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Paleo-sea ice distribution and polynya variability on the Kara Sea shelf during the last 12 ka

  • Tanja Hörner
  • Ruediger Stein
  • Kirsten Fahl
Original Article
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The Kara Sea is an important area for paleo-climatic research since sea ice and brine formation take place on its shelf—two processes inducing supra-regional climatic implications and thereby connecting regional environmental variability with global climatic conditions. To gain information about past sea ice coverage and variations, three sediment cores distributed in the southern and central parts of the marginal Sea were investigated. By applying the sea ice biomarker IP25 and the PIP25 index [phytoplankton biomarker (dinosterol)-IP25 index] post-glacial sea ice variability could be detected in the central Kara Sea (Core BP00-36/4), with most intense sea ice cover between 12.4 and 11.8 ka coinciding with the Younger Dryas (12.9–11.6 ka), and reduced sea ice cover between 10 and 8 ka during the Holocene Thermal Maximum. During the last ~ 7 ka, increasing sea ice indicators might indicate a Holocene cooling trend, probably induced by declining summer insolation. Furthermore, temporal changes in the fast ice—polynya distribution in the southern Kara Sea were detected: expanding fast ice during the late Holocene and a cyclic short-term Holocene climate variability documented by abrupt changes in the sea ice coverage at the BP00-07/7 core site. Core BP99-04/7 from the Yenisei estuary recorded consistently seasonal sea ice cover since ~ 9.3 ka, apart from five short phases of fast ice expansion to the core site. The strong influence of river run-off as well as estuary processes might prevent the detection of (short-term) climatic signals at this study site.


Sea ice Arctic Ocean Kara Sea Biomarker Holocene Polynya 


Environmental background and scientific relevance

As sea ice acts as a key component (amplifier of climate change) in terms of the current global warming [1, 2, 3, 4], it is essential to understand its behavior on environmental and climatic changes. During the last decades, Arctic sea ice has retreated enormously with unexpected extreme reductions in the years 2007 and 2012 [5, 6, 7, 8, 9, 10, 11, 12; National Snow and Ice data center,]. As recently reported, cargo ships are able to pass through the Arctic without icebreakers that were usually necessary to open up the way.

Arctic sea ice is mainly produced shallow Siberian shelves where rivers drain high amounts of freshwater to the adjacent ocean. The low-saline nutrient-rich water entry plays an important role for the Arctic oceanic circulation that affects the global circulation system. In the Kara Sea, the two major rivers Ob and Yenisei (with input rates of 404 and 620 km3/a, respectively) are responsible for 40% of the total Arctic river discharge [13]. Moreover, the freshwater input is significant for sea ice formation. Huge amounts of riverine water create a freshwater layer at the water surface, causing a stratification of the Kara Sea water column that is favorable for sea ice, and thus brine formation (Dmitrenko [14]; Eicken et al. [15]). This process is significant for deepwater formation along the continental margin that propels the global circulation system [16].

The freshwater input also controls the salinity distribution and the sedimentation processes on the shelf. Close to the estuary, the so-called Marginal Filter (Fig. 1) catches more than 90% of the suspended matter within this area due to estuary flocculation and aggregation processes [17]. This results in the deposition of a thick Holocene sediment sequence on the Kara Sea shelf [18, 19, 20]. The Kara Sea environment was strongly affected by the post-glacial eustatic sea level rise [21] (Fig. 1b, c), for instance, resulting in the migration of the main depocenter [18, 19] and changing nutrient supply. The shelf became flooded since 16 ka and modern conditions were reached at ~ 7.5 ka (in the eastern Kara Sea) [19, 22].

Fig. 1

a Overview map of the Arctic Ocean with positions of the studied sediment cores and the bathymetric map of the Kara Sea (IBCAO grid, Jakobsson et al. [23]) with averaged monthly sea ice distribution (1978–2010) ( and the location of the Marginal Filter [17]. b, c Show the sea level at ~ 9 and ~ 11 ka [24]

Sea ice distribution is complex on the Kara Sea shelf. Seasonal sea ice achieves its maximum extent in March and the minimum extent in September (Fig. 1 a). Fast ice is formed around the coast and islands (Fig. 2b) in the eastern Kara Sea (Volkov et al. [25]). It builds up in autumn, increases during January and March and lasts until July. Its formation is mainly controlled by winter surface temperatures (Fig. 2c) and winds. Within or very close to the estuary, river run-off also has an effect on the formation process [26]. In autumn, strong and cold offshore winds separate the fast ice from seasonal (drift) ice leading to ice-free areas (polynyas) where an important heat exchange takes place that contributes to deepwater formation [27, 28, 29, 30]. New sea ice is formed at the northern border of the polynya [26]. During the formation process, high amounts of whirled sediment is incorporated within the ice. Both, (drift) ice and sediments are transported from the Kara Sea shelf via the Transpolar Drift (TPD) across the central Arctic to the Fram Strait [31, 32].

Fig. 2

a Position of leads (averaged January-April from 2003 to 2015, maximum in 2012 and minimum in 2003) according to Willmes and Heinemann [33]. b Average fast ice distribution map in March, June and September (1953–1990) taken from Divine et al. [26]. c Monthly average surface air temperature in March, June and September (1951–2015) ( The studied sediment cores are marked in all maps by green-black dots

As described above, on the shallow Siberian shelves substantial as well as climate-relevant environmental processes proceed, e.g., sea ice formation, riverine freshwater input, biological production and brine formation. Modifications of these environmental parameters might affect the TPD and further global Thermohaline Circulation (THC). Increased discharge by Siberian rivers might shift the Drift farther on the Eurasian side to include the increased freshwater and ice fluxes. But probably a greater influence is given by variations in sea ice production and in size and position of the polynyas off Siberia during the Holocene [29]. It is suggested that variations in sea ice and river run-off (freshwater flux) alter the North Atlantic circulation that in turn affects directly the climate in Europe and North America [34, 35].

Given the past variability in the TPD system during the Holocene (e.g., Dyke et al. [36], Funder et al. [37] and Darby et al. [38]) and the current climate change coupled with the alarming sea ice loss, it is important to get information about potential future scenarios, as ecologic but also economic consequences might be severe.

Polynyas are also to major parts responsible for sea ice production and are important spots for detecting regional climate variability since their distribution is affected by climate changes [39].

Nowadays, polynyas are located in front of land fast ice in the Kara and Laptev Sea at the ~ 20–30 m isobaths. The Holocene transgression most probably varied the polynya location and size. Moreover, the Holocene transgression might have led to a modulated sea ice export that eventually influenced Atlantic water inflow. Therefore, Werner et al. [40] hypothesized that the flooding might be responsible for enhanced Atlantic water inflow detected in the Fram Strait 5 ka ago. As the sea level rise had a considerable effect on the ecosystem (e.g., Bauch et al. [22]; Stein et al. [19]) it is important to reconstruct sea ice distribution during this time.

According to the scientific relevance described above, this study presents IP25-(ice proxy with 25 carbon atoms; Belt et al. [41]) and PIP25-(phytoplankton biomarker-IP25 index; Müller et al. [42]) based reconstruction of sea ice distribution from the southern (Core BP99-04/7 and Core BP00-07/7) to the central Kara Sea (Core BP00-36/4) over the last 12.4 ka.

Biomarkers for Arctic sea ice and organic matter source

The sea ice biomarker IP25 (highly-branched isoprenoid (HBI)-monoene) is exclusively produced by Arctic sea ice algae and directly indicative for the absence or presence of sea ice above the study site [41, 43]. IP25 concentrations in surface sediment samples correlate well with satellite-sea ice data validating its application as sea ice proxy [42, 44, 45]. However, IP25 as single proxy does not allow the distinction between permanent sea ice cover and ice-free conditions. The molecule is absent when sea ice is missing offering no habitat for sea ice algae growth, but it is also absent if permanent thick sea ice covers the ocean preventing light penetration underneath the ice layer, necessary for sea ice algae growth. To distinguish between these two scenarios, the PIP25 index was introduced by Müller et al. [42]. It combines a phytoplankton biomarker (that is absent under permanent sea ice cover but abundant under ice-free conditions) with IP25. In this approach, the open-water phytoplankton biomarkers brassicasterol and dinosterol were used. Using the PIP25 index, Müller et al. [42] were able to classify different sea ice situations: A PIP25 index of 0 is indicative of ice-free conditions; values lower than 0.5 reflect variable sea ice cover; the range between 0.5 and 0.75 equals seasonal sea ice or ice-edge conditions and an index greater than 0.75 points to a permanent sea ice cover. More recently, Smik et al. [46] introduced an HBI-III alkene as a phytoplankton biomarker replacing the sterols in the PIP25 calculation. This modified PIP25 approach is based on biomarkers from the same group of compounds (i.e., HBIs) with more similar diagenetic sensitivity. Furthermore, it is far less dependent on the so-called “balance factor c” that is needed for PIP25 calculation (see “Organic geochemical analyses and proxy calculation” for calculation procedure). In studies of sediment cores from the northern Norwegian Sea, the western Barents Sea and the northern Barents Sea continental margin, Belt et al. [47] and Stein et al. [48] have calculated PIP25 values using brassicasterol and HBI-III as a phytoplankton biomarker. Importantly, these authors could demonstrate that both approaches yielded similar outcomes if core-specific balance factors were used (for further information and a critical discussion of the PIP25 approach as well as the different phytoplankton source indicators see Stein et al. [49]; Belt and Müller [50]; Belt et al. [47]; Xiao et al. [45]; Smik et al. [46]).

Meanwhile, IP25 as well as the PIP25 index are quite well established within the scientific community investigating paleo-sea ice cover in the central Arctic Ocean and its marginal seas from Quaternary, Pliocene to even Miocene time intervals (e.g., Müller et al. [51, 52]; Vare et al. [53]; Fahl and Stein [54]; Cabedo-Sanz et al. [55]; Stein and Fahl [56]; Berben et al. [57]; Knies et al. [58]; Müller and Stein [59]; Belt et al. [47]; Xiao et al. [60]; Hoff et al. [61]; Méheust et al. [62]; Stein et al. [48, 63, 64]).

As indicator for different sources of the organic matter specific sterols have been used in this study (for more information about the use of organic-geochemical proxies including biomarkers, we refer to Meyers [65] and Stein and Macdonald [66]). β-sitosterol and campesterol are predominantly produced by higher land plants (e.g., Huang and Meinschein [67]; Volkman [68]) and are commonly used as indicators for terrigenous sources (e.g., Fahl and Stein [69, 70, 71]; Xiao et al. [45, 60, 72]). Dinosterol and brassicasterol are produced by phytoplankton (e.g., Volkman [68, 73]; Volkman et al. [74]). However, it has to be considered, that brassicasterol is less definite in a riverine-dominated regime such as the Kara Sea [75]. Therefore, in this study the PIP25 index is calculated with dinosterol (see “Organic geochemical analyses and proxy calculation”), since brassicasterol is not a reliable marine organic matter tracer in our study area.

Material and method


The three investigated sediment cores were taken during RV Akademik Boris Petrov expeditions BP99 and BP00 in 1999 and 2000 [76, 77], carried out within the German–Russian research project Siberian River Run-off (SIRRO) [24]. These cores are located on a South–North transect on the Kara Sea shelf (Fig. 1a): Core BP99-04/7 is positioned in the Yenisei Estuary in a water depth of 32 m, ~ 112 km distant from the river mouth close to the Dickson Island. Core BP00-07/7 is situated ~ 255 km off the Yenisei delta between two small islands in a water depth of 38 m. Core BP00-36/4 were taken in the central Kara Sea in a water depth of 66 m, ~ 245 km further north from Core BP00-07/7. At site BP99-04, the first 30 cm of the sediment sequence of Core BP99-04/7 were disturbed during the coring procedure. Thus, a multicorer core taken from the same location (BP99-04/5), was correlated with the gravity core providing a complete record over the last ~ 10 ka (see Fahl and Stein [71]).

Chronology and lithology

The age models for Core BP99-04/7, Core BP00-07/7 and Core BP00-36/4 are based on 13, 9 and 6 radiocarbon AMS 14C (accelerator mass spectrometry) ages, respectively, measured on marine bivalves [19, 71, 78; see supplementary data Figs. S1–S3]. Due to the over-penetration of Core BP99-04/7 during the coring procedure, the multicorer Core BP99-04/5 (with one separate age determination) was added on top of the BP99-04/7 sediment sequence [19]. The AMS 14C ages were calibrated after Stuiver et al. [79; INTCAL98 radiocarbon age calibration], assuming a reservoir age of 440 years for the southern Kara Sea [80]. It should be considered that the reservoir age might vary during the last 12.4 ka [81]. However, we applied the published age model for a direct data comparison. Based on these age models Core BP00-36/4 dates back to 12.4 ka, Core BP00-07/7 back to 8.2 ka and Core BP99-04/7 back to 10.2 ka.

The main lithology of all three cores, described in detail by Stein et al. [19], consists of bioturbated silty clay to clayey silt. Bioturbation reduces the accuracy of the age model. This should be considered for the paleo-environmental reconstruction.

Organic geochemical analyses and proxy calculation

The total organic carbon (TOC) content was determined by means of the ELTRA Analyser. The biomarkers were extracted from 4 to 5 g of the freeze-dried, homogenized sediment using an accelerated solvent extractor (DIONEX, ASE 200; 100 °C, 5 min, 1000 psi) with dichlormethane:methanol (2:1 v/v) as solvent.

For quantification the internal standards [7-hexylnonadecane (0.076 µg/sample), squalane (2.4 µg/sample) and cholest-5-en-3β-ol-D6 (10.5 µg/sample)] were added prior to any analytical treatment.

For analysis of the sterols and hydrocarbons, the fractions were separated by means of open column chromatography (SiO2) using n-hexane (5 ml) as solvent for the hydrocarbons and ethylacetate:n-hexane 20:80 v/v (7 ml) for the sterols. Sterols were derivatized with 200 µl BSTFA (bis-trimethylsilyl-trifluoroacet-amide) for 2 h (60 °C). The biomarker analysis were performed with a Gas Chromatograph (GC) Agilent Technologies GC series 6850 and 7850 (30 m DB-1MS column, 0.25 mm inner diameter, 0.25 µm film, helium as carrier gas) coupled with an Agilent Technologies 5975C VL and 5977 A Extractor MSD mass selective detector (with Triple-Axis Detector, 70 eV constant ionization potential, Scan 50–550 m/z, 1 scan/s, ion source temperature 230 °C), respectively, by applying specific temperature programs: for the hydrocarbons, 60 °C (3 min), 150 °C (rate 15°C/min), 320 °C (rate 10°C/min), 320 °C (15 min isothermal) and for the sterols, 60 °C (2 min), 150 °C (rate 15°C/min), 320 °C (rate 3°C/min), 320 °C (20 min isothermal). The identification and quantification of the sterols [24-methylcholesta-5,22E-dien-3β-ol (brassicasterol), 4α–23,24-trimethyl-5α-cholest-22E-en-3β-ol (dinosterol), 24-methylcholest-5-en-3β-ol (campesterol), 24-ethylcholest-5-en-3β-ol (β-sitosterol) and the hydrocarbon (IP25)] were performed as described in Fahl and Stein [54].

The PIP25 index was calculated according to Müller et al. [42], applying dinosterol as phytoplankton biomarker (PIP25). As in our study the dinosterol concentrations are conspicuously higher compared to the IP25 concentrations a balance factor C has been introduced to compensate these diverged amounts (cf., Müller et al. [42]; Eq. 2).
$${\text{PI}}{{\text{P}}_{25}}={\text{ I}}{{\text{P}}_{25}}{\text{/}}\left( {{\text{I}}{{\text{P}}_{25}}+{\text{ }}\left( {{\text{phytoplankton biomarker}} \times C} \right)} \right).$$
$$C = {\text{ mean conc}}.{\text{ }}\left( {{\text{IP}}_{{25}} } \right){\text{/mean conc}}.{\text{ }}\left( {{\text{phytoplankton biomarker}}} \right).$$

All data are available online at  Pangaea—Data Publisher for and Environmental Science (, doi:10.1594/PANGAEA.887426).


Core BP99-04/7 from the Yenisei estuary

The content of TOC varies between 0.5 and 1.8% in the entire core (Fig. 3; cf., Stein et al. [19]).

Fig. 3

Biomarker records in µg/g sediment and the TOC content in percent of Core BP99-04/7 versus depth. Black stars mark the 14C AMS datings by Stein et al. [24] and Fahl and Stein [71]. The vertical line was set arbitrarily to highlight the reducing trends of the terrigenous biomarkers

Except in the bottom layer, IP25 is absent (or below the detection limit) in the lowest 40 cm (1.3 ka) of the sediment sequence (Fig. 3). In the following parts, the concentration raises to ~ 0.003 µg/g sediment, the average concentration up to 120 cm (2 ka), apart from four layers of higher concentrations at 590 cm (8.5 ka), 572 cm (8.3 ka), 420 cm (7.3 ka) and 255 cm (5 ka; 0.007, 0.007, 0.009 and 0.005 µg/g sediment, respectively). At 120 cm (2 ka), IP25 is absent, but it increases again towards the top to 0.005 µg/g sediment. In the multicorer core, IP25 reaches maximum values of 0.005–0.024 µg/g sediment.

Dinosterol is highly variable throughout the entire record, brassicasterol concentrations show a strong variation only in the upper 450 cm (7.5 ka) (Fig. 3). Brassicasterol steeply increases to 0.4 µg/g sediment within the lowest meter, whereas dinosterol increases gradually on average from 0.04 µg/g sediment to 0.43 µg/g sediment between 780 and 560 cm (10–8.2 ka). Then, it decreases to values of 0.15 µg/g sediment and fluctuates strongly around this value up to 120 cm (2 ka), similar to brassicasterol (around 0.1 µg/g sediment). In the upper 120 cm (2 ka), the concentrations of both sterols are in the range of 0.2 µg/g sediment for brassicasterol and 0.3 µg/g sediment for dinosterol.

The distribution patterns of β-sitosterol and campesterol are very similar (Fig. 3). However, the concentration of β-sitosterol is ten times higher than the concentration of campesterol. The concentrations of both sterols are higher between 780 and 680 cm (10–9 ka; on average 6.5 µg/g sediment for β-sitosterol and 1 µg/g sediment for campesterol), and lower between 680 and 590 cm (9–8.5 ka; on average 3.2 µg/g sediment for β-sitosterol and 0.6 µg/g sediment for campesterol). From 600 to 120 cm (8.6–2 ka), β-sitosterol and campesterol decrease from max. 5.7 and 0.9 µg/g sediment to 0.7 and 0.1 µg/g sediment, followed by a slight increase on average to 1.4 and 0.3 µg/g sediment towards the core top.

Core BP00-07/7 from the southern Kara sea

Generally, all biomarkers are distinctly variable in Core BP00-07/7 (Fig. 4). The TOC is mostly about 1%, apart from seven intervals of TOC contents higher than 1.3% at 700 cm (8.3 ka), 590–530 cm (7.7–7.2), 450 cm (6.5 ka), 370 cm (5.7 ka), 260–230 cm (4.1–3.8 ka), 80 cm (1.5 ka) and 40–0 cm (from 0.7 ka on).

Fig. 4

Biomarker records in µg/g sediment and the TOC content in percent of Core BP00-07/7 versus depth. Black stars mark the 14C AMS datings by Stein et al. [24], Simstich et al. [78] and Fahl and Stein [71]. The vertical line was set arbitrarily to highlight the trends

Brassicasterol and dinosterol vary between 0.05 and 0.7 µg/g sediment and 0.1–0.5 µg/g sediment, respectively. They increase from 200 cm (3.4 ka) to the core top, whereas β-sitosterol and campesterol (with concentrations between 0.2 and 1.4 µg/g sediment and 0.1 and 0.3 µg/g sediment, respectively) decrease throughout the record.

The IP25 concentrations are between 0.002 and 0.012 µg/g sediment. Higher values (more than 0.009 µg/g sediment) occur between 600 and 480 cm (7.8 and 6.8 ka). Afterwards, IP25 shows a slightly decreasing trend to 0.002 µg/g sediment up to 50 cm (0.9 ka). Over the entire record, prominent peaks of high IP25 concentrations (~ 0.006–0.009 µg/g sediment) occur at ~ 450 cm (6.5 ka), 380 cm (5.7 ka), 310 cm (4.8 ka), 110 cm (2.0 ka), 70 cm (1.3 ka), 25 cm (0.2 ka) and 10 cm (0,09 ka) and less high peaks (0.4 µg/g sediment) at ~ 360 cm (5.5 ka), 200 cm (3.4 ka), 150 cm (2.8 ka), 130 cm (2.4 ka), 95 cm (1.8 ka) and 50 cm (0.4 ka). These peaks coincide with low brassicasterol and dinosterol concentrations.

All biomarkers are highly variable within the upper 180 cm (3.2 ka) reaching both, the maximum and minimum concentrations.

Core BP00-36/4 from the central Kara sea

In Core BP00-36/4, the TOC content decreases from 1 to 0.55% between 560 and 370 cm (12.2 and 10.5 ka) (Fig. 5). Then it increases to 1.05% within 30 cm and decreases to 0.55% until 200 cm (9.7 ka). In the last 200 cm (9.7 ka), the content rises again to 1.05% at 130 cm (8.9 ka) and lowers towards the top to 0.6%.

Fig. 5

Biomarker records in µg/g sediment and the TOC content in percent of Core BP00-36/4 versus depth. Black stars mark the 14C AMS datings by Stein et al. [24], Simstich et al. [78] and Fahl and Stein [71]. The vertical lines were set arbitrarily to highlight the trends

At the bottom, IP25 first decreases from 0.006 to 0.003 µg/g sediment at 540 cm (11.8 ka)and then increases again up to 0.006 µg/g sediment (Fig. 5). From 520 to 350 cm (11.3–10.3 ka), the IP25 concentration lowers to 0.001 µg/g sediment and increases afterwards towards the core top to 0.0125 µg/g sediment [with a steep rise in the last 50 cm (6 ka)], except for a layer of absent IP25 at 260 cm (9.9 ka).

At 550 cm (12 ka), sterols are very low (brassicasterol 0 µg/g sediment; dinsterol 0.005 µg/g sediment; campesterol 0.01 µg/g sediment; β-sitosterol 0.25 µg/g sediment; Fig. 5). Dinosterol fluctuates on average 0.04 µg/g sediment between 520 and 350 cm (11.3 and 10.3 ka). Its concentration is higher between 350 and 320 cm (10.3 and 10 ka). (0.075 µg/g sediment). Brassicasterol is absent from the bottom to 275 cm (9.95 ka), then increases together with dinosterol [increase of dinosterol starts at 320 cm (10 ka)], to values of 0.05 µg/g sediment for brassicasterol and 0.07 µg/g sediment for dinosterol, apart from a phase between 175 and 130 cm (9.3 and 8.9 ka), where both sterol concentrations are lower (on average between 0 and 0.01 µg/g sediment for brassicasterol and down to 0.015 µg/g sediment for dinosterol). β-sitosterol shows a clear decreasing trend throughout the entire record, from 0.52 to minimal 0.1 µg/g sediment. Campesterol is more variable and fluctuates between 0.01 and 0.05 µg/g sediment within the entire sediment sequence.


On the Kara Sea shelf sediment deposition is mainly controlled by riverine terrigenous matter input [24, 69, 70, 82]. Furthermore, the freshwater as well as nutrients discharged by the rivers influence considerably the environmental conditions. Therefore, the distance to the river mouth determines to some parts the environment of a setting on the Kara Sea shelf. This is reflected in the South-North decrease in accumulation of terrigenous organic biomarkers. Areas that are situated more distant form the coast, are less supplied by terrigenous input and nutrients (Fig. 6). Therefore, marine productivity is also affected. This is indicated by decreasing concentrations of phytoplankton biomarkers from the southern to the central Kara Sea (Figs. 3, 4, 5). During the last 7 ka, when the sea level rise is completed and modern conditions are reached, on average 160 g/cm3 ka of terrigenous organic matter were deposited in the Yenisei Estuary, whereas the accumulation decreased by a factor of ~ 4–35 g/cm3 ka in the southern Kara Sea and down to 0.8 g/cm3 ka in the central Kara Sea.

Fig. 6

Accumulation of the terrigenous biomarkers (sum of β-sitosterol and campesterol) versus age in the cores BP99-04/7 (32 m water depth), BP00-07/7 (38 m water depth) and BP00-36/4 (66 m water depth), illustrated in a South-North transect on the Kara Sea shelf. The dotted lines (below) mark the sea level at ~ 9 and ~ 11 ka ago according to Stein et al. [24]

Deglacial to Holocene environment on the Kara sea shelf

During the last glacial maximum (LGM), the northern Kara Sea was covered by an ice-sheet, whereas the ice-sheet most probably did not extend to the southern Kara Sea (Svendsen et al. [83]; Polyak et al. [84]; Stein et al. [85]; Jakobsson et al. [86] and references therein). The shelf was exposed due to lowered sea level (~ 120 m; Fairbanks [21]). Consequently, huge ice sheets melted and the shelves were flooded during deglaciation between 16 and 7.5 ka (e.g., Bauch et al. [22]; Stein et al. [19]; Fig. 1a, b). The Holocene transgression changed significantly the environment of the Kara Sea shelf. The coastline and the main depocenter migrated southward during the sea level rise [18, 19, 20, 22, 87; Fig. 1b, c]. The sediment accumulation decreased at core sites positioned at the outer shelf, initiated by the southward shift of the main depocenter [19; Fig. 7a–c]. The post-glacial sea level rise is also reflected by the strong reduction of terrigenous matter accumulation observed in all studied cores. The distance between core sites and the river mouth grew while the shelves were flooded and less terrigenous material reached the core sites (Fig. 6). Moreover, some sharp reductions can be observed in the accumulation rates probably indicating abrupt climate changes (see Stein et al. [19]). While a decrease in the proportion of terrigenous biomarkers is recognizable in Core BP00-36/4 during this time, the proportion of marine biomarkers increased in the central Kara Sea (Fig. 5). This trend of the biomarkers documents the establishment of a marine milieu as also described by Fahl and Stein [71]. The two cores from the southern Kara Sea investigated here, might have been stronger influenced by the Holocene transgression. In both cores, the environmental transition is also represented by decreasing amounts of terrigenous and synchronously increasing marine biomarkers until ~ 7.5–6 ka (Figs. 3, 4). The core location of BP99-04/7 was flooded at ~ 9.3 ka (Fig. 1b, c) and modern conditions were reached approximately 2 ka later [19]. The lithological features of the sediment core (laminated subunit Ib, Stein et al. [19]) as well as the distinct change in the biomarker concentrations between 10.2 and 9.3 ka document the flooding of the core site (Fig. 3).

Fig. 7

I Schematic illustration of sea ice distribution in the southern Kara Sea (see also the legend on the right) and positions of the studied sediment cores. II IP25 concentrations (red dots mark dated samples), PIP25 index and accumulation rates of (a) Core BP99-04/7, (b) Core BP00-07/7 (the black dashed rectangle frame a period of distinct short-term sea ice variability; Hörner et al. [88]) and (c) Core BP00-36/4 versus age. Our own biomarker data are compared with the proportion of sea ice diatoms [89] and the magnetic susceptibility record as indicator for Siberian river discharge [20] of Core BP99-04/7, the incoming summer insolation at 75°N [90] and the NGRIP oxygen stable isotope record [91]. The horizontal black dashed line indicates the modern sea level stand in the Kara Sea [22, 24]. Red and blue hatched bars mark warmer and colder phases. The long blue arrows highlight general trends of the sea ice biomarkers

The sea ice records of Core BP00-36/4 represent deglacial climate changes in the central Kara Sea (Fig. 7 c). Between 12.4 and 11.2 ka, the PIP25 index fluctuates between 0.5 and 0.8. According to the classification of the PIP25 index by Müller et al. [42], seasonal to permanent sea ice cover was present at the core site within this interval. IP25 as well as phytoplankton biomarkers are present which points to a seasonal ice cover allowing phytoplankton growth. This changed temporarily to periods of a permanent ice cover with strongly restricted phytoplankton growth due to the lasting surficial light-shielding ice layer. These are most intense sea ice conditions in the central Kara Sea recorded by Core BP00-36/4. This time interval coincides with the Younger Dryas cold period (12.9–11.6 ka; Stuiver et al. [92]). The climatic deterioration was also recorded on Greenland by decreasing values in δ18O ice core data (NGRIP; Andersen et al. [91]; Fig. 7c) and pan-Arctic in several paleo-sea ice records. Synchronously, more intense sea ice cover was detected in the Bering Sea/North Pacific (Méheust [62]), in the Fram Strait [59], off northern Norway [55], in the western Barents Sea and northwest off Norway [47] and in the Laptev Sea at the western shelf as well as close to the interception with the Lomonosov Ridge [54].

IP25 concentrations and PIP25 values are low (less IP25 biomarkers and increased phytoplankton biomarkers) between 11 and 8 ka (Core BP00-36/4; Fig. 7c) reflecting less sea ice cover in the central Kara Sea after the cold interval. This warmer phase designated as the Holocene Thermal Maximum (HTM), was caused by the insolation maximum in the Northern Hemisphere (e.g., Koç et al. [93]; Bauch et al. [22]; Hald et al. [94]; Risebrobakken et al. [95]; Fig. 7c). The warmer climate was also documented in other parts of the high latitudes by reduced sea ice cover in the eastern and western Laptev Sea [54, 96] and by warmer and moister climate in the Siberian hinterland [97, 98]. Moreover, warmest atmospheric temperatures were recorded in Greenland ice cores during this interval [92, 99].

Generally, the sea ice indicators are strongly variable in Core BP00-36/4 between 12.4 and 8 ka comprising the colder (Younger Dryas) and warmer (HTM) climate phases (Fig. 7c). This suggests a dynamic system that reacts on climatic turnovers. In addition, three pulses of substantial changes in the accumulation rates occurred at ~ 11, 10 and 9 ka (Core BP00-36/4; Fig. 7c) reflecting climatic-induced changes [19]. These pulses are accompanied by alterations in sea ice cover (fluctuation in the PIP25 index; Fig. 7c). That means, sea ice cover changed conspicuously (also on shorter time scales) during the late Deglacial and the early Holocene in the central Kara Sea, most probably caused by climatic alterations.

Belt et al. [47] studied three sediment cores that are located in areas of different sea ice settings during modern times—from the Andfjorden, northwest off Norway; the Kveithola Trough in the western Barents Sea and from the Olga Basin in the northern Barents Sea. As mentioned above, the YD cold phase is documented in the cores from northwest off Norway and the western Barents Sea between 12.9 and 11.9 ka through extensive and a prolonged seasonal sea ice cover. Even the transition to the warmer HTM is reflected by a trend to ice-free conditions. Comparatively, in the Kara Sea, approximately 1700 km further east, seasonal sea ice prevailed throughout the HTM. Ice-free conditions at the entrance area of the Arctic Ocean might be caused by the influence of intensified Atlantic water inflow to the Barents Sea during the HTM. Moreover, Core BP00-36/4 is situated 2° further north and obviously prone to harsher climatic conditions.

However, there were no strong variations in the core from the northern Barents Sea during the entire Holocene, most probably reasoned by the more northern position of the core—located at ~ 78°N.

The mid Holocene environment: towards the modern sea level

The mid Holocene is also documented by the southern cores, Core BP99-04/7 and Core BP00-07/7 representing the last 9.3 and 8.2 ka, respectively. The sea ice records (IP25 concentration and PIP25 index) of the two cores show generally a different pattern as that presented by Core BP00-36/4 from the central part of the Kara Sea (Fig. 7). On a first view, this induces the impression of different environmental evolution in the southern and central Kara Sea during the mid to late Holocene. To interpret the IP25 and PIP25 records it is necessary to consider the modern sea ice distribution in the southern Kara Sea that is more complex compared to the central Kara Sea. Nowadays during spring, the site of Core BP99-04/7 is directly located at the border of coastal fast ice and the adjacent polynya (cf. Fig. 8I). A similar sea ice distribution prevails at the site of Core BP00-07/7, where fast ice is attached at the proximal islands and a polynya is formed in front (Figs. 2a, b; 8I). Fast ice extent in spring is mainly influenced by winter surface air temperature (Fig. 2c) and winds. Therefore, fast ice expansion is linked to climatic and atmospheric circulation changes [26, 100]. The climatic influence on fast ice distribution is also obvious from modern observations (see minimum and maximum fast ice distribution in the years 2012 and 2003; Fig. 2a). As IP25 and the PIP25 index are indicative for sea ice conditions during spring [101] when fast ice is formed [26], changes in the fast ice—polynya distribution are most likely recorded at cores BP99-04/7 and BP00-07/7.

Fig. 8

Schematic box models of sea ice distribution in the southern Kara Sea. The scenario I displays a normal situation, II shows an extreme cold environment with fast ice expansion to the core sites in the southern Kara Sea, III illustrates warmer conditions resulting in fast ice retreat and reduced sea ice

As both cores document a phase of lower IP25 concentrations and PIP25 values (low IP25 concentrations in combination with higher phytoplankton biomarker concentrations) between ~ 8 and 7.5 ka, the sea ice indicators may reflect a reduced sea seasonal drift ice cover at site BP00-07 and reduced fast ice at site BP99-04 (Figs. 7a, b; 8III). In addition, a warmer interval, i.e., the Mid-Holocene Climatic Optimum, characterized by warmer and moister climate was observed in the Siberian hinterland between 8.2 and 7.3 ka [102, 103]. It was also recorded by other studies from the southern Kara Sea. These studies reported a warmer climate determined by lower δ18O values measured on ostracods [78], higher river discharge [20] and enhanced biological productivity (also documented by higher phytoplankton biomarker concentrations in the southern cores; Figs. 3, 4) that was probably amplified by higher nutrient supply [71].

Intensifying seasonal sea ice cover is detectable since 8 ka at site BP00-07 (Fig. 8I; black arrow), indicated by an increased IP25 concentration and an elevated PIP25 index pointing to colder conditions. In the central Kara Sea (Core BP00-36/4) sea ice cover build-up started synchronously to the cooling determined at site BP00-07.

Sea ice variability during the late holocene

At Core BP00-36/4 accumulation rates dropped down at 8.5 ka due to the post-glacial sea level rise causing a migration of the main depocenter [22, 24]. Thus, the temporal resolution of the sea ice record is extremely reduced at Core BP00-36/4 during the late Holocene. However, the IP25 concentration and the PIP25 index display a general increase during the last 7 ka, indicating a cooling trend that allowed the intensified IP25 deposition due to stronger sea ice cover and harder conditions for phytoplankton assemblages. Most probably this was triggered by the decreasing incoming summer solar radiation [90; Fig. 7c]. The late Holocene cooling was observed in several studies from the Arctic region with colder climate on northern Greenland [91; Fig. 7c], a colder and drier environment in the Kara Sea region [102, 104], increasing sea ice cover in the Fram Strait [52], around Svalbard [105, 106, 107, 108] and in the Laptev Sea [54, 96, 109]. Moreover, the cooling trend is also documented in paleo-climate records from lower latitudes (see compilation by Wanner et al. [110]). As the environmental conditions were no longer affected by the post-glacial sea level rise, the increase of IP25 and PIP25 is solely controlled by climate change. A late Holocene centennial climatic variability observed globally in numerous studies (see summary by Wanner et al. [110]) cannot be resolved within the low-resolution sedimentary record of Core BP00-36/4 (Fig. 7c).

When the modern sea level conditions were established in the southern Kara Sea, approximately 7.5 ka ago [19; Fig. 7, dashed line], the PIP25 index measured on Core BP99-04/7 fluctuated between 0.75 and 0.2 (Fig. 7a). Following the classification by Müller et al. [42], these values reflect predominantly a seasonal sea ice cover (including sea ice edge conditions) at the core site during the whole time period. As the biomarker records most probably document the signal of the fast ice and the polynya, the fluctuations of the PIP25 index might indicate slight changes in extent or location of the polynya (Fig. 8). The lower PIP25 index (indicative for less sea ice cover) might reflect occasionally open-water conditions at site BP99-04 (Fig. 8III). But at five intervals, at ~ 7.3, 4.8, 4.4, 3.3 and 2 ka the PIP25 index is higher than 0.75, indicative for permanent sea ice cover. These phases might reflect long-lasting fast ice expansion around the coast to the core site during a temporally extremely cold environment with stronger winds (Fig. 8II). The phases at ~ 7.3 and 2 ka coincide with major changes in accumulation rates indicating a climatic-induced change [19; Fig. 7a]. Additionally, at 2 ka, a change of all measured proxies could be detected (Core BP99-04/7; Fig. 3). An intensified cooling in sub-Arctic regions, called Sub-Atlantic cooling (Fig. 7a, b), was also recognized by several environmental reconstructions in the Kara Sea region (e.g., Hahne and Melles [111]; Andreev and Klimanov [102]; Hubberten et al. [112]; Kraus et al. [113]), by increasing sea ice extent off northern Iceland [114] and glacier advances in northern Norway [115].

In general, the IP25 concentration of Core BP99-04/7 (this study) does not show a correlation to sea-ice diatom assemblages determined in the same sediment core by Polyakova and Stein [89] (Fig. 7a). In addition, a correlation between the phytoplankton biomarkers and the marine diatoms as well as between the terrigenous biomarkers and the freshwater diatoms could be also not identified. These differences may result from silica dissolution processes that reduced the amounts of the diatoms and may distort the results.

According to the modern distribution, a polynya situation was established close to the core site BP00-07 when the modern sea level stand was reached (Fig. 7I). Remarkably, the PIP25 index (values between 0.3 and 0.7) shows a decreasing trend during the last 7 ka indicative for reduced sea ice cover (Fig. 7b). This is in opposite to other late Holocene IP25 and PIP25 records mirroring an extended sea ice cover [47, 52, 54, 64, 96]. However, as discussed in Hörner et al. [88], the decrease in IP25 and PIP25 in Core BP00-07/7 follow the general cooling trend. At this setting, colder air temperature provided an expansion of the fast ice and a northward migration of the polynya (cf. Fig. 8II). Consequently, a more frequent polynya situation occurred at the core site BP00-07 leading to lower PIP25 values (and lower IP25 concentration).

In addition, distinct peaks (maxima) in the PIP25 index (and also in sterol concentrations; Fig. 4) could be detected at Core BP00-07/7 documenting Holocene centennial short-term climate variability (Fig. 7b). Most prominent PIP25 maxima of > 0.7 and up to 0.9 occurred at 5.9, 5.7, 2.9, 1.9, 1.2, 0.2 ka, additional distinct peaks (0.6–0.65) at 6.7, 6.3, 5.3, 5, 4.7, 3.5 ka. These short-term phases of intensified sea ice cover correlate with prominent cold short-term phases documented by various proxies within the Arctic, i.e., with phases of reduced river run-off [20] and less precipitation in the Kara Sea [104] (for detailed elucidation see Hörner et al. 2017). Short cold phases are also observed in North America and Fennoscandia [116, 117] and by phases of glacier advances on Franz Josef Land [118]. Between 4.5 and 3 ka, on the other hand, a phase characterized by less sea ice cover, i.e., low IP25 concentrations and PIP25 values at Core BP00-07/7, is recorded. Synchronously, increased precipitation [102, 104], stronger river discharge [20] and warmer climate in Siberia [102, 116, 117] as well as in Fennoscandia and northern America [116, 117; Fig. 7b, red bar] are observed.

The PIP25 peaks at ~ 3 and 2 ka record permanent sea ice cover at the core location (Core BP00-07/7; Fig. 7b, blue bars). Thus, most probably fast ice expands around the islands to the core site under extremely cold conditions. An amplified Holocene cooling trend during the last 3 ka is supported by several studies from different (sub-) Arctic areas (e.g., Grootes et al. [119]; Nesje et al. [115]; Kraus et al. [113]; Andrews et al. [120]).

The climatic characteristics of the Holocene, i.e., the cooling trend as well as the centennial short-term variability, are documented in Core BP00-07/7, but are not clearly recorded in Core BP99-04/7. Directly in the estuary where Core BP99-04/7 is located, river run-off has an effect on fast ice formation [26]. The immense freshwater supply as well as estuary processes proceeding at site BP99-04 may influence sea ice formation or sea ice algae growth and most probably overprint the climate signal at that location. In addition, the centennial climatic fluctuations are not documented in the sedimentary records from the Barents Sea [47], highlighting the climatic sensitivity of the marginal shelves.


The main results of this study can be summarized as follows:

  • Core BP00-36/4 from the central Kara Sea documents Arctic (global) post-glacial climate variability, whereas the southern cores (Core BP99-04/7, Core BP00-07/7) represent changes in the regional fast ice—polynya distribution.

  • In the central Kara Sea (Core BP00-36/4), sea ice cover was most intense (up to permanent sea ice cover) between 12.4 and 11.8 ka coinciding with the prominent Younger Dryas cold phase (12.9–11.6 ka). Between 10 and 8 ka, minimum sea ice was present (Holocene Thermal Maximum). The time period afterwards is documented in a low resolution due to decreased sedimentation rates. Although, increasing sea ice indicators imply the Holocene cooling trend during the last 7 ka that most probably followed the decreasing summer insolation.

  • Core BP00-07/7 records the Holocene cooling by decreasing IP25 and PIP25 values, due to the polynya situation at the core site. Moreover, centennial Holocene short-term (cyclic) climate variability is documented by changes in the sea ice distribution. PIP25 maxima up to 0.9 occurred at 5.9, 5.7, 2.9, 1.9, 1.2, 0.2 ka and distinct peaks (0.6–0.65) at 6.7, 6.3, 5.3, 5, 4.7, 3.5 ka. Consequently, this location is sensitive for abrupt short-term climatic variability that has a profound influence on climatic changes.

  • Core BP99-04/7 reflect changes in fast ice and polynya distribution. However, the strong influence of river run-off close to the Yenisei Estuary might prevent to some extent the detection of climatic signals. During the last ~ 9.3 ka (when the core site was flooded) seasonal sea ice cover was present, apart from five short phases at 7.2, 4.8, 4.4, 3.3 and 2 ka of permanent sea ice cover, i.e., most probably fast ice expansion to the core site (,,



We thank all members of the expeditions with RV Akademik Boris Petrov (BP99 and BP00; part of a German–Russian research project Siberian River Run-off (SIRRO), funded by the Federal Ministry of Education and Research) for providing the sediment material. Financial support by the Federal Ministry of Education and Research (Transdrift, Grant-No. 03G0833B) and the Alfred Wegener Institute is gratefully acknowledged. Thanks to Simon Belt and colleagues (Biogeochemistry Research Centre, University of Plymouth) for providing the internal standard for IP25 analysis.

Compliance with ethical standards

Conflict of interest

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

Supplementary material

41063_2018_40_MOESM1_ESM.pdf (94 kb)
Supplementary material 1 (PDF 93 KB)
41063_2018_40_MOESM2_ESM.pdf (87 kb)
Supplementary material 2 (PDF 87 KB)
41063_2018_40_MOESM3_ESM.pdf (93 kb)
Supplementary material 3 (PDF 93 KB)


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Copyright information

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

Authors and Affiliations

  1. 1.Helmholtz Centre for Polar and Marine ResearchAlfred Wegener InstituteBremerhavenGermany
  2. 2.Helmholtz Centre Potsdam, GFZ German Research Centre for GeosciencesPotsdamGermany
  3. 3.MARUM and Faculty of GeosciencesUniversity of BremenBremenGermany

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