, 4:22 | Cite as

A 190-ka biomarker record revealing interactions between sea ice, Atlantic Water inflow and ice sheet activity in eastern Fram Strait

  • A. Kremer
  • R. Stein
  • K. Fahl
  • H. Bauch
  • A. Mackensen
  • F. Niessen
Original Article
Part of the following topical collections:
  1. PAST Gateways


The northeastern Fram Strait at the entrance to the Arctic Ocean represents a key observatory for sea ice reconstructions as it sensitively reacts to environmental changes. A combined biomarker approach (HBIs, sterols, alkenones) was carried out on Core PS93/006-1 from the western Svalbard margin to reconstruct sea ice conditions related to glacial–interglacial cycles of the last 190 ka. The continuous presence of sea ice demonstrates the strong influence of polar water masses in the eastern Fram Strait. Glacial intervals are characterised by extended sea ice conditions with perennial sea ice cover during early MIS 6, the Penultimate Glacial Maximum, the interstadial MIS 5d, MIS 4 and the Last Glacial Maximum. Less severe, yet highly variable, sea ice conditions with more frequent summer melt dominated the interglacial stages. The opposing sea ice conditions along the western and northern Svalbard margin highlight the different regional impact of various environmental forces in eastern Fram Strait. Thus, the major expansion of the Svalbard Barents Sea Ice Sheet favoured the formation of perennial sea ice west of Svalbard while it triggered the establishment of marginal ice cover on the Yermak Plateau.


Sea ice Ice sheets IP25 Biomarker Arctic ocean Fram strait 


The Arctic realm is a central area in the ongoing debate around climate change as it mirrors the global warming trend strongly amplified [77, 98, 100]. In turn, the Arctic Ocean with its permanent to seasonal sea ice cover plays a critical role in maintaining climatic stability worldwide [130]. Arctic sea ice significantly influences the Earth’s energy budget through albedo feedback mechanisms at the ocean’s surface [41]. The release of brines during the formation of seasonal sea ice contributes to deepwater convection, an important trigger of the thermohaline circulation [20, 30, 94]. The dramatic decline of Arctic sea ice observed in recent decades [75, 112] is, therefore, alarming, particularly as it exceeds the forecasts of common climate models [99]. Consequently, a completely ice-free Arctic Ocean during summer may be attained by the end of this century [86, 111]. The related freshening of the North Atlantic has the potential to disrupt or slow down the global conveyor belt [40, 53, 131]. To fully assess the far-reaching consequences of this trend, the reconstruction of past sea ice variations is key. The sea ice proxy IP25 (C25 HBI [highly branched isoprenoids] monoene = IP25 [[10]) is nowadays indispensable for sea ice reconstructions in the Arctic Ocean. Biosynthesised by specific Arctic sea ice diatoms [11, 21], this molecule is predominantly found in areas covered by seasonal sea ice [10, 129]. By combining IP25 with marine open-water phytoplankton biomarkers, in form of the so-called PIP25 index (phytoplankton marker—IP25 = PIP25; [80]), an even more quantitative estimation of sea ice conditions is possible.

The Fram Strait is a major gateway connecting the Arctic Ocean with the North Atlantic, hence the world oceans (Fig. 1). Herein, the sea ice distribution is decisively influenced by the variable strength of the inflowing Atlantic and the outflowing polar water masses. Therefore, this highly dynamic system represents a key observatory for sea ice reconstructions, as it sensitively reacts to any alteration of environmental boundary conditions. Previous paleoenvironmental studies in this region provided important information on the last glacial to Holocene (middle to late MIS 1) sea ice variability (e.g., [23, 62, 79, 81, 82]). However, details on sea ice conditions over longer timescales are rare (e.g., [59, 60, 106, 109]) but absolutely necessary to close existing knowledge gaps concerning sea ice interactions in the climate system. In this regard, our study focuses on the reconstruction of eastern Fram Strait sea ice variations in the course of glacial–interglacial cycles during the last 190 ka. Measurements of specific biomarkers (HBIs, sterols, alkenones) were carried out on Core PS93/006-1 from the Western continental margin of Svalbard (Fig. 1) to complement the understanding of driving mechanisms behind sea ice variability. A comparison with further sea ice reconstructions from the Arctic Gateway aims to give a more comprehensive picture of sea ice responses related to major environmental changes.

Fig. 1

Overview of the oceanographic setting in the Fram Strait and recent SST conditions (inset). The white arrows refer to the inflowing Atlantic and outflowing Polar Waters. The West Spitsbergen Current (WSC), the Svalbard Branch (SB), the Yermak Branch (YB), the Return Atlantic Current (RAC) and the East Greenland Current (EGC) are indicated in white. The position of the September sea ice margin for the time interval 1979–1983 and for the years 2007, 2012 and 2017 is indicated by yellow lines ( Core locations are marked by diamonds, the herein investigated Core PS93/006-1 is highlighted in red

Reconstruction of Arctic sea ice cover

For the calculation of the PIP25 index, several options are available in the selection of a suitable phytoplankton counterpart to IP25. The most commonly used plankton markers are the sterols brassicasterol (i.e., PBIP25) and dinosterol (i.e., PDIP25) that are produced by a variety of phytoplankton genera like dinoflagellates, diatoms and haptophytes [17, 91, 123]. Both PBIP25 and PDIP25 yield a positive correlation with modern satellite-based sea ice observations [80, 129] and were successfully applied in various paleoenvironmental studies of Quaternary, Pliocene and even late Miocene sedimentary sections (e.g., [12, 49, 50, 61, 81, 82, 108, 109, 110]). However, potential limitations should be considered when using brassicasterol and dinosterol for PIP25-based sea ice reconstructions. As these sterols are biosynthesised by a relatively broad group of marine phytoplankton, their sedimentary signal might represent ambiguous environmental conditions. Furthermore, the structural differences between the IP25 and sterol compounds may affect the PIP25 index due to selective biomarker degradation (for a detailed review of potential limitations see [14, 83, 106]).

A just recently proposed phytoplankton marker for the calculation of the PIP25 index is a tri-unsaturated HBI lipid (HBI III, i.e., PIIIIP25). Its stable isotope signature reveals a polar phytoplankton origin [11, 72], supposedly from marine diatoms that thrive proximal to the winter sea ice edge in the Marginal Ice Zone [9, 12, 13, 92, 102]. In surface sediments from the Barents Sea, PIIIIP25 values provided realistic estimates of spring sea ice concentrations (SpSIC) with a threshold of > 0.8 linked to the presence of summer sea ice (> 5% summer sea ice concentration [SuSIC]; [102]).

Discussed as an additional approach for sea ice reconstructions is the di-unsaturated HBI molecule (HBI II [10]) that is structurally related to IP25. The co-occurrence of HBI II and IP25 in Arctic sediments and its isotopic composition point to a sea ice-related origin [10, 11, 119, 128, 129]. In the Arctic Ocean, previous studies applied the HBI II/IP25 ratio (i.e., DIP25) to determine sea ice variability [22] and SST tendency [35, 128].

Regional setting

As the Arctic’s only deepwater passage, the Fram Strait decisively regulates the exchange of water masses with the North Atlantic [4, 36, 78]. There are two counteractive current systems: the northward-flowing West Spitsbergen Current (WSC) and the southward-flowing East Greenland Current (EGC). Considered as the northernmost continuation of the North Atlantic Current, the WSC carries warm and saline Atlantic Water along the western continental margin of Spitsbergen into the Arctic Ocean [1, 25, 27, 96]. Between 78 and 80°N, the WSC separates into the Svalbard and the Yermak Branch [69, 70]. The Svalbard Branch streams northeasterly, staying close to the continental margin of Spitsbergen and eventually enters the Arctic Ocean [2, 27]. The Yermak Branch initially follows the western flank of the Yermak Plateau, partly detaches from it north of 80°N, turns westward and recirculates southward as the Return Atlantic Current (RAC; [19, 42, 71]). The EGC, on the other hand, flows southward along the eastern coast of Greenland, thereby driving the major outflow of polar freshwater and sea ice from the interior Arctic [3, 29, 40, 93, 97, 101]. The oceanic situation in the Fram Strait is subject to a strong intra- and interannual variability that substantially controls the distribution of sea ice [58, 76, 101, 121]. The entire Svalbard archipelago is enclosed by sea ice during the winter months from November to April. Lowest sea ice concentrations prevail, however, along the western margin due to the continuous inflow of Atlantic Water [120]. The fjord systems and sheltered coastal areas around Spitsbergen remain covered by landfast ice until June [39, 114], with a recent tendency to stay ice free throughout the year [5, 28, 85]. When the minimum sea ice extent is reached in September, open-water conditions dominate in the eastern Fram Strait (National Snow and Ice Data Center,

Materials and methods

Sediment core PS93/006-1 was recovered from the northwestern continental margin of Spitsbergen in ~ 1600 m water depth during RV Polarstern expedition PS93 in 2015 [107; Fig. 1].

The core provides a continuous sedimentary record, mainly composed of greyish to brownish silty clays. The lowermost part of the core (785–600 cm) contains coal fragments. An increased content of dropstones can be observed for the intervals 785–640, 360–300 and 200–65 cm (107). A preliminary shipboard stratigraphy for Core PS93/006-1 suggested an age of MIS 6 to MIS 1. This age model was established by correlating the lithology and physical properties (wet bulk density, magnetic susceptibility) to Core PS2839-4 [107].

Stable oxygen and carbon isotope analyses were carried out on planktic foraminifera (N. pachyderma sin.; 150–250 µm; 5–32 specimen) using a Finnigan MAT 253 isotope ratio mass spectrometer coupled to a carbonate preparation device Kiel IV. Measurements were calibrated against the international NBS-19 standard and reported in δ-notation versus Vienna Peedee Belemnite (VPDB). The long-term precision of δ18O and δ13C, based on an internal laboratory standard (Solnhofen limestone) measured over a 1-year period together with samples, was better than ± 0.08 and ± 0.06‰, respectively.

For bulk organic-geochemical analyses, subsamples were taken in an interval of ~ 5 cm and stored in glass vials at − 20 °C. In the further course, the sediment was freeze-dried, ground and total organic carbon (TOC) contents were determined using a Carbon–Sulfur Analyser (CS-125, Leco).

Biomarker analyses

For highly branched isoprenoids (HBIs) and sterol analyses (every 5–10 cm), 5 g of sediment was extracted with ultrasonication using dichloromethane:methanol (2:1 vol/vol) as solvent. Beforehand, the internal standards 7-hexylnonadecane (7-HND; 0.076 µg/sample) and cholesterol-d6 (cholest-5-en-3β-ol-D6; 10.1 µg/sample) were added for biomarker quantification. Hydrocarbons and sterols were separated via open column chromatography using SiO2 as stationary phase and 5 ml of n-hexane followed by 9 ml of ethylacetate:n-hexane (2:8 vol/vol) as eluent, respectively. Sterols were silylated with 200 µl bis-trimethylsilyl-trifluoroacetamide (BSTFA; 60 °C, 2 h) in the next step. Compound identification was carried out with coupled gas chromatography–mass spectrometry (GC–MS; Agilent Technologies). GC measurements were carried out with the following temperature setup: 60 °C (3 min), 150 °C (heating rate: 15 °C/min), 320 °C (heating rate: 10 °C/min), 320 °C (15 min isothermal) for the hydrocarbons and 60 °C (2 min), 150 °C (heating rate: 15 °C/min), 320 °C (heating rate: 3 °C/min), and 320 °C (20 min isothermal) for the sterols. Helium served as carrier gas (1 ml/min constant flow). Specific compound identification was based on the comparison of retention times and mass spectra with literature references (sterols [17, 122]; HBIs [9, 56]). The concentration of each biomarker was calculated by setting its individual GC–MS ion responses in relation to those of respective internal standards. For the quantification of the sterols (quantified as trimethylsilyl ethers), the molecular ions m/z 470 for brassicasterol (as 24-methylcholesta-5,22E-dien-3β-ol), m/z 472 for campesterol (as 24-methylcholest-5-en-3β-ol), m/z 486 for β-sitosterol (as 24-ethylcholest-5-en-3β-ol) and m/z 500 for dinosterol (as 4α,23,24R-trimethyl-5α-cholest-22E-en-3β-ol) were used in relation to the molecular ion m/z 464 for the internal standard cholesterol-d6. For HBI quantification, their molecular ion (m/z 350 for IP25, m/z 348 for HBI II and m/z 346 for HBI III) was compared to the fragment ion m/z 266 for the internal standard 7-HND. The different responses of these ions were balanced by an external calibration curve (see [35]). All biomarker concentrations were corrected to the amount of extracted sediment and organic carbon content.

PIP25 indices (phytoplankton biomarker combined with IP25) were calculated following the equation of Müller et al. [80]: PIP25 = IP25/(IP25 + (plankton marker × c)), where c is a balance factor to compensate possible concentration differences between IP25 and the plankton marker (c = mean IP25 concentration/mean plankton marker concentrations). Brassicasterol, dinosterol and HBI III were used as phytoplankton markers to compare between PBIP25, PDIP25 and PIIIIP25, respectively.

In addition, long-chain C37 alkenones were quantified in selected samples to evaluate the production of haptophyte algae. For the sake of simplicity, alkenones (C37:2, C37:3, C37:4) were eluted from the sterol fraction (every 10–20 cm). The separation of compounds was performed by open column chromatography using SiO2 as stationary phase and 3 ml of n-hexane:dichloromethane (1:1 vol/vol) followed by 5 ml of dichloromethane as eluent. After saponification with 0.1 N KOH (9:1 vol/vol) at 80 °C for 2 h, GC measurements were carried out using an HP 6890 GC equipped with a cold injection system, a DB-1MS fused silica capillary column and a flame ionisation detector. Individual alkenone identification was based on retention times and the comparison with an external standard.

All data reported in this study are available online on the PANGAEA database (


The records of the stable oxygen and carbon isotopes are discontinuous for certain intervals (780–710, 680–630, 540–500, 440–390, 340–310, 200–150 cm; Fig. 2), owing to an absence of planktic foraminifera. Other than that, values range from 1.3 to 4.6‰ for δ18O and from − 2.8 to 0.7‰ for δ13C.

Fig. 2

TOC content (wt%) and biomarker (IP25, HBI II, HBI III, brassicasterol, dinosterol, campesterol, β-sitosterol) contents (µg/g OC) of Core PS93/006-1 against depth (cm). The black lines illustrate the floating average of three data points. The presence/absence of alkenones is indicated by filled/empty circles

The TOC content shows a succession of intervals with either low contents around 0.5% (780–600, 420–300, 200–140 cm) or high contents around 0.8% (in between; Fig. 2). The only exception of this general pattern is a short-term drop (~ 0.3%) of otherwise high percentages at 500 cm.

For IP25, minor contents around 0.08 µg/g OC can be observed in the lowermost part of the core (780–600 cm; Fig. 2). The overlying section (600–340 cm) holds enhanced IP25 values around 0.25 µg/g OC, except for short-term decreases around 500 and 370 cm. Moderate IP25 contents (~ 0.15 µg/g OC) occur between 340 and 190 cm followed by an interval completely absent of IP25 (190–150 cm). The uppermost part of the core is characterised by relatively high IP25 contents (up to 0.50 µg/g OC) around 110 cm that decrease towards the core top (~ 0.15 µg/g OC; Fig. 2).

The records of the di- and tri-unsaturated HBIs (HBI II and HBI III, respectively) show minimal contents in the lowermost part from 780 to 600 cm (Fig. 2). Above, HBI II is slightly enhanced with pronounced peaks around 460 cm (~ 5.04 µg/g OC) and 30 cm (~ 2.56 µg/g OC). HBI III shows a stronger fluctuating signal with enhanced values for the intervals from 490 to 390 cm (~ 2.90 µg/g OC), 300 to 230 cm (~ 3.67 µg/g OC) and 110 cm to the core top (up to 19.26 µg/g OC). Likewise to the IP25 record, both HBI II and HBI III are absent between 190 and 150 cm (Fig. 2).

The sterols brassicasterol, dinosterol, β-sitosterol and campesterol show comparable trends; however, brassicasterol records a notably higher variability (Fig. 2). Between the base of the core and 150 cm, mean values of 10.5 µg/g OC for brassicasterol, 2.9 µg/g OC for dinosterol, 3.2 µg/g OC for campesterol and 11.9 µg/g OC for β-sitosterol can be observed. Pronounced single peaks occur around 673, 590, 480, 403 and 283 cm for brassicasterol, at 480 cm for campesterol and at 403 cm for dinosterol and β-sitosterol. The uppermost part of the core is characterised by a shift to enhanced values of up to 57.8, 12.7, 22.2 and 67.0 µg/g OC for brassicasterol, dinosterol, campesterol and β-sitosterol, respectively, around 110 cm with a decreasing tendency towards the core top (Fig. 2).

The C37:2 and C37:3 alkenones were abundant throughout the core, except for the samples at 743, 723, 638, 618, 510, 393, 313, 283, 173 and 83 cm core depth (Fig. 2).


Age model

The revised age model of Core PS93/006-1 is based on δ18O stratigraphy complemented by core correlation and biostratigraphy (Fig. 3; Table 1).

Fig. 3

Stratigraphic framework of Core PS93/006-1. The age tie points (yellow asterisks) derive from the correlation of δ18O values to the global LR04 stack [67], the occurrence of the biostratigraphic marker Turborotalita quinqueloba indicating MIS event 5.5 (e.g., [84]) and the correlation of the MS record to the MS stack compiled by Jessen et al. [55] for the western Svalbard margin. An additional verification of this age model is given by the tentative correlation of the MS and the carbonate record to corresponding records of the neighbouring cores PS1535-8 [104], PS1294, PS1295 and PS1297 [46]. The ages of specific MIS events are adopted in accordance to Thompson and Goldstein [116]. Meltwater events are characterised by concurrent deflections of the δ18O and the δ13C records towards lighter values and are indicated by light blue arrows

Table 1

Age fix points of the age model of Core PS93/006-1

Depth (cmbsf)

Age (cal. ka B.P.)

Marine isotope stage/event

Fix point origin



















MIS 3.31




4/3 boundary












5/6 boundary






Origin of the age fix points: (1) correlation of the MS record to the stacked MS record compiled by Jessen et al. [55] for the western Svalbard slope, (2) correlation of the δ18O record to the global LR04 stack [67], (3) occurrence of the biostratigraphic marker Turborotalita quinqueloba for event 5.5 in the polar North Atlantic [84], (4) occurrence of coal fragments, indicative for MIS 6 deposits (cf., [15])

When using δ18O stratigraphy in the Arctic Ocean, local processes like sea ice formation and freshwater discharge might complicate the correlation to the global isotope signal (e.g., [104]). The isotope record of Core PS93/006-1 shows several gaps due to low or missing amounts of planktic foraminifera. However, the identification of several Marine Isotope Stage (MIS) events and boundaries was possible (Fig. 3). The base of the core is marked by relatively light δ18O values that are associated with latest MIS 7 (~ 193 ka). Planktic foraminifera are predominantly absent in the overlying sequence (780–630 cm), probably indicating stage 6 deposits. The concurrent abundance of coal fragments further supports the allocation of MIS 6 to this core interval (Fig. 3) as Bischof et al. [15] identified an enhanced transport of coal fragments from Siberia towards the Fram Strait and Nordic Seas during MIS 6. An abrupt transition from very heavy to light δ18O values around 603 cm is most probably linked to major deglaciation in the course of Termination II. The MIS 6/5 boundary (130 ka) is placed at the midpoint from heavier to lighter values. Two deflections towards lighter oxygen isotope values at 573 and 463 cm are associated with the MIS events 5.5 (~ 123 ka) and 5.33 (~ 104 ka; Thompson and Goldstein, [116]), respectively. The occurrence of the subpolar biostratigraphic marker Turborotalita quinqueloba at 570 cm further indicates MIS event 5.5 (e.g., [84]; Fig. 3). This planktic foraminifer species is regarded to indicate the inflow of Atlantic Water to the Fram Strait [124] and occurs in enhanced abundances during MIS 5e [6, 18, 88, 132]. The gaps in the δ18O record around 515 and 410 cm might represent the substages MIS 5d and 5b, respectively. These substages are associated with colder temperatures and possibly more severe sea ice conditions that are unfavourable for the production of planktic foraminifera [124]. A pronounced shift from heavier to lighter δ18O values marks the MIS 4/3 boundary (57 ka), with its midpoint lying at 298 cm. Above, the MIS 3.31 event (~ 53.8 ka; Thompson and Goldstein [116]) indicates the general shift to lighter oxygen isotope values after deglaciation (Fig. 3).

In the upper part of Core PS93/006-1, the chronostratigraphy is based on the correlation of the magnetic susceptibility (MS) to the MS stack compiled by Jessen et al. [55] for the eastern Fram Strait (Fig. 3). These authors combined the MS records of eleven sediment cores positioned along the western continental margin of Svalbard between 76° and 80°N to develop a chronology tool for the region. Two consecutive lows in the MS stack around 23.8 and 14.5 ka can be correlated to corresponding trends in the MS record of Core PS93/006-1 around 200 and 115 cm core depth, respectively. For the Holocene sequence, Jessen et al. [55] described an initial rise of MS values that peak around 8.4 ka and gradually decreases thereafter towards the core top. A similar trend can be observed in the MS record of Core PS93/006-1, culminating in peak values around 76 cm (Fig. 3).

An additional verification of this age model is given by the tentative correlation of the MS and the carbonate record to corresponding records of neighbouring cores. More precisely, the MS record of Core PS1535-8 [104] and the carbonate records of the cores PS1294, PS1295 and PS1297 [46] show fluctuations comparable to those observed in the records of Core PS93/006-1 (Fig. 3).

According to this age model, Core PS93/006-1 represents late MIS 7 to MIS 1 (Fig. 4). In general, the tentative shipboard stratigraphy by Stein [107] can be supported.

Fig. 4

Age–depth model and sedimentation rates (cm/ky) of Core PS93/006-1. Yellow stars indicate age tie points obtained from stable isotope stratigraphy, biostratigraphy and correlation of the magnetic susceptibility record to the MS stack of Jessen et al. [55]. A linear interpolation is used to calculate ages in between these age tie points. MIS boundaries are indicated by dashed lines, grey shading refers to glacial intervals

Evaluation of sea ice proxies applied in this study

In Core PS93/006-1, all three PIP25 indices yield similar outcomes regarding the prevalence of permanent versus seasonal sea ice cover (Fig. 5). Especially, the PBIP25 and PDIP25 indices resemble each other to a great extent, most likely owing to the related group of source organisms. The PIIIIP25 index shows slightly higher values than the PBIP25 and PDIP25 indices, implicating more extensive sea ice conditions. At the transition from permanent (PIP25 ~ 1; [80]) to seasonal sea ice cover (PIP25 << 1; [80]), the decrease of PIIIIP25 lags slightly behind the PBIP25 and PDIP25 indices (dashed lines in Fig. 5). Moreover, the PBIP25 and PDIP25 indices sharply decrease in response to the changing environment, while the PIIIIP25 index reacts in a less pronounced, more gradual manner due to rather hesitantly rising HBI III contents (Fig. 5). This might be attributable to the higher selectivity of HBI III towards marine genera [9, 92], whereas dino- and brassicasterol are biosynthesised by a broader group of phytoplankton occupying an overall longer bloom season [17, 91, 123]. The slight deviations between brassicasterol and HBI III become further apparent from the correlation of IP25 versus both phytoplankton markers, while the general pattern is similar (Fig. 6a, c). Therefore, the HBI III compound and the applicability of PIIIIP25 need to be verified by further studies including sediment traps and broader surface datasets. Moreover, the subtle differences between the PBIP25 and PIIIIP25 indices are certainly better captured by prospective higher resolution studies.

Fig. 5

Fluctuations of TOC (black), δ18O (red), IP25 (black), brassicasterol (green), dinosterol (blue), HBI II (yellow) and HBI III (grey) against age. Accumulation rates of IP25 are indicated as grey colouration. The presence or absence of alkenones is highlighted by filled or empty circles, respectively. Age tie points are marked by black triangles. Grey shading refers to glacial intervals. PBIP25, PDIP25 and PIIIP25 indices are indicated in green, blue and grey, respectively, with the coloured area highlighting most extensive sea ice conditions (e.g. [80, 102]). The vertical dashed lines indicate the small offset between the PBIP25/PDIP25 indices and the PIIIP25 index at the transition from permanent to seasonal sea ice cover

Fig. 6

IP25 (µg/g OC) versus brassicasterol (µg/g OC) for a Core PS93/006-1 and b Core PS92/039-2, and IP25 (µg/g OC) versus HBI III (µg/g OC) for c Core PS93/006-1 and d Core PS92/039-2. The classification of the different sea ice conditions for the crossplots of IP25 and brassicasterol (i.e., less, variable/seasonal, extended and permanent ice cover) follows Müller et al. [80] and Xiao et al. [129]. Different intervals (PGM, LIG, LGM, Holocene) are highlighted by colouration

Apart from that, the sea ice conditions at the western Svalbard margin are resembled equally well by the different PIP25 indices. This verifies the applicability of the PIIIIP25 index for sea ice reconstructions in eastern Fram Strait and again the use of the PBIP25 and PDIP25 indices in this region (e.g., [12, 109]). Hence, PBIP25 does not seem to suffer from an ambiguous origin of brassicasterol as it can be observed for the Kara and Laptev Sea regions influenced by enhanced river runoff [33, 34, 50].

In general, the partly in-phase fluctuations of IP25 and the phytoplankton markers, likewise described for the eastern Fram Strait in paleorecords (e.g., [24, 81]) and surface sediments [103], do not impede the PIP25 approach. However, the PIP25 indices should always be discussed alongside the individual biomarker profiles to avoid misleading sea ice reconstructions (e.g., [80, 106]).

Regarding the HBI II compound, there is no correlation between IP25 and HBI II in Core PS93/006-1 (Fig. 5). Instead, HBI II correlates positively with the HBI III compound revealing a somehow related origin. Considering the observations of previous studies, the HBI II compound may, therefore, derive from both sea ice and open-water phytoplankton assemblages. The non(sea ice)-specific origin of this compound in Core PS93/006-1 leads to an exclusion of the HBI II signal in the sea ice reconstruction of this study.

Sea ice variability in eastern Fram Strait over the last 190 ka

The cyclic glacial–interglacial climate variability of the last 190 ka strongly impacted the oceanic and related sea ice conditions in the Arctic Gateway [26, 45, 46, 57, 59, 60, 63, 104, 117, 126]. Interestingly, sea ice must have been present in the eastern Fram Strait throughout the entire time interval, as an almost continuous input of the sea ice proxy IP25 is recorded in Core PS93/006-1 (Fig. 5). Apart from that, IP25 and further biomarkers show a strong variability that can be linked to shifts of glacial–interglacial modes. Especially, the transition from peak glacial (i.e., Penultimate Glacial Maximum [PGM], Last Glacial Maximum [LGM]) to peak interglacial stages (i.e., the Last Interglacial [LIG], Holocene) reveals profoundly changing environmental conditions and forces.

Glacial conditions

In the Arctic Ocean, glacial intervals are mainly associated with the build-up of large ice sheets on the continental shelves. Four major Eurasian ice sheets have been formed within the last 200 ka: a most extensive one during the late Saalian (> 140 ka), followed by two early (~ 110 ka, ~ 90 ka), a middle (~ 60 ka) and a late Weichselian glaciation (~ 20 ka; [51, 52, 54, 64, 68, 115]). Along western Svalbard, the Svalbard Barents Sea Ice Sheet (SBIS) advanced up to 40 km off the coast during major glaciations (e.g., [54, 55, 113, 115, 126]). Glacial intervals are further characterised by a reduced (temperate), yet persistent, inflow of Atlantic Water to the Arctic Ocean [32, 48, 59, 132].

The biomarker records of Core PS93/006-1 reveal a prevalence of severe sea ice conditions during glacial intervals at the western continental margin of Svalbard (Fig. 7). Decreased fluxes of the sea ice proxy IP25 and the phytoplankton biomarkers brassicasterol, dinosterol and HBI III indicate suboptimal conditions for both sea ice and open-water algal growth. Accordingly, the PIP25 indices show elevated values implicating a pervasive sea ice cover (Fig. 7). The concurrently diminished contents of organic carbon (Fig. 5) and the terrigenous biomarkers (campesterol, β-sitosterol) indicate the overall hampered particle flux to the sea floor. However, a seasonal ice break-up must have been present most of the time due to the nearly steady, albeit very low, input of biomarkers and the presence of alkenones (Fig. 5). The abundance of dropstones suggests floating icebergs in the eastern Fram Strait and in turn at least temporally open-water conditions. The summer ice edge might have been positioned in the vicinity of the core position, resulting in a relatively short season of ice melt and primary production. These overall severe sea ice conditions are also consistent with previous sea ice reconstructions along the western Svalbard margin (i.e., MSM05/5-712-2, [81]) and the western Yermak Plateau (i.e., PS2837-5, [79, 82, 106]; Fig. 7). They clearly indicate the glacial strength of the EGC that exports sea ice from the interior Arctic Ocean towards the Fram Strait. Cold freshwater is carried along with this flow that possibly fostered local sea ice formation. At the same time, the persistent inflow of Atlantic Water to the Arctic Ocean ceases the formation of a permanent sea ice cover along western Svalbard [32, 44, 48, 59, 132]. It is suggested that these seasonally ice-free regions in the Nordic Seas during the glacial intervals represent an essential moisture source for the final growth of Northern Hemisphere ice sheets [31, 44].

Fig. 7

PBIP25 indices for the cores PS93/006-1, PS92/039-2, PS2138-1 [109] and PS2837-5 [79, 82, 106; all in black], and MSM5/5-712-2 [81; in grey] against age. In the record of Core PS93/006-1, the dashed line around 40 ka displays an alternative estimation of sea ice conditions based on the PDIP25 and PIIIIP25 indices that contrast the PBIP25 in this interval (Fig. 5). The paleoproductivity record for Core PS2138-1 is based on benthic foraminifera and indicated in green [127]. Age tie points are indicated by black triangles. The mean summer insolation at 79°N [65] and the global sea level [105] are illustrated for the last 200 ka. Blue shading refers to glacial stages with an extended Svalbard Barents Sea Ice Sheet (SBIS). Peak interglacial stages (i.e., LIG, Holocene) are indicated by red shading

Indications for perennial sea ice conditions along western Svalbard can only be observed during absolute insolation minima, i.e., early MIS 6, the PGM, the interstadial MIS 5d, partly MIS 4 and the LGM (Fig. 7). This is reflected in the biomarker records of Core PS93/006-1 (and the cores MSM05/5-712-2 and PS2837-5 during the LGM) as concurrently absent (or negligible) input of IP25 and the open-water phytoplankton biomarkers (Fig. 6a, c). The insufficient light and nutrient supply associated with a closed sea ice cover limits the primary production in the Arctic Ocean [87, 95]. The production of planktic foraminifera is strongly affected by these hostile conditions (124), as indicated by gaps in the δ18O curve (Fig. 5). When sea ice is not melting during the summer months, no or little organic matter is released to the water column and transported towards the sea floor (decreased TOC and terrigenous biomarker contents; Fig. 5). Hence, the summer ice edge must have been positioned south of 79°N (or east of ~ 5°E), leaving the core position ice covered year-round (Fig. 8a, c). The predominance of permanent sea ice cover coincides with major advances of the SBIS along western Svalbard (Fig. 7; [54, 68, 115]). The enhanced eastward influence of polar water masses combined with the expansion of the ice sheet onto the outer shelf area likely fostered the formation of landfast ice on the continental margin. Under such heavy ice conditions, the inflowing Atlantic Water would have submerged beneath the cold freshwater at the surface and continued farther north in intermediate depth [43, 93].

Fig. 8

Overview map and schematic illustration of the sea ice conditions west and north of Spitsbergen for different intervals: a the PGM, b the LIG, c the LGM and d the Holocene. Core locations are marked by diamonds, the herein investigated Core PS93/006-1 is highlighted in red. The proposed (based on the biomarker records) positions of the winter and summer ice edge are indicated by dashed and dotted black lines, respectively. Red arrows indicate Atlantic Water entering the Arctic Ocean via the Fram Strait

Quite the opposite scenario can be observed when following the continental margin of the Svalbard Archipelago in a northeastern direction into the interior Arctic Ocean. At the eastern Yermak Plateau (PS92/039-2), simultaneously enhanced accumulation of IP25, the marine phytoplankton (brassicasterol, dinosterol, HBI III) and the terrigenous (campesterol, β-sitosterol) biomarkers point to the presence of marginal sea ice cover during intervals of enhanced ice sheet activity (Figs. 6b, d, 7; [63]). A combination of katabatic winds from the protruded SBIS and upwelling of warm, subsurface Atlantic Water along its shelf break (e.g., [37, 66]) triggered the formation of a coastal polynya along the northern Barents Sea margin with the parallel formation of a stationary ice margin on the eastern Yermak Plateau [63; Fig. 8c]. Such polynya-type conditions have also been proposed from studies of other core sites along the major circum-Arctic ice sheets in MIS 6 [59, 60, 109]. Only during the PGM, the eastern Yermak Plateau experienced a perennial ice cover most probably due to a short-term extension of the ice sheet onto the plateau that merged with the sea ice cover (Figs. 7, 8a). Continuing farther east along the continental margin north of Svalbard (i.e., PS2138-1), biomarker proxy records reveal an extended ice cover with intermittent seasonally open water allowing some phytoplankton and ice diatom production (Figs. 7, 8a; [109, 127]).

Interglacial conditions

The transition from glacial to interglacial stages is accompanied by the retreat of major glaciation on the continental shelves and a strengthened inflow of Atlantic Water. Unconstrained by the presence of extended ice sheets and a low sea level stand, the warm water streams along the western continental margin of Svalbard and intrudes the Arctic Ocean. The strongest inflow of Atlantic Water is supposed for the LIG and the Holocene, while MIS 5c, MIS 5a and MIS 3 are associated with less pronounced advection [48, 59, 73, 74, 104, 127].

In Core PS93/006-1, the onset of interglacial intervals is marked by rising contents of IP25 and the phytoplankton markers (Fig. 5), implying improved conditions for sea ice and open-water algae production. Hence, a reduced sea ice cover with more frequent summer melt probably prevailed during interglacials at the western Svalbard slope at 79°N. The predominance of a seasonal sea ice cover is further displayed in the PIP25 records as a shift from previously maximum to moderate or low values (Fig. 7). Furthermore, the constant presence of alkenones implies a regular production of haptophyte algae during summer (Fig. 5). The establishment of seasonal sea ice cover can further be observed in the nearby cores MSM05/5-712 [81], PS2837-5 [79, 82, 106] and PS2138-1 [109; Fig. 7]. Meanwhile, the more interior parts of the Arctic Ocean remained covered by predominantly permanent sea ice [109] with the summer sea ice boundary positioned southward of Core PS92/039-2 at the eastern Yermak Plateau [63; Figs. 6d, 7, 8].

Although the interglacial stages are likewise characterised by the presence of seasonal sea ice cover along the Svalbard margin, subtle differences within and between individual biomarker records reveal highly dynamic sea ice conditions. This becomes particularly apparent when examining the LIG and the Holocene intervals in more detail, as both phases have been attributed with comparable environmental conditions such as orbital configuration and Atlantic Water inflow [38, 47].

At the western Svalbard margin (i.e., PS93/006-1), the LIG is characterised by high IP25 fluxes, while the phytoplankton sterols and the HBI III compound show only moderate input (Fig. 5). This suggests an enhanced production of sea ice diatoms while the open-water phytoplankton experienced rather disadvantageous living conditions. Therefore, it is likely that the core position experienced a long seasonal sea ice cover with at least partial occurrence of sea ice during summer. The disintegration of the enormous ice sheets covering the continents around the Arctic Ocean during MIS 6 probably decelerated the sea ice retreat during the subsequent LIG. The associated huge freshwater discharge would have fostered sea ice formation and concurrently repressed the influence of Atlantic Water in eastern Fram Strait [118]. The northward advection of warm water masses to the Arctic Ocean is clearly indicated by the occurrence of blue mussels in western Svalbard fjords [68]. However, the inflow of Atlantic Water must have been laterally restricted to the easternmost Fram Strait as recorded by various studies along the western Svalbard margin [6, 7, 8, 89, 90, 132]. The LIG is associated with strong westerly winds that might have induced an enhanced influence of cold, polar water in eastern Fram Strait [16, 89, 90].

In contrast, the seasonal sea ice cover was significantly reduced along the northern Svalbard margin as recorded by lowest PBIP25 values [109] and highest paleoproductivities in Core PS2138-1 [127; Fig. 7]. The restricted lateral expansion of Atlantic Water in eastern Fram Strait may have led to an intensified heat transport towards the Yermak Plateau.

Unlike the LIG, the early Holocene shows simultaneously high contents of IP25, the phytoplankton biomarkers and the terrigenous sterols in Core PS93/006-1 (Fig. 5). This is likely related to stable (spring) ice edge conditions above the core position with partly ice-free summers. Two short-term meltwater events following the disintegration of the SBIS might have caused slight delays of the sea ice retreat along western Svalbard [55]. These are clearly visible as deflections in the δ18O and the δ13C records (Fig. 3) towards lighter values indicating the influence of isotopically lighter meltwater and reduced surface water ventilation [104].

In the further course of the Holocene, however, IP25 and the phytoplankton sterols decrease to moderate values while HBI III steadily rises (Fig. 5). Hence, sea ice likely retreated during this period resulting in shorter seasonal ice cover and gradually longer summer seasons. Maximum HBI III contents might indicate the presence of the winter ice edge proximal to the core position of PS93/006-1 (Fig. 8d [12, 102]). This interval probably represents the early Holocene Climate Optimum (~ 8 ka), associated with less severe sea ice cover in the eastern Fram Strait [81, 125]. Due to the insufficient age control of this part of the record, it is, however, difficult to allocate this with certainty. Either way, the divergent interglacial sea ice conditions in the Arctic Gateway demonstrate the highly complex oceanic system of this region.


Our multi-biomarker study from the western Svalbard margin complements the overall picture of past sea ice distribution in eastern Fram Strait. Moreover, it provides valuable information about the complex relation between sea ice cover and various environmental forces (i.e., ice sheet activity, Atlantic Water inflow) throughout glacial–interglacial cycles during the last 190 ka.

The following statements can be made:

  • The continuous presence of sea ice at the western Svalbard margin indicates the strong influence of polar water masses in eastern Fram Strait throughout glacial and interglacial intervals.

  • During glacial intervals, the western Svalbard margin experienced extended sea ice conditions with ephemeral ice break-up during summer due to the persistent, yet temperate, inflow of Atlantic Water.

  • Perennial sea ice cover prevailed during absolute insolation minima with extensive expansion of the Svalbard Barents Sea Ice Sheet (early MIS 6, the Penultimate Glacial Maximum, MIS 5d, MIS 4 and the Last Glacial Maximum).

  • The opposing sea ice variations north (i.e., PS92/039-2) and west (i.e., PS93/006-1) of Svalbard highlight the diverse impact of ice sheet activity in the region. While the expansion of the Svalbard Barents Sea Ice Sheet triggered the formation of perennial sea ice west of Svalbard, it led to the establishment of marginal ice cover north of Svalbard.

  • A reduced sea ice cover with more frequent summer melt dominated the interglacial intervals along the western Svalbard margin.

  • Peak interglacials, i.e., the LIG and the Holocene interval, differ with regard to the intensity of seasonal sea ice conditions due to a variable inflow of Atlantic Water to the Arctic Ocean.

  • As the sea ice conditions in the eastern Fram Strait are resembled equally well by the different PIP25 indices (i.e., PBIP25, PDIP25, PIIIIP25), the applicability of each index for sea ice reconstructions in this region can be verified.



We thank the captain and the crew of R/V Polarstern for excellent cooperation during the cruise PS93 in 2015. Thanks to Walter Luttmer and Lisa Schönborn for technical support during the laboratory work. Thanks to Simon Belt and colleagues (Biogeochemistry Research Centre, University of Plymouth) for providing the internal standard for the IP25 analyses. The paper is a contribution to the German–Chinese project with the title “Natural variability of Arctic sea ice and its significance for global climate change and organic carbon cycle”. Financial support was given by the Federal Ministry of Education and Research (BMBF, Project no. 01DO14004). The authors would like to thank the editor and two anonymous reviewers for their thorough comments that helped to improve the manuscript.

Compliance with ethical standards

Conflict of interest

On behalf of all authors, the corresponding author states that there is no conflict of interest.


  1. 1.
    Aagaard K (1982) Inflow from the Atlantic Ocean to the Polar Basin. In: Rey L (ed) The Arctic Ocean. Comité Arctique International, Monaco, pp 69–82Google Scholar
  2. 2.
    Aagaard K, Foldvik A, Hillman SR (1987) The West Spitsbergen Current—disposition and water mass transformation. J Geophys Res Oceans 92:3778–3784Google Scholar
  3. 3.
    Aagaard K, Coachman L (1968) The East Greenland Current north of Denmark Strait: part II. Arctic 21:181–200Google Scholar
  4. 4.
    Aagaard K, Greisman P (1975) Toward new mass and heat budgets for the Arctic Ocean. J Geophys Res 80:3821–3827Google Scholar
  5. 5.
    Alkire MB, Morison J, Andersen R (2015) Variability in the meteoric water, sea-ice melt, and Pacific water contributions to the central Arctic Ocean, 2000–2014. J Geophys Res Oceans 120:1573–1598Google Scholar
  6. 6.
    Bauch HA, Erlenkeuser H, Fahl K, Spielhagen RF, Weinelt MS, Andruleit H, Henrich R (1999) Evidence for a steeper Eemian than Holocene sea surface temperature gradient between Arctic and sub-Arctic regions. Palaeogeogr Palaeoclimatol Palaeoecol 145, 95–117Google Scholar
  7. 7.
    Bauch HA, Kandiano ES, Helmke JP (2012) Contrasting ocean changes between the subpolar and polar North Atlantic during the past 135 ka. Geophys Res Lett 39:L11604Google Scholar
  8. 8.
    Bauch HA, Erlenkeuser H (2008) A “critical” climatic evaluation of Last Interglacial (MIS 5e) records from the Norwegian Sea. Polar Res 27:135–151Google Scholar
  9. 9.
    Belt ST, Allard WG, Massé G, Robert JM, Rowland SJ (2000) Highly branched isoprenoids (HBIs): Identification of the most common and abundant sedimentary isomers. Geochim Cosmochim Acta 64:3839–3851Google Scholar
  10. 10.
    Belt ST, Massé G, Rowland SJ, Poulin M, Michel C, LeBlanc B (2007) A novel chemical fossil of palaeo sea ice: IP25. Org Geochem 38:16–27Google Scholar
  11. 11.
    Belt ST, Massé G, Vare LL, Rowland SJ, Poulin M, Sicre M-A, Sampei M, Fortier L (2008) Distinctive 13C isotopic signature distinguishes a novel sea ice biomarker in Arctic sediments and sediment traps. Mar Chem 112:158–167Google Scholar
  12. 12.
    Belt ST, Cabedo-Sanz P, Smik L, Navarro-Rodriguez A, Berben SMP, Knies J, Husum K (2015) Identification of paleo Arctic winter sea ice limits and the marginal ice zone: optimised biomarker-based reconstructions of late Quaternary Arctic sea ice. Earth Planet Sci Lett 431:127–139Google Scholar
  13. 13.
    Belt ST, Brown TA, Smik L, Tatarek A, Wiktor J, Stowasser G, Assmy P, Allen CS, Husum K (2017) Identification of C25 highly branched isoprenoid (HBI) alkenes in diatoms of the genus Rhizosolenia in polar and sub-polar marine phytoplankton. Org Geochem 110:65–72Google Scholar
  14. 14.
    Belt ST, Müller J (2013) The Arctic sea ice biomarker IP25: a review of current understanding, recommendations for future research and applications in palaeo sea ice reconstructions. Quat Sci Rev 79:9–25Google Scholar
  15. 15.
    Bischof J, Koch J, Kubisch M, Spielhagen RF, Thiede J (1990) Nordic seas surface ice drift reconstructions: Evidence from ice-rafted coal fragments during oxygen isotope stage 6. In: Dowdeswell JA, Scourse JD (eds). Glacimarine environments: processes and sediments (vol 53, pp 235–251). Geological Society of Special Publication, LondonGoogle Scholar
  16. 16.
    Blindheim JV, Borovkov B, Hansen SA, Maimberg WR, Turrell, Osterbus S (2000) Upper layer cooling and freshening in the Norwegian Sea in relation to atmospheric forcing. Deep Sea Res 47:655–680Google Scholar
  17. 17.
    Boon JJ, Rijpstra WIC, Delange F, Deleeuw JW, Yoshioka M, Shimizu Y (1979) Black sea sterol—molecular fossil for dinoflagellate blooms. Nature 277:125–127Google Scholar
  18. 18.
    Born A, Nisancioglu KH, Risebrobakken B (2011) Late Eemian warming in the Nordic Seas as seen in proxy data and climate models. Paleoceanogr Paleoclimatol 26:PA2207Google Scholar
  19. 19.
    Bourke R, Weigel A, Paquette R (1988) The westward turning branch of the West Spitsbergen Current. J Geophys Res Oceans 93:14065–14077Google Scholar
  20. 20.
    Broecker WS (1997) Thermohaline circulation, the Achilles heel of our climate system: Will man-made CO2 upset the current balance? Science 278:1582–1588Google Scholar
  21. 21.
    Brown T, Belt S, Tatarek A, Mundy C (2014) Source identification of the Arctic sea ice proxy IP25. Nat Commun 5:4197Google Scholar
  22. 22.
    Cabedo-Sanz P, Belt ST, Knies J (2013) Identification of contrasting seasonal sea ice conditions during the Younger Dryas. Quat Sci Rev 79:74–86Google Scholar
  23. 23.
    Cabedo-Sanz P, Belt ST, Jennings AE, Andrews JT, Geirsdóttir Á (2016) Variability in drift ice export from the Arctic Ocean to the North Icelandic Shelf over the last 8000 years: a multi-proxy evaluation. Quat Sci Rev 146:99–115Google Scholar
  24. 24.
    Cabedo-Sanz P, Belt ST (2016) Seasonal sea ice variability in eastern Fram Strait over the last 2000 years. Arktos 2:1–12Google Scholar
  25. 25.
    Carmack E, Polyakov I, Padman L, Fer I, Hunke E, Hutchings J, Jackson J, Kelley D, Kwok R, Layton C, Melling H, Perovich D, Persson O, Ruddick B, Timmermanns M-L, Toole J, Ross T, Vavrus S, Winsor P (2015) Towards quantifying the increasing role of oceanic heat in sea ice loss in the new Arctic. Bull Am Meteor Soc 96(12):2079–2105Google Scholar
  26. 26.
    Chauhan T, Rasmussen TL, Noormets R (2016) Palaeoceanography of the Barents Sea continental margin, north of Nordaustlandet, Svalbard, during the last 74 ka. Boreas 45:76–99Google Scholar
  27. 27.
    Coachman LK, Aagaard K (1974) Physical Oceanography of Arctic and Subarctic Seas. In: Herman Y (ed) Marine geology and oceanography of the Arctic Seas. Springer, Berlin, pp 1–72Google Scholar
  28. 28.
    Cottier FR, Nilsen F, Inall ME, Gerland S, Tverberg V, Svendsen H (2007) Wintertime warming of an Arctic shelf in response to large-scale atmospheric circulation. Geophys Res Lett 34:L10607Google Scholar
  29. 29.
    De Steur L, Hansen E, Mauritzen C, Beszczynska-Möller A, Fahrbach E (2014) Impact of recirculation on the East Greenland Current in Fram Strait: results from moored current meter measurements between 1997 and 2009. Deep Sea Res I 92:26–40Google Scholar
  30. 30.
    Dieckmann GS, Hellmer HH (2008) The Importance of Sea Ice: An Overview. In: Thomas DN, Diekmann GS (eds) Sea Ice: an introduction to its physics, chemistry, biology, and geology. Blackwell Science, Oxford, pp 1–21Google Scholar
  31. 31.
    Dokken TM, Hald M (1996) Rapid climatic shifts during isotope stages 2–4 in the Polar North Atlantic. Geology 24:599–602Google Scholar
  32. 32.
    Ezat MM, Rasmussen TL, Groeneveld J (2014) Persistent intermediate water warming during cold stadials in the southeastern Nordic seas during the past 65 k.y. Geology 42:663–666Google Scholar
  33. 33.
    Fahl K, Stein R, Gaye-Haake B, Gebhardt C, Kodina LA, Unger D, Ittekkot V (2003) Biomarkers in surface sediments from the Ob and Yenisei estuaries and southern Kara Sea: evidence for particulate organic carbon sources, pathways, and degradation. In: Stein R, Fahl K, Fütterer DK, Galimov EM, Stepanets OV (eds) Siberian river run-off in the Kara Sea: characterisation, quantification, variability, and environmental significance. Proceedings in Marine Sciences, vol 6. Elsevier, Amsterdam, pp 329–348Google Scholar
  34. 34.
    Fahl K, Stein R (1999) Biomarkers as organic-carbon-source and environmental indicators in the Late Quaternary Arctic Ocean: ‘‘Problems and perspectives’’. Mar Chem 63:293–309Google Scholar
  35. 35.
    Fahl K, Stein R (2012) Modern seasonal variability and deglacial/Holocene change of central Arctic Ocean sea-ice cover: new insights from biomarker proxy records. Earth Planet Sci Lett 351:123–133Google Scholar
  36. 36.
    Fahrbach E, Meincke J, Osterhus S, Rohardt G, Schauer U, Tverberg V, Verduin J (2001) Direct measurements of volume transports through Fram Strait. Polar Res 20:17–224Google Scholar
  37. 37.
    Falk-Petersen S, Pavlov V, Berge J, Cottier F, Kovacs KM, Lydersen C (2015) At the rainbow’s end: high productivity fueled by winter upwelling along an Arctic shelf. Polar Biol 38:5–11Google Scholar
  38. 38.
    Fischer N, Jungclaus JH (2010) Effects of orbital forcing on atmosphere and ocean heat transports in Holocene and Eemian climate simulations with a comprehensive Earth system model. Clim Past 6:155–168Google Scholar
  39. 39.
    Gerland S, Renner AHH (2007) Sea ice mass balance in an Arctic fjord. Ann Glaciol 46:435–442Google Scholar
  40. 40.
    Haine TW, Curry B, Gerdes R, Hansen E, Karcher M, Lee C, Rudels B, Spreen G, de Steur L, Stewart KD, Woodgate R (2015) Arctic freshwater export: status, mechanisms, and prospects. Glob Planet Change 125:13–35Google Scholar
  41. 41.
    Hall A (2004) The role of surface albedo feedback in climate. J Clim 17:1550–1568Google Scholar
  42. 42.
    Hattermann T, Isachsen PE, von Appen W-J, Albretsen J, Sundfjord A (2016) Where eddies drive recirculation of Atlantic Water in Fram Strait. Geophys Res Lett 7:3406–3414Google Scholar
  43. 43.
    Haugan PM (1999) Structure and heat content of the West Spitsbergen Current. Polar Res 18(2):183–188Google Scholar
  44. 44.
    Hebbeln D, Dokken T, Andersen ES, Hald M, Elverhøi A (1994) Moisture supply for northern ice-sheet growth during the Last Glacial Maximum. Nature 370:357–359Google Scholar
  45. 45.
    Hebbeln D, Henrich R, Baumann K-H (1998) Paleoceanography of the last interglacial/glacial cycle in the Polar North Atlantic. Quat Sci Rev 17:125–153Google Scholar
  46. 46.
    Hebbeln D, Wefer G (1997) Late Quaternary paleoceanography in the Fram Strait. Paleoceanography 12:65–78Google Scholar
  47. 47.
    Helmke JP, Bauch HA (2003) Comparison of glacial and interglacial conditions between the polar and subpolar North Atlantic region over the last five climatic cycles. Paleoceanography. CrossRefGoogle Scholar
  48. 48.
    Henrich R (1998) Dynamics of Atlantic water advection to the Norwegian-Greenland Sea—time-slice record of carbonate distribution in the last 300 ky. Mar Geol 145:95–131Google Scholar
  49. 49.
    Hoff U, Rasmussen TL, Stein R, Ezat MM, Fahl K (2016) Sea ice and millennial-scale climate variability in the Nordic seas 90 kyr ago to present. Nature Communications 7:12247Google Scholar
  50. 50.
    Hörner T, Stein R, Fahl K, Birgel D (2016) Post-glacial variability of sea ice cover, river run-off and biological production in the western Laptev Sea (Arctic Ocean)—a high-resolution biomarker study. Quat Sci Rev 143:133–149Google Scholar
  51. 51.
    Hughes ALC, Gyllencreutz R, Lohne ØS, Mangerud J, Svendsen JI (2016) The last Eurasian ice sheets—a chronological database and time-slice reconstruction, DATED-1. Boreas 45:1–45Google Scholar
  52. 52.
    Ingólfsson Ó, Landvik JY (2013) The Svalbard–Barents Sea ice-sheet—historical, current and future perspectives. Quat Sci Rev 64:33–60Google Scholar
  53. 53.
    Ionita M, Scholz P, Lohmann G, Dima M, Prange M (2016) Linkages between atmospheric blocking, sea ice export through Fram Strait and the Atlantic Meridional Overturning Circulation. Sci Rep 6:32881Google Scholar
  54. 54.
    Jakobsson M, Andreassen K, Bjarnadóttir LR, Dove D, Dowdeswell JA, England JH, Funder S, Kelly Hogan K, Ingólfsson Ó, Jennings AE, Larsen NK, Kirchner N, Landvik JY, Mayer LA, Mikkelsen N, Möller P, Niessen F, Nilsson J, O’Regan M, Polyak L, Nørgaard- Pedersen N, Stein R (2014) Arctic Ocean glacial history. Quat Sci Rev 92:40–67Google Scholar
  55. 55.
    Jessen SP, Rasmussen TL, Nielsen T, Solheim A (2010) A new Late Weichselian and Holocene marine chronology for the western Svalbard slope 30,000–0 cal years BP. Quat Sci Rev 29:1301–1312Google Scholar
  56. 56.
    Johns L, Wraige EJ, Belt ST, Lewis CA, Masseu G, Robert J-M, Rowland SJ (1999) Identification of C25 highly branched isoprenoid (HBI) dienes in Antarctic sediments, sea- ice diatoms and laboratory cultures of diatoms. Org Geochem 30:1471–1475Google Scholar
  57. 57.
    Kandiano ES, Bauch HA, Müller A (2004) Sea surface temperature variability in the North Atlantic during the last two glacial-interglacial cycles: comparison of faunal, oxygen isotopic and Mg/Ca-derived records. Palaeogeogr Palaeoclimatol Palaeoecol 204:145–164Google Scholar
  58. 58.
    Kawasaki T, Hasumi H (2016) The inflow of Atlantic water at the Fram Strait and its interannual variability. J Geophys Res Oceans 121:502–519Google Scholar
  59. 59.
    Knies J, Vogt C, Stein R (1999) Late Quaternary growth and decay of the Svalbard/Barents Sea ice sheet and paleoceanographic evolution in the adjacent Arctic Ocean. Geo Mar Lett 18:195–202Google Scholar
  60. 60.
    Knies J, Nowaczyk N, Müller C, Vogt C, Stein R (2000) A multiproxy approach to reconstruct the environmental changes along the Eurasian continental margin over the last 150 000 years. Mar Geol 163:317–344Google Scholar
  61. 61.
    Knies J, Cabedo-Sanz P, Belt ST (2014) The emergence of modern sea ice cover in the Arctic Ocean. Nat Commun 5:5608Google Scholar
  62. 62.
    Kolling HM, Stein R, Fahl K, Perner K, Moros M (2017) Short-term variability in late Holocene sea ice cover on the East Greenland Shelf and its driving mechanisms. Palaeogeogr Palaeoclimatol Palaeoecol 485(1):336–350Google Scholar
  63. 63.
    Kremer A, Stein R, Fahl K, Ji Z, Yang Z, Wiers S, Matthiessen J, Forwick M, Löwemark L, O’Regan M, Chen J, Snowball I (2018) Changes in sea ice cover and ice sheet extent at the Yermak Plateau during the last 160 ka—reconstructions from biomarker records. Quat Sci Rev (in review) Google Scholar
  64. 64.
    Landvik JY, Bondevik S, Elverhøi A, Fjeldskaar W, Mangerud J, Salvigsen O, Siegert MJ, Svendsen JI, Vorren TO (1998) The last glacial maximum of Svalbard and the Barents Sea area: Ice sheet extent and configuration. Quat Sci Rev 17:43–75Google Scholar
  65. 65.
    Laskar J, Robutel P, Joutel F, Gastineau M, Correia A, Levrard B (2004) A long-term numerical solution for the insolation quantities of the Earth. Astron Astrophys 428:261–285Google Scholar
  66. 66.
    Lind S, Ingvaldsen R (2012) Variability and impacts of Atlantic Water entering the Barents Sea from the north. Deep Sea Res 62(I):70–88Google Scholar
  67. 67.
    Lisiecki LE, Raymo ME (2005) A Pliocene–Pleistocene stack of 57 globally distributed benthic δ18O records. Paleoceanography 20:PA1003Google Scholar
  68. 68.
    Mangerud J, Dokken T, Hebbeln D, Heggen B, Ingólfsson O, Landvik JY, Mejdahl V, Svendsen JI, Vorren TO (1998) Fluctuations of the Svalbard–Barents Sea Ice Sheet during the last 150 000 years. Quat Sci Rev 17:11–42Google Scholar
  69. 69.
    Manley T, Bourke R, Hunkins K (1992) Near-surface circulation over the Yermak Plateau in northern Fram Strait. J Mar Syst 3:107–125Google Scholar
  70. 70.
    Manley T (1995) Branching of Atlantic Water within the Greenland-Spitsbergen Passage: an estimate of recirculation. J Geophys Res Oceans 100:20627–20634Google Scholar
  71. 71.
    Marnela M, Rudels B, Houssais MN, Beszczynska-Möller A, Eriksson PB (2013) Recirculation in the Fram Strait and transports of water in and north of the Fram Strait derived from CTD data. Ocean Sci 9(3):499–519Google Scholar
  72. 72.
    Massé G, Belt ST, Crosta X, Schmidt S, Snape I, Thomas DN, Rowland SJ (2011) Highly branched isoprenoids as proxies for variable sea ice conditions in the Southern Ocean. Antarct Sci 23:487–498Google Scholar
  73. 73.
    Matthiessen J, Knies J, Nowaczyk NR, Stein R (2001) Late Quaternary dinoflagellate cyst stratigraphy at the Eurasian continental margin, Arctic Ocean: indications for Atlantic water inflow in the past 150,000 years. Glob Planet Change 31:65–86Google Scholar
  74. 74.
    Matthiessen J, Knies J (2001) Dinoflagellate cyst evidence for warm interglacial conditions at the northern Barents Sea margin, during marine isotope stage 5. J Quat Sci 16:727–737Google Scholar
  75. 75.
    Meier WN, Hovelsrud GK, van Oort BEH, Key JR, Kovacs KM, Michel C, Haas C, Granskog MA, Gerland S, Perovich DK, Makshtas A, Reist JD (2014) Arctic sea ice in transformation: a review of recent observed changes and impacts on biology and human activity. Rev Geophys 52:185–217. CrossRefGoogle Scholar
  76. 76.
    Meyer A, Sundfjord A, Fer I, Provost C, Villacieros-Robineau N, Koenig Z, Onarheim IH, Smedsrud LH, Duarte P, Dodd PA, Graham RM, Schmidtko S, Kauko HM (2017) Winter to summer oceanographic observations in the Arctic Ocean north of Svalbard. J Geophys Res Oceans 122(8):6218–6237Google Scholar
  77. 77.
    Miller GH, Alley RB, Brigham-Grette J, Fitzpatrick JJ, Polyak L, Serreze M, White JWC (2010) Arctic Amplification: can the past constrain the future? Quat Sci Rev 29:1779–1790Google Scholar
  78. 78.
    Mosby H (1962) Water, salt and heat balance of the North Polar Sea and the Norwegian Sea. Geophys Publ 24 (11):289–313Google Scholar
  79. 79.
    Müller J, Massé G, Stein R, Belt ST (2009) Variability of sea-ice conditions in the Fram Strait over the past 30,000 years. Nat Geosci 2:772–776Google Scholar
  80. 80.
    Müller J, Wagner A, Fahl K, Stein R, Prange M, Lohmann G (2011) Towards quantitative sea ice reconstructions in the northern North Atlantic: a combined biomarker and numerical modelling approach. Earth Planet Sci Lett 306:137–148Google Scholar
  81. 81.
    Müller J, Werner K, Stein R, Fahl K, Moros M, Jansen E (2012) Holocene cooling culminates in sea ice oscillations in Fram Strait. Quat Sci Rev 47:1–14Google Scholar
  82. 82.
    Müller J, Stein R (2014) High-resolution record of late glacial and deglacial sea ice changes in Fram Strait corroborates ice–ocean interactions during abrupt climate shifts. Earth Planet Sci Lett 403:446–455Google Scholar
  83. 83.
    Navarro-Rodriguez A, Belt ST, Knies J, Brown TA (2013) Mapping recent sea ice conditions in the Barents Sea using the proxy for palaeo sea ice reconstructions. Quat Sci Rev 79:26–39Google Scholar
  84. 84.
    Nørgaard-Pedersen N, Mikkelsen N, Lassen SJ, Kristoffersen Y, Sheldon E (2007) Reduced sea ice concentrations in the Arctic Ocean during the last interglacial period revealed by sediment cores off Northern Greenland. Paleoceanography 22:PA1218Google Scholar
  85. 85.
    Onarheim I, Smedsrud L, Ingvaldsen R, Nilsen F (2014) Loss of sea ice during winter north of Svalbard. Tellus A 66:23933Google Scholar
  86. 86.
    Overland JE, Wang M (2013) When will the summer arctic be nearly sea ice free? Geophys Res Lett 40:2097–2101Google Scholar
  87. 87.
    Popova EE, Yool A, Coward AC, Dupont F, Deal C, Elliott S, Hunke E, Jin M, Steele M, Zhang J (2012) What controls primary production in the Arctic Ocean? Results from an intercomparison of five general circulation models with biogeochemistry. J Geophys Res 117:C00D12Google Scholar
  88. 88.
    Rasmussen TL, Thomsen E, Troelstra SR, Kuijpers A, Prins MA (2003) Millennial-scale glacial variability versus Holocene stability: Changes in planktic and benthic foraminifera faunas and ocean circulation in the North Atlantic during the last 60,000 years. Mar Micropaleontol 47:143–176Google Scholar
  89. 89.
    Risebrobakken B, Dokken T, Jansen E (2005) Extent and variability of the Meridional Atlantic Circulation in the Eastern Nordic Seas during Marine Isotope Stage 5 and its influence on the inception of the Last Glacial. The Nordic Seas: an integrated perspective. merican Geophysical Union, Geophysical Monograph Series, Washington DC, pp 323–339Google Scholar
  90. 90.
    Risebrobakken B, Dokken T, Otterå OH, Jansen E, Gao Y, Drange H (2007) Inception of the Northern European ice sheet due to contrasting ocean and insolation forcing. Quat Res 67:128–135Google Scholar
  91. 91.
    Robinson N, Eglinton G, Brassell SC, Cranwell PA (1984) Dinoflagellate origin for sedimentary 4α-methylsteroids and 5α(H)-stanols. Nature 308:439–442Google Scholar
  92. 92.
    Rowland SJ, Allard WG, Belta ST, Massae G, Robert JM, Blackburn S, Frampton D, Revill AT, Volkman JK (2001) Factors influencing the distributions of polyunsaturated terpenoids in the diatom, Rhizosolenia setigera. Phytochemistry 58:717–728Google Scholar
  93. 93.
    Rudels B, Friedrich HJ, Quadfasel D (1999) The Arctic circumpolar boundary current. Deep Sea Res Part II 46:1023–1062Google Scholar
  94. 94.
    Rudels B, Korhonen M, Schauer U, Pisarev S, Rabe B, Wisotzki A (2015) Circulation and transformation of Atlantic water in the Eurasian Basin and the contribution of the Fram Strait inflow branch to the Arctic Ocean heat budget. Prog Oceanogr 132:128–152Google Scholar
  95. 95.
    Sakshaug E (2004) Primary and secondary production in the Arctic Seas. In: Stein R, Macdonald RW (eds) The organic carbon cycle in the Arctic Ocean. Springer, Berlin, pp 57–82Google Scholar
  96. 96.
    Schauer U, Fahrbach E, Osterhus S, Roharardt R (2004) Arctic warming through the Fram Strait: oceanic heat transport from 3 years of measurements. J Geophys Res 109:C06026Google Scholar
  97. 97.
    Schlichtholz P, Houssais M-N (1999) An inverse modeling study in Fram Strait. Part I: dynamics and circulation. Deep Sea Res Part II 46:1083–1135Google Scholar
  98. 98.
    Screen JA, Simmonds I (2010) The central role of diminishing sea ice in recent Arctic temperature amplification. Nature 464:1334–1337Google Scholar
  99. 99.
    Serreze MC, Barry RG (2011) Processes and impacts of Arctic amplification: a research synthesis. Glob Planet Change 77:85–96Google Scholar
  100. 100.
    Serreze MC, Francis JA (2006) The arctic amplification debate. Clim Change 76:241–264Google Scholar
  101. 101.
    Smedsrud LH, Halvorsen MH, Stroeve JC, Zhang R, Kloster K (2017) Fram Strait sea ice export variability and September Arctic sea ice extent over the last 80 years. Cryosphere 11:65–79Google Scholar
  102. 102.
    Smik L, Cabedo-Sanz P, Belt ST (2016) Semi-quantitative estimates of paleo Arctic sea ice concentration based on source-specific highly branched isoprenoid alkenes: a further development of the PIP25 index. Org Geochem 92:63–69Google Scholar
  103. 103.
    Smik L, Belt ST (2017) Distributions of the Arctic sea ice biomarker proxy IP25 and two phytoplanktonic biomarkers in surface sediments from West Svalbard. Org Geochem 105:39–41Google Scholar
  104. 104.
    Spielhagen RF, Baumann KH, Erlenkeuser H, Nowaczyk NR, Norgaard-Pedersen N, Vogt C, Weiel D (2004) Arctic Ocean deep-sea record of northern Eurasian ice sheet history. Quat Sci Rev 23:1455–1483Google Scholar
  105. 105.
    Spratt RM, Lisiecki LE (2016) A Late Pleistocene sea level stack. Clim Past 12:1079–1092Google Scholar
  106. 106.
    Stein R, Fahl K, Müller J (2012) Proxy reconstruction of Arctic Ocean sea ice history: from IRD to IP25. Polarforschung 82:37–71Google Scholar
  107. 107.
    Stein R (2016) The Expedition PS93.1 of the Research Vessel POLARSTERN to the Greenland Sea and the Fram Strait in 2015. Reports on polar and marine research, Bremerhaven, Alfred Wegener Institute for Polar and Marine ResearchGoogle Scholar
  108. 108.
    Stein R, Fahl K, Schreck M, Knorr G, Niessen F, Forwick M, Gebhardt C, Jensen L, Kaminski M, Kopf A, Matthiessen J, Jokat W, Lohmann G (2016) Evidence for ice-free summers in the late Miocene central Arctic Ocean. Nat Commun 7:1–13Google Scholar
  109. 109.
    Stein R, Fahl K, Gierz P, Niessen F, Lohmann G (2017) Arctic Ocean sea ice cover during the penultimate glacial and the last interglacial. Nat Commun 8:373Google Scholar
  110. 110.
    Stein R, Fahl K (2013) Biomarker proxy shows potential for studying the entire Quaternary Arctic sea ice history. Org Geochem 55:98–102Google Scholar
  111. 111.
    Stocker TF, Qin, Plattner G-K, Tignor M, Allen SK, Boschung J, Nauels A, Xia Y, Bex V, Midgley PM (2013) Climate change 2013: the physical science basis. Contribution of Working Group I to the Fifth Assessment Report of the Intergovern- mental Panel on Climate Change (IPCC). Cambridge University Press, CambridgeGoogle Scholar
  112. 112.
    Stroeve JC, Serreze MC, Holland MM (2012) The Arctic’s rapidly shrinking sea ice cover: a research synthesis. Clim Change 110:1005–1027Google Scholar
  113. 113.
    Svendsen JI, Mangerud J, Elverhoi A, Solheim A, Schuttenhelm RTE (1992) The late Weichselian glacial maximum on western Spitsbergen inferred from offshore sediment cores. Mar Geol 104:1–17Google Scholar
  114. 114.
    Svendsen H, Beszczynska-Möller A, Hagen JO, Lefauconnier B, Tverberg V, Gerland S, Bischof K, Papucci C, Zajaczkowski M, Azzolini R, Bruland O, Wiencke C, Winther JG, Dallmann W (2002) The physical environment of Kongsfjorden-Krossfjorden, an Arctic fjord system in Svalbard. Polar Res 21(1):133–166Google Scholar
  115. 115.
    Svendsen JI, Alexanderson H, Astakhov VI, Demidov I, Dowdeswell JA, Funder S, Gataullin V, Henriksen M, Hjort C, Houmark-Nielsen M, Hubberten HW, Ingólfsson O, Jakobsson M, Kjaer KH, Larsen E, Lokrantz H, Lunkka JP, Lysa A, Mangerud J, Matiouchkov A, Murray A, Moller P, Niessen F, Nikolskaya O, Polyak L, Saarnisto M, Siegert C, Siegert MJ, Spielhagen RF, Stein R (2004) Late Quaternary ice sheet history of northern Eurasia. Quat Sci Rev 23:1229–1271Google Scholar
  116. 116.
    Thompson WG, Goldstein SL (2006) A radiometric calibration of the SPECMAP timescale. Quat Sci Rev 25:3207–3215Google Scholar
  117. 117.
    Tütken T, Eisenhauer A, Wiegand B, Hansen BT (2002) Glacial–interglacial cycles in Sr and Nd isotopic composition of Arctic marine sediments triggered by the Svalbard/Barents Sea ice sheet. Mar Geol 182:351–372Google Scholar
  118. 118.
    Van Nieuwenhove N, Bauch H, Eynaud F, Kandiano E, Cortijo E, Turon J-L (2011) Evidence for delayed poleward expansion of North Atlantic surface waters during the last interglacial (MIS 5e). Quat Sci Rev 30:934–946Google Scholar
  119. 119.
    Vare LL, Masé G, Gregory TR (2009) Sea ice variations in the central Canadian Arctic Archipelago during the Holocene. Quat Sci Rev 28:1354–1366Google Scholar
  120. 120.
    Vinje T (1985) Sea ice distribution 1971-80. Norsk Polarinstutt SkrifterGoogle Scholar
  121. 121.
    Vinje T (2001) Anomalies and trends of sea-ice extent and atmospheric circulation in the Nordic Seas during the period 1864–1998. J Clim 14:255–267Google Scholar
  122. 122.
    Volkman JK (1986) A review of sterol markers for marine and terrigenous organic-matter. Org Geochem 9:83–99Google Scholar
  123. 123.
    Volkman JK, Barrett SM, Blackburn SI, Mansour MP, Sikes EL, Gelin F (1998) Microalgal biomarkers: a review of recent research developments. Org Geochem 29:1163–1179Google Scholar
  124. 124.
    Volkmann R (2000) Planktic foraminifers in the outer Laptev Sea and the Fram Strait—modern distribution and ecology. J Foramin Res 30:157–176Google Scholar
  125. 125.
    Werner K, Spielhagen RF, Bauch D et al (2013) Atlantic Water advection versus sea ice advances in the eastern Fram Strait during the last 9 ka: multi proxy evidence for a two-phase Holocene. Paleoceanography 28:283–295Google Scholar
  126. 126.
    Winkelmann D, Schafer C, Stein R, Mackensen A (2008) Terrigenous events and climate history of the Sophia Basin, Arctic Ocean. Geochem Geophys Geosyst 9:Q07023Google Scholar
  127. 127.
    Wollenburg JE, Kuhnt W, Mackensen A (2001) Changes in Arctic Ocean paleoproductivity and hydrography during the last 145 kyr: the benthic foraminiferal record. Paleoceanography 16:65–77Google Scholar
  128. 128.
    Xiao X, Fahl K, Stein R (2013) Biomarker distributions in surface sediments from the Kara and Laptev seas (Arctic Ocean): indicators for organic-carbon sources and sea ice coverage. Quat Sci Rev 79:40–52Google Scholar
  129. 129.
    Xiao XT, Fahl K, Müller J, Stein R (2015) Sea-ice distribution in the modern Arctic Ocean: Biomarker records from trans-Arctic Ocean surface sediments. Geochim Cosmochim Acta 155:16–29Google Scholar
  130. 130.
    Yang S, Christensen JH (2012) Arctic sea ice reduction and European cold winters in CMIP5 climate change experiments. Geophys Res Lett 39:L20707Google Scholar
  131. 131.
    Yang H, Lohmann G, Wei W, Dima M, Ionita M, Liu J (2016) Intensification and poleward shift of subtropical western boundary currents in a warming climate. J Geophys Res Oceans 121:4928–4945Google Scholar
  132. 132.
    Zhuravleva A, Bauch HA, Spielhagen RF (2017) Atlantic water transfer through the Arctic Gateway (Fram Strait) during the Last Interglacial. Global Planet Change 157:232–243Google Scholar

Copyright information

© Springer-Verlag GmbH Germany, part of Springer Nature 2018

Authors and Affiliations

  • A. Kremer
    • 1
  • R. Stein
    • 1
  • K. Fahl
    • 1
  • H. Bauch
    • 1
  • A. Mackensen
    • 1
  • F. Niessen
    • 1
  1. 1.Alfred-Wegener-Institut Helmholtz-Zentrum fur Polar- und MeeresforschungBremerhavenGermany

Personalised recommendations