al., 2016). 4Δ14Cinitial=(F14Ceλt-1)×1000‰ where F¹⁴C is the fraction modern carbon, t is the time of deposition, and λ the decay constant for radiocarbon. Furthermore, the pre-depositional age (years) was subsequently corrected through Eq. (5) (Schefuß et al., 2016). 5Pre-depositional age=-8033×ln⁡1+(Δ14Cinitial10001+Δ14Catm1000

The Dansgaard-Oeschger (DO) events are characterized by rapid increases in atmospheric and sea surface temperatures, along with significant changes in hydrology (Dansgaard et al., 1993; Jensen et al., 2018). Given the age model uncertainty, the DO-3 event (28–27.5 kyr B.P.), or the aftermath of it, is likely captured at the base of 31-PC (spanning 27.2 ± 2.3/-1.4 kyr B.P., 90 % confidence interval; Muschitiello et al., 2020; Fig. 3). The relatively low PIP₂₅ during this time suggest that the DO-3 likely led to approximately ∼ 1 kyr of seasonal sea ice cover, persisting until the transition into the LGM at 26 kyr B.P. This contrasts with earlier findings from the Southern Lomonosov Ridge (PS2767-4) that reported increased sea ice preceding the LGM (Xiao et al., 2015). However, this earlier study was based on a low-resolution core using a less well-constrained age model that was subsequently re-interpreted based on the improved dating of 31-PC (Muschitiello et al., 2020). Conclusively attributing seasonal sea ice cover during/after DO-3 to increased temperatures is still challenging, given that the onset of DO-3 is not captured in this record. Nevertheless, the amplified phytoplankton productivity following DO-3 was likely associated with seasonal sea ice cover and/or marginal ice zones. Nutrient input from melting sea ice may have contributed to the enhanced IP₂₅, dinosterol, and brassicasterol (Fig. 2: Belt, 2018), and the increased input of marine OC to the core location (Fig. 5; Martens et al., 2020). However, nutrient supply from terrestrial sources could also be an important contributor to the observation of increased productivity. As DO-3 progressed into its later stages, the increasing PIP₂₅ values indicate an expanding sea ice cover approaching the LGM (Figs. 2 and 4).

Succeeding the DO-3, 31-PC recorded an abrupt increase in terrestrial biomarkers, consistent with the high input of OC observed in core 31-PC (Figs. 4 and 5; Martens et al., 2020). The terrestrial OC is characterized by young pre-depositional ages of ∼ 4–7 kyrs (Fig. 5). Accordingly, previous research attributed the significant input of terrestrial OC to active layer thawing, caused by warming during DO-3 (Martens et al., 2020). However, our records of reduced sea ice may suggest additional mechanisms causing active layer thawing, associated with heat and moisture transport during times of open water conditions. Specifically due to the proximity of the core location to the paleo-shoreline at the time (Fig. 1). The reduced sea ice after DO-3 may have led to higher air temperatures, which, combined with increased moisture sources, could result in heavy autumn snowfall. This snow could insulate the permafrost from low winter temperatures, leading to further thawing, as was displayed by Vaks et al. (2020). Consequently, the more frequent open water conditions observed in core 31-PC following DO-3, persisting for about 1 kyr or longer, may have contributed to the remobilization of permafrost carbon and observed enhancement of terrestrial OC input. Taken together, core 31-PC represents the first high-resolution record of sea ice history following DO-3 in this region, providing valuable insights into the sea ice impact on permafrost carbon remobilization.

During the cold stadial of LGM and HS-1 (Heinrich Stadial-1), the concentrations of IP₂₅ and marine phytoplankton biomarkers were below detection limits (Fig. 2), indicating the presence of permanent sea ice. The absence of these biomarkers suggests that thick sea ice inhibited the growth of open-water phytoplankton, but also hindered ice-algae by limiting light penetration through the ice. In contrast to studies that artificially set the PIP₂₅ index to 1 in the absence of both sea ice and marine phytoplankton biomarkers, our records emphasize the absence of these biomarkers to highlight the presence of permanent sea ice (Fig. 2). The onset of perennial sea ice cover during the LGM is consistent with recently published proxy records from the central Arctic Ocean (Pavia et al., 2025). However, these results contrast with studies from the Laptev Sea, which suggest the existence of seasonal sea ice or even ice-free conditions during the HS-1 (Hörner et al., 2016; Fahl and Stein, 2012). Thus, the different sea ice conditions between the Southern Lomonosov Ridge and the nearby Laptev Sea may be attributed to the influence of Atlantic Water (Bradley and England, 2008), but it should be noted that the differences in sea ice cover can also be explained by the presence of polynyas in the Laptev Sea (Taldenkova et al., 2013). It is possible that this warm current has reduced sea ice in the Laptev Sea, while the Southern Lomonosov Ridge and the adjacent East Siberian Sea were less affected by this warm inflow and maintained permanent sea ice (Xiao et al., 2015). This could also explain the near absence of pre-Holocene sediment records in the East Siberian Sea. However, further studies are necessary to validate this hypothesis.

The input of the terrestrial biomarkers campesterol and β-sitosterol was significantly reduced during the LGM and HS-1 (Fig. 2). Previous studies have attributed this decline to the stabilization of permafrost under the cold climate (Martens et al. 2020). After 26 kyr B.P., the decreasing pre-depositional age and low ICD-OC flux suggest that permanent sea ice may have acted as a barrier for transport of terrestrial matter to the study site. It is also worth noting that the presence of permanent sea ice likely reduced the fetch distance and wave activity, thereby protecting coastal areas from erosion and minimizing the input of older carbon sources such as those from ICD-PF. While limited coastal erosion was evident at the Southern Lomonosov Ridge site, a seasonal sea ice cover in the Western Laptev Sea resulted in a stronger input of terrestrial organic matter (Hörner et al., 2016) compared to 31-PC. During the HS-1, 31-PC was rather close to the coastline (75 km), indicating that the core should provide a record of coastal erosion of the continental margin and terrestrial export of organic matter. However, low fluxes of OC and terrestrial biomarkers were present in 31-PC (Figs. 2 and 5). In the Western Laptev Sea, β-sitosterol and campesterol concentrations were around four times higher during the HS-1 (Hörner et al., 2016) compared to 31-PC. This highlights the role of perennial sea ice in limiting terrestrial carbon export during the HS-1 in 31-PC. Additionally, the proximity to the coastline of both cores captures how seasonal sea ice in the Western Laptev Sea permitted higher terrOC input through coastal erosion. Conversely, perennial sea ice cover over the Southern Lomonosov Ridge limited erosion of the nearby continental margin.

During the LGM, the reduction of wave-induced coastal erosion is apparent despite the proximity to the ICD-coastline (Fig. 1). Interestingly, our record showed a minor increase in terrestrial biomarkers and brassicasterol between ∼ 19–22 kyr B.P., while such an increase was absent in dinosterol (Fig. 2). Given that brassicasterol may have lacustrine sources (Hörner et al., 2016), this increase may indicate a moderate rise in material from riverine input. However, the underlying cause of this increase remains unclear, especially considering the cold climate, presence of permanent sea ice (Fig. 2), and relatively stable sea levels (Lambeck et al., 2014).

During the YD stadial, and despite the warming climate of the B/A, the IP₂₅ and open-water algae biomarkers remained undetectable, comparable to the levels observed during the LGM (Fig. 2). This suggests that permanent sea ice persisted from the LGM until approximately 11 kyr BP. This sea ice condition contrasts with the patterns observed in the Kara Sea, Laptev Sea, and the Beaufort Sea, where the sea ice cover was reduced during the B/A and slightly increased during the YD, but generally remained seasonal (Fig. 4; Hörner et al., 2016, 2018; Fahl and Stein, 2012; Wu et al., 2020). Lin et al. (2025) proposed that during this time, Atlantic Water inflow inhibited sea ice formation in the Laptev Sea. Similar to the conditions during the LGM and HS-1. Therefore, the persisting sea ice cover in the Southern Lomonosov Ridge and Eastern Laptev Sea is possibly due to the less profound impact of warm Atlantic Water inflow to our study area.

While permanent sea ice persisted during the B/A, there was a modest increase in terrestrial input, coupled with a notable increase in the pre-depositional age of OC, reaching approximately 15 kyr (Figs. 2 and 5). This implies the involvement of additional mechanisms that remobilized older permafrost carbon. Typically, permanent sea ice is expected to reduce wave activity, thereby limiting coastal erosion, and to decrease heat and moisture transport, which reduces inland permafrost thawing (Barnhart et al., 2014; Vaks et al., 2020). Thus, the observed increase in permafrost carbon remobilization is likely linked to the shelf flooding induced by sea-level rise during the meltwater pulse 1A (MWP-1A; Lambeck et al., 2014; Martens et al., 2020). Indeed, profound sea level rise was observed in the nearby Laptev Sea during MWP-1A (Bauch et al., 2001). This indicates that permafrost carbon remobilization during the B/A was primarily influenced by sea-level rise and atmospheric warming, rather than by sea ice conditions at our study location. It is important to note that the magnitude of permafrost remobilization recorded in core 31-PC was moderate compared to the DO-3 and PB events, as well as compared to the Sea of Okhotsk, Bering Sea, North Pacific, and Beaufort Sea, where extensive remobilization during the B/A has been widely documented (Winterfeld, et al., 2018; Meyer et al., 2019; Wu et al., 2022). This suggests a significant role of the permanent sea ice limiting coastal erosion during B/A of the continental margin close to the Southern Lomonosov Ridge. Indeed, the OC-fluxes were much larger in 31-PC when seasonal sea ice was present after DO-3 and the PB (Figs. 2 and 5). Additionally, nearby cores in the Eastern (PS2458; Fahl and Stein, 2012) and Western Laptev Sea (PS51-154/11; Hörner et al., 2016) display open water/seasonal sea ice conditions concurrent to input of pre-aged OC/n-alkanoic acids, but in several orders of magnitude higher fluxes than in 31-PC (Fig. 5; Nicolas et al., 2025). This suggests that the presence of perennial sea ice at the Southern Lomonosov Ridge, similar to the HS-1, served as a barrier to limit the transport of terrestrial derived material to the core location, and that coastal erosion of the continental margin was limited close to the site. In contrast, regional variability was evident in the Laptev Sea, where seasonal sea ice permitted substantially greater transport of terrestrial organic carbon, partly facilitated by the sea ice conditions permitting more coastal erosion (Fahl and Stein, 2012; Hörner et al., 2016). Granted that permanent sea ice likely limited the transport of terrOC to the core site, the actual magnitude of terrOC input driven by sea-level rise in this region remains uncertain and warrants further investigation.

During the YD stadial, terrestrial input returned to low levels, comparable to the marine contributions as indicated by the marine phytoplankton biomarkers and carbon source apportionment (Fig. 2; Martens et al., 2020). Overall, the low terrestrial input and young pre-depositional age suggest that the cold climate of the YD likely contributed to the stabilization of permafrost, and the permanent sea ice may have had an additional effect on this stabilization.

During the PB/EH transition, there was a marked increase in IP₂₅ and marine phytoplankton biomarkers, with PIP₂₅ values indicating the presence of seasonal sea ice (Fig. 3). Although the emergence of seasonal sea ice aligns with the warmer climate of this period (Shakun et al., 2012), large-scale Arctic sea ice records suggest that the elevated temperatures were not drivers of the breakup of permanent sea ice. This is evident from the observations that sea ice in regions like the Beaufort Sea and Laptev Sea was reduced during the cold deglaciation and expanded during the warm Early Holocene (Fig. 4; Wu et al., 2020; Fahl and Stein, 2012; Hörner et al., 2016). Instead, the transition was likely driven by summer solar insolation at 70° N reached its maximum during this time (Laskar et al., 2004; Stranne et al., 2014). Comparisons with sea ice records from the nearby Laptev Sea reveal contrasting patterns during the early Holocene: while sea ice over the Southern Lomonosov Ridge transitioned from permanent to seasonal sea ice, sea ice coverage in the Laptev Sea increased (Fig. 4; Hörner et al., 2016; Lin et al., 2025). This suggests that the solar insolation, rather than warm Atlantic Water, became the dominant factor influencing sea ice dynamics during this period.

During the PB/EH, there was a substantial increase in terrestrial biomarkers of campesterol and β-sitosterol, even surpassing levels observed during the DO-3 (Fig. 2). Source apportionment of OC suggest that most of the OC in 31-PC came from ICD at the time (Fig. 5; Martens et al., 2020). Due to cross-shelf transport of 0.6 to 2.6 kyr of the OC (Martens et al., 2020), the mobilization could be linked to coastal erosion driven by sea-level rise during the MWP-1B (Lambeck et al., 2014). This finding is consistent with massive permafrost carbon remobilization recorded in the Laptev Sea, Chukchi Sea, Beaufort Sea, Okhotsk Sea, Bering Sea, and North Pacific (Tesi et al., 2016; Martens et al., 2019; Wu et al., 2022; Winterfeld et al., 2018; Meyer et al., 2019; Lin et al., 2025). However, a role of concurrent sea ice reduction cannot be excluded. Seasonal sea ice can promote wave-induced coastal erosion of ICD by increasing fetch distance under frequent open water conditions (Barnhart et al., 2014). Additionally, open water conditions can further elevate heat and moisture transport, potentially leading to deeper thaws of permafrost (Vaks et al., 2020). Therefore, it is plausible that sea-level rise and sea ice reduction acted synergistically to increase destabilization of coastal permafrost, although their relative contributions remain unclear.

Following the rapid decline in sea ice cover during the PB/EH transition, the Mid-Late Holocene is characterized by relatively stable environmental conditions. Starting around 7.5 kyr B.P., phytoplankton and sea ice algae productivity gradually diminished, indicating a buildup of a more extended sea ice cover (Fig. 3). The trend of expanding Arctic sea ice towards the Late Holocene, as observed in our record, is consistent with findings from the Beaufort Sea, Kara Sea, Chukchi Sea, Laptev Sea, and East Siberian Sea (Fig. 5; Fahl and Stein, 2012; Hörner et al., 2016, 2018; Stein et al., 2017; Wu et al., 2020; Dong et al., 2022). Dong et al. (2022) found that the primary driver for mid-Holocene sea ice reduction in the East Siberian Sea was increased solar insolation, driving heat discharge from Arctic rivers. In accordance, the sea ice cover in our record as well as Dong et al. (2022) expanded in the Late Holocene, attributed to reduced summer insolation and resulting heat discharge from nearby rivers.

Coinciding with the expanding sea ice cover, the input of campesterol and β-sitosterol gradually decreased throughout the Holocene, implying a decline in terrOC contribution in tandem with reduced marine productivity (Figs. 2 and 4). This trend of minimal terrOC input and OC deposition is consistently observed in the Laptev Sea and East Siberian Sea (Tesi et al., 2016; Keskitalo et al., 2017; Martens et al., 2020). Although the OC input remained low during the Holocene, the relative influence of marine sources became more important towards the Late Holocene (Fahl and Stein, 2012; Martens et al., 2020). The terrOC remobilization of permafrost pools reduced, likely due to the relatively stable climatic conditions that did not trigger significant OC remobilization, as had occurred during earlier periods of rapid temperature rise.