Introduction
The northward flowing North Atlantic Current (NAC) is the most
important source of heat and salt in the Arctic Ocean (Gammelsrod and Rudels,
1983; Aagaard et al., 1987; Schauer et al., 2004; Fig. 1b). The main stream
of Atlantic Water (AW) flowing north to Fram Strait in the form of the West
Spitsbergen Current (WSC) causes a dramatic reduction of the sea-ice extent
and thickness via the warming of the intermediate water layer in this region
of the Arctic Ocean (Quadfasel et al., 1991; Serreze et al., 2003).
Palaeoceanographic (e.g. Spielhagen et al., 2011; Dylmer et al., 2013) and
instrumental (Walczowski and Piechura 2006, 2007; Walczowski et al., 2012)
investigations provide evidence of a recent intensification of the flow of AW
in the Nordic Seas and the Fram Strait.
The Svalbard archipelago is influenced by two water masses: AW flowing
northward from the North Atlantic and Arctic Water (ArW) flowing southwest
from the northern Barents Sea (Fig. 1b). An oceanic front arising at the
contact of different bodies of water is an excellent area for research of
contemporary and past environmental changes. Intensification of AW flow and
associated climate warming result in decreased sea-ice cover in the Svalbard
fjords during winter (Berge et al., 2006) and an increased sediment
accumulation rate (Zajączkowski et al., 2004; Szczuciński et
al., 2009) and influence the pelage–benthic carbon cycling (Zajączkowski
et al., 2010).
Palaeoceanographic records indicate that AW was present along the western
margin of Svalbard, at least during the last 12 000 years (e.g.
Ślubowska-Woldengen et al., 2007; Werner et al., 2011; Rasmussen et
al., 2013), and occasionally reached the Hinlopen Trough and Kvitøya
Trough, thus transporting warmer and more saline water to the eastern portion
of Svalbard from the north (Ślubowska-Woldengen et al., 2007;
Ślubowska-Woldengen et al., 2008; Kubischta et al., 2010; Klitgaard
Kristensen et al., 2013). Periods of enhanced inflow of AW during the
Holocene led to the expansion of marine species that are absent or only
rarely occurring at present. These species include the mollusc
Mytilus edulis whose fossil remains are widely distributed in raised
beach deposits on the western and northern coasts of Svalbard (e.g.
Feyling-Hanssen and Jørstad, 1950; Hjort et al., 1992). Mytilus edulis spawn at temperatures above 8–10 ∘C (Thorarinsdóttir and
Gunnarson, 2003) and thus are considered to indicate higher surface-water
temperature related to stronger AW inflow during the early Holocene
(11 000–6800 cal yr BP) (Feyling-Hanssen, 1955; Salvigsen et al., 1992;
Hansen et al., 2011). Although the progressive development of Mytilus edulis is well documented by periods of warming and inflow of AW to the
Hinlopen Trough, the presence of this species in Storfjorden (W Edgeøya;
Fig. 1) is unclear. Hansen et al. (2011) suggested that a small branch of
warm AW could have reached eastern Spitsbergen from the south at that time.
In the 1980s and 1990s, Storfjorden was thought to be exclusively influenced
by the East Spitsbergen Current (ESC), which carries cold and less saline ArW
from the Barents Sea (Quadfasel et al., 1988; Piechura et al., 1996). More
recent studies suggested that the hydrography in Storfjorden is affected by
the production of brine-enriched shelf waters (e.g. Haarpaintner et
al., 2001; Rasmussen and Thomsen, 2009), the creation of a coastal polynya
(e.g. Skogseth et al., 2005; Geyer et al., 2010) or the overflow of dense
waters to the continental shelf (e.g. Fer et al., 2003). However,
hydrological data obtained from conductivity–temperature sensors attached to
a Delphinapterus leucas showed a substantial and topographically
steered inflow of AW to Storfjorden through the Storfjordrenna (Lydersen et
al., 2002). Recently, Akimova et al. (2011) reviewed typical water masses for
Storfjorden, where the AW was located between 50 and 70 m.
Storfjordrenna is a sensitive boundary area (Fig. 1) where two contrasting
water masses form an oceanic polar front separating the colder, less saline
and isotopically lighter ArW from warmer, highly saline and δ18O-heavier AW. An abrupt cooling (e.g. Younger Dryas, Little Ice Age)
and warming (e.g. early Holocene warming) of the European Arctic might be
linked to relatively small displacements of this front (Sarnthein et
al., 2003; Hald et al., 2004; Rasmussen et al., 2014, 2015).
Location map (a) showing the core site from this study
(JM09-020-GC) and core site of JM02-460 (Rasmussen et al., 2007). The inlet
map (b) shows the modern surface oceanic circulation in Nordic Seas
and location of a core NP94-51 (Ślubowska et al., 2005). Abbreviations:
NAC: Norwegian–Atlantic Current; WSC: West Spitsbergen Current; ESC: East
Spitsbergen Current; EGC: East Greenland Current; NC: Norwegian Current. The
cores JM02-460 and NP94-51 are discussed in the text.
Two sediment cores collected at the mouth of Storfjordrenna reveal a
continuous inflow of AW to the southwestern Svalbard shelf since the
deglaciation of Svalbard–Barents Ice Sheet (Rasmussen et al., 2007), whereas
the inner Storfjorden basins underwent a shift from occupation by continental
ice to an ice proximal condition (Rasmussen and Thomsen, 2014, 2015).
Nevertheless, a limited amount of palaeoceanographical data are available
from this region, and thus the reconstruction of the Svalbard–Barents Ice
Sheet retreat and the further development of Storfjordrenna oceanography are
often speculative.
In this paper, we present results from multi-proxy analyses of a sediment
core retrieved 100 km east of the mouth of Storfjordrenna (Fig. 1a). We
provide a new age for the retreat of the last Svalbard–Barents Sea Ice Sheet
from Storfjordrenna and discuss the interaction of oceanography and
deglaciation as well as the postglacial history of Atlantic Water inflow onto
the shelf off of southern Svalbard. Because the studied sediment core was
retrieved from an oceanographic frontal zone, which is sensitive to
larger-scale changes, we believe that the presented data show the general
climatic/oceanographic trends in the eastern Arctic.
Oceanographic setting
Storfjorden is an approximately 190 km long and up to 190 m deep glacial
trough located between the landmasses of Spitsbergen to the west, Edgeøya
and Barentsøya to the east, and the shallow Storfjordenbanken to the
southeast (Fig. 1a). It is not a fjord sensu stricto because the sounds of
Heleysundet and Freemansundet to the north and northeast, respectively,
connect the head of Storfjorden to the northwestern Barents Sea. A sill of
120 m depth crosses the mouth of Storfjorden. The 254 km long
Storfjordrenna, a continuation of the trough that extends towards the shelf
break, is located beyond this sill. The bottom depth along the trough axis
varies between 150 and 420 m (Pedrosa et al., 2011).
The water column of Storfjorden and Storfjordrenna is composed of two main
water masses transported with currents from the east and south and mixed
waters that are formed locally (Table 1 after Skogseth et al., 2005). Warm
and saline AW enters Storfjordrenna in a cyclonic manner
(Schauer, 1995; Fer et al., 2003), flowing into the trough parallel to its
southern margin and flowing towards the trough mouth along its northern
slope. The AW occurs between 50 and 70 m in Storfjorden and extends to a
depth of 200 m in Storfjordrenna (Akimova et al., 2011). The origin of AW
entering Storfjordrenna is an eastward branch of the North Atlantic Current
(NAC) following the topography of the Barents Sea Shelf Break. However,
approximately 50 % of the AW flowing northward also penetrates into
Bjørnøyrenna (Smedsrud et al., 2013; for location, see Fig. 1). The AW
in Storfjordrenna is cooler and fresher than in Bjørnøyrenna as an
effect of the distance and mixing processes (O'Dwyer et al., 2001). The AW
may occasionally propagate even further east of Svalbard, where it fills
depressions below 180 m (Schauer, 1995). Relatively cold ArW
is transported to Storfjorden and Storfjordrenna by the ESC. The ESC enters the fjord through the tidally influenced sounds
of Heleysundet and Freemansundet in the north and northeast (Norges
Sjøkartverk, 1988) as well as from the southeast with a coastal current
flowing near Edgøya (Loeng, 1991). The AW and ArW mix to form Transformed
Atlantic Water (TAW), which dominates on the shelf off of West Spitsbergen
(Svendsen et al., 2002; Table 1). Dense, brine-enriched Shelf Water (BSW) in
Storfjorden is produced through high polynya activity and results from
intense formation of sea-ice (Haarpaintner et al., 2001; Skogseth et
al., 2004, 2005). The BSW fills the fjord to the top of the sill (120 m) and
initiates a gravity-driven overflow (Quadfasel et al., 1988; Schauer, 1995;
Schauer and Fahrbach, 1999; Fer et al., 2003, 2004; Skogseth et al., 2005).
The BSW is characterised by a salinity value greater than 34.8 and a
temperature at or slightly above the freezing point (Table 1). Surface Water
(SW) in the upper 50 m is cold and fresh during the autumn and warm and
fresh during the summer due to ice melting. In winter, the water column in
Storfjorden is homogenised due to wind and tidal mixing and is considered to
have a temperature close to the freezing point (Skogseth et al., 2005).
Water mass characteristics in Storfjorden and Storfjordrenna
(Skogseth et al., 2005, modified). The two main water masses are in bold.
Water mass names
Water mass
characteristics
Temperature
Salinity
(∘C)
Atlantic Water (AW)
>3.0
>34.95
Arctic Water (ArW)
<0.0
34.3–34.8
Brine-enriched Shelf Water (BSW)
< -1.5
> 34.8
Surface Water (SW)
> 0.0
< 34.4
Transformed Atlantic Water (TAW)
> 0.0
> 34.8
Materials and methods
Multi-proxy analyses of the gravity core JM09-020-GC provided the foundation
for this study. The core was retrieved with R/V Jan Mayen
(University of Tromsø – The Arctic University of Norway, UiT) in November
2009 from the Storfjordrenna (76∘31489′ N, 19∘69957′ E)
at a bottom depth of 253 m (Fig. 1a). The coring site was located in an area
above the continuous presence of BSW and was selected after an echo-acoustic
investigation to identify the greatest possible area of flat bottom with a
minimum disturbance of sediments. Conductivity–temperature–depth (CTD)
measurements were performed prior to coring (Fig. 2a) and in summer 2013
(Fig. 2b).
Prior to sediment core opening, the magnetic susceptibility (MS) was measured
using a loop sensor installed on a GEOTEK Multi Sensor Core Logger at the
Department of Geology, UiT. Core sections were stored in the laboratory for
one day prior to measurements, thus allowing the sediments to adjust to room
temperature and avoiding measurement errors related to temperature changes
(Weber et al., 1997). The X-radiographs and digital images were collected
from half of the core to define the sedimentary and biogenic structures. The
sediment colour was defined according to the Munsell Soil Color Charts
(Munsell Products, 2009). Qualitative element geochemical measurements were
performed with an Avaatech X-ray fluorescence (XRF) core scanner using the
following settings: 10 kV, 1000 µA, 10 s measuring time, and no
filter. Both core halves were subsequently cut into 1 cm slices and
transported to the Institute of Oceanology at the Polish Academy of Sciences
in Sopot for further analyses.
Sediment samples for foraminiferal analyses were freeze-dried, weighed, and
wet sieved using sieves with mesh sizes of 500 and 100 µm. The
residues were dried, weighed again and subsequently split on a dry
micro-splitter. Where possible, at least 300 specimens of foraminifera were
counted in every 5 cm of sediment. Species identification under a binocular
microscope (Nikon SMZ1500) was supported using the classification of Loeblich
and Tappan (1987), with few exceptions, and percentages of the eight
indicator species were applied. The number of species per sample and
Shannon–Wiener index were calculated using the program Primer 6. The benthic
foraminiferal abundance and ice-rafted debris (IRD; grains
> 500 µm) were counted under a stereo-microscope and expressed
as flux values (number of specimens/grains cm-2ka-1) using the
bulk sediment density and sediment accumulation rate.
Stable oxygen and carbon isotope compositions of tests of the infaunal
foraminifer species Elphidium excavatum f. clavata were
determined at the Department of Geological Sciences, University of Florida
(Florida, USA). All values are calibrated to the Pee Dee Belemnite (PDB) scale
and corrected for ice volume changes. In our study, we discuss the
δ18O and δ13C record as a relative measure for changes in
the water mass characteristics (temperature–salinity) and/or the supply of
meltwater/freshwater to the area. Moreover, no reliable vital effect
correction has been created for E. excavatum f. clavata
(Bauch et al., 2004; Ślubowska-Woldengen et al., 2007), and therefore we
did not correct the isotopic values for vital effect.
Grain size (< 2 mm) analyses were performed every 1 cm using a
Malvern Mastersizer 2000 laser particle analyser and presented as volume
percent. To examine the relative variability in the near-bottom currents,
the mean grain-size distribution of the < 63 µm fraction was
calculated to avoid the effect of ice-rafted coarse fraction. The mean grain
size was calculated using the program GRADISTAT 8.0 according to the
geometric method of moments (Blott and Pye, 2001).
The chronology for this study is based on high-precision AMS 14C
measurements of fragments from nine calcareous bivalve shells. Measurements
were performed in the Poznań Radiocarbon Laboratory, which is equipped
with a 1.5 SDH-Pelletron Model “Compact Carbon AMS” (Czernik and Goslar,
2001; Goslar et al., 2004). The surface layer of shells was scraped off to
avoid contamination with younger carbonate encrustation. The AMS 14C
dates were converted into calibrated ages using the calibration program
CALIB 6.1 (Stuiver and Reimer, 1993; Stuiver et al., 2005) and the Marine13
calibration curve (Reimer et al., 2013). The difference ΔR in
reservoir age correction of the model ocean and region of Svalbard was
reported by Mangerud et al. (2006) as 105±24 or 111±35, and we used
the first value. The calibrated ages are presented in Table 2. It should be
noted that the reservoir age is based on a few data points from western
Spitsbergen, and the age may be different for the eastern coast. However, no
data are available from the latter region.
AMS 14C dates and calibrated ages.
Sample No
Depth
Lab No
Raw AMS 14C
Calibrated years
Cal yr BP
Dated material
(cm)
BP
BP ± 2σ
used in
age model
St 20A 5/6
5
Poz-46955
1835±30
1200–1365
1285
Ciliatocardium ciliatum
St 20A 39
38.5
Poz-46957
2755±30
2245–2470
Not used
Astarte crenata
St 20 78/79
78
Poz-46958
2735±30
2177–2429
2320
Astarte crenata
St 20 110
109.5
Poz-46959
3450±30
3079–3323
3220
Astarte crenata
St 20 142
141.5
Poz-46961
6580±40
6850–7133
6970
Astarte crenata
St 20A 152
151.5
Poz-46962
7790±40
8018–8277
8160
Astarte crenata
St 20 157
156.5
Poz-46963
8610±50
8989–9288
9120
Bathyarca glacialis
St 20 251/252/253
252
Poz-46964
10 200±60
10 895–11 223
11 230
Thracia sp.
St 20 396
395.5
Poz-46965
12 570±60
13 780–14 114
13 950
Bivalvia shell
Temperature and salinity versus depth, measured on 5 November 2009
(a) and on 13 August 2013 (b) at the site of core
JM09-020GC. SW: Surface Water, TAW: Transformed Atlantic Water, BSW:
Brine-enriched Shelf Water.
Results
Modern hydrology
In November 2009, the SW at the coring site (upper ∼ 27 m)
had already cooled (1.24 ∘C; Fig. 2a); however, its salinity was
still low (34.24). Transformed AW was observed in the layer between 60 and
160 m. The lowermost portion of the water column shows evidence of gradual
cooling that reached a minimum temperature of 0.76 ∘C near the
bottom. The lack of BSW at the bottom indicates gradual water mixing during
summer and fall. In August 2013, the SW had a slightly lower
salinity, but the temperature was ∼ 5 ∘C higher than in
November 2009 (Fig. 2b). The TAW occupied the same depths as in 2009.
However, an almost 50 m thick layer of BSW was present close to the
seafloor.
Age model
The 14C ages and calibrated ages are reported in Table 2. The
calibration gives an age distribution and not a single value; thus, the
2-sigma range is presented, and Fig. 3 shows the age probability distribution
curves. The ages of the samples generally increase with sediment depth except
in the case of one sample, namely St 20A 39, which provided an older age than
the sample below. That shell was most likely re-deposited and thus was not
used for the age model. However, because all of the samples used for dating
were shell fragments, it must be noted that it is possible that more samples
could be subjected to re-deposition, but based on the available data, it is
not possible to confirm. The age model is based on the assumption of linear
sediment accumulation rates between data points. The highest probability
peaks from the calibrated age ranges were used as input values for the model.
For the lowermost and uppermost regions of the core, we adopted sediment
accumulation rates for the neighbouring region. It is common to observe the
loss of the sediment surface layer during coring with heavy gravity cores. In
the case of core JM09-020-GC, it is likely that at least the top 40 cm of
sediments were lost during coring. This conclusion is supported by analysis
of a box corer collected prior to coring (Łącka et al., 2015). The
extrapolated age model for the sediment surface is therefore
1200 cal yr BP.
Age–depth relationship for JM09-020-GC based on eight AMS 14C
calibrated ages with 2-sigma age probability distribution curves. The
chronology is established by linear interpolation between the calibrated
ages.
Sedimentological and geochemical parameters
The core JM09-020-GC is 426 cm long and consists of four lithological units:
L1 (bottom of the core to 370 cm; > 13 450 cal yr BP), L2
(370–272 cm; ∼ 13 450 to ∼ 11 500 cal yr BP), L3
(272–113 cm; ∼ 11 500 to ∼ 3600 cal yr BP) and L4 (113 cm
to core top; ∼ 3600 to ∼ 1200 cal yr BP). The lithological log
was created based on the X-radiographs, grain-size analysis data and
foraminiferal flux (Fig. 4). Grains > 2 mm are referred to as “clasts”
and are marked in the lithological logs as individual features.
Unit L1 consists of compacted massive dark grey (5Y 4/1) sandy mud with
various amounts of clasts. Bioturbation and foraminifera were generally
absent. However, one shell fragment was found at approximately 395 cm.
Unit L2 contains massive dark grey (5Y 4/1) sandy mud with an amount of
coarser material and generally lower amounts of clasts than unit L1. The mean
grain size (< 63 µm) ranged from 7 to 10 µm. The
highest IRD flux and Fe / Ca ratio for the entire core occur in this unit.
The mass accumulation rate (MAR) is 0.043 gcm-2yr-1. The
first signs of bioturbation occur in this unit, and the flux of foraminifera
increases rapidly up to ∼ 5700 individualscm-2ka-1
(Fig. 4).
The unit L3 is composed of massive dark olive grey mud (5Y 3/2) and is
characterised by decreasing MAR values (0.019–
0.002 gcm-2yr-1), moderate sand content and clearly
increasing mean grain size (< 63 µm). The IRD flux is low, and
the Fe / Ca ratio decreases gradually until ca. 9200 cal yr BP and
remains low (between 3 and 4; Fig. 4) Continuous bioturbation and variable
foraminiferal fluxes are observed, with maxima in the intervals 9000–8000
and 6000–5500 cal yr BP.
The uppermost unit L4 is primarily composed of the same material as the
underlying unit, i.e. massive dark olive grey mud (5Y 3/2). However, the
sand content is occasionally higher. The MAR increases to
0.024 gcm-2yr-1. The mean grain size
(< 63 µm) throughout this interval is even higher than that in
L3 and reaches up to 15 µm; the Fe / Ca ratio is increasing.
The bioturbation continues, numerous shell fragments are present, and the
foraminifera flux reaches high values throughout the entire unit.
Lithological log of core JM09-020GC. Lithology, 14C dates,
occurrence of bioturbation, mass accumulation rates, mean grain size in the
range of 0–63 µm, sand content, ice-rafted debris flux, magnetic
susceptibility, foraminifera flux as well as Fe / Ca ratio and water
content. The results are presented with lithostratigraphic units (L1–L4),
versus calendar years (cal kyr BP) and core depth (cm).
Foraminiferal fauna
A total of 54 calcareous and six agglutinated species were identified. The
foraminiferal assemblages were dominated by calcareous fauna. Agglutinated
species occurred only in 14 sediment samples, and their abundance did not
exceed 4 %. The only exception is the sample dated to ca.
11 350 cal yr BP (262.5 cm depth) with 25 % of agglutinated
foraminiferal fauna. However, in this sample, the total foraminifera
abundance was low (13 specimensg-1sediment). In general,
species richness, number of agglutinated foraminifera, and rare and fragile
species increase towards the top of the core. Benthic foraminiferal fauna is
dominated by Elphidium excavatum f. clavata,
Cassidulina reniforme, Nonionellina labradorica,
Melonis barleeanum, Islandiella spp. (Islandiella norcrossi/Islandiella helenae) and Cibicides lobatulus.
Percentages of E. excavatum f. clavata show an inverse
relationship to C. reniforme with the almost constant dominance of
the latter species in the periods ∼ 12 450 to ∼ 12 000
cal yr BP and ∼ 9600 to ∼ 2800 cal yr BP (Fig. 5).
Planktonic foraminifera are represented by three species:
Neogloboquadrina pachyderma (sinistral), Neogloboquadrina pachyderma (dextral) and Turborotalita quinqueloba. However, the
two latter species are quite rare. In general, the abundance of planktonic
fauna is low in the older portions of the core and slightly increases at
approximately 10 000 cal yr BP, reaching maximum values ca.
2000 cal yr BP (Fig. 5).
Based on the most significant changes in the foraminiferal species
abundances, species diversity, and δ18O and δ13C in
E. excavatum f. clavata tests, the core was divided into
four foraminiferal zones F1–F4: ∼ 13 450–11 500 cal yr BP (F1);
11 500–9200 cal yr BP (F2); 9200–3600 cal yr BP (F3); and
3600–1200 cal yr BP (F4) (Fig. 5). The zones correspond to lithological
divisions. The age of unit F4 is the same as L4, units F3 and F2 correspond
to L3, and unit F1 is linked to unit L2. In unit L1, foraminifera are rare to
absent.
Zone F1 is dominated by the opportunistic E. excavatum
f. clavata and C. reniforme. The latter species dominates
more than E. excavatum f. clavata between 12 250 and
11 950 cal yr BP. High percentages of C. lobatulus (up to
57 %) and Astrononion gallowayi (up to 2.5 %) occur
occasionally. The planktonic foraminifera flux was low at the beginning of
this section (mean value of nine specimenscm-2ka-1) and
completely disappeared for nearly 1500 years from approximately
11 500 cal yr BP (Fig. 5). The species richness and the Shannon–Wiener
index show low biodiversity compared with the upper portion of the core (mean
values of 8 and 1.26, respectively). Furthermore, maxima of δ18O
and δ13C occur in this interval.
Percentage distributions (upper scale; black line) of the most
dominant benthic species, fluxes (no. cm-2ka-1; bottom scale;
grey shading) of benthic and planktonic foraminiferal species, diversity
parameters (species richness and Shannon–Wiener index) and stable oxygen and
carbon isotope data (δ18O and δ13C) plotted versus
thousands of calendar years with indicated foraminiferal zonation (zones
F1–F4) and lithostratigraphic units (L1–L4). Foraminiferal taxa are grouped
based on their ecological tolerances described in the text.
In zone F2, the contribution of E. excavatum f. clavata and
C. reniforme is slightly lower, and N. labradorica becomes
the most abundant species (Fig. 5). There is also an increase in
Islandiella spp. percentage. Planktonic foraminifera appeared again
ca. 10 000 cal yr BP. Biodiversity significantly increased, and
δ18O reached its minimum value of 2.61 ‰ vs. Vienna Pee Dee Belemnite at
approximately 10 000 cal yr BP.
Zone F3 is characterised by the minimum mass accumulation rates of sediment
and consequent low temporal resolution. C. reniforme dominates over
E. excavatum f. clavata throughout. M. barleeanum
has its maximum abundance in this zone, and N. labradorica is
abundant in the lower portions of this zone, decreasing at approximately
7000 cal yr BP. Islandiella spp. increases upcore. Planktonic
foraminifera occur in the entire zone, and the fluxes are higher than those
of previous units (Fig. 5). Biodiversity remains high in this zone, and
δ18O and δ13C remain generally stable; however, marked
peaks occurred at approximately 6800, 6500 and 5700 cal yr BP,
respectively.
A consistently high foraminiferal flux of up to
∼ 4900 specimenscm-2ka-1 characterises zone F4. The
fluxes of Islandiella spp. and Buccella spp. increase
significantly, and from 2850 cal yr BP, Islandiella spp.
dominated the assemblage with E.excavatum f. clavata.
Additionally, the fluxes of C. lobatulus and A. gallowayi
increase; however, their abundances are lower than those of zone F2. A
maximum abundance of planktonic foraminifera occurs in this unit.
Foraminifera biodiversity continues to increase towards the core top (up to
2.33; Fig. 5), and δ18O and δ13C increase slightly with
numerous fluctuations.
Discussion
The European Arctic includes continental slope strongly influenced by
northward flowing Atlantic water and large shelf of the Barents Sea
characterised by less saline and colder water. The available broad range of
studies concerning palaeoceanography of the European Arctic focus on its
marginal sites: westernmost (e.g. Rasmussen et al., 2007; Eldevik et
al., 2014; Sternal et al., 2014), northern (Wollenburg et al., 2004;
Klitgaard Kristensen et al., 2013) and eastern (Polyak and Solheim, 1994),
while the border zone lying between the slope of continental shelf and
central Barents Sea is poorly studied. The lack of well-defined and
sufficiently complete palaeoceanographic record containing the signal from
both of these environments encouraged the authors to study a sediment core
retrieved inside Storfjordrenna, especially in the light of current
hydrological changes in this area (e.g. Lydersen et al., 2002; Skogseth et
al., 2005; Akimova et al., 2011). This location should present the general
trends in the eastern Arctic, including Svalbard glacier activities, pack-ice
in the Arctic Ocean and North Atlantic water circulation, moreover it avoids
the local (fjordic) condition. We decided to discuss the presented record
chronologically as a postglacial interplay between two hydrological regimes.
Based on the most pronounced changes in sedimentological and foraminiferal
data as well as comparisons with previous studies from adjacent areas, we
distinguish five units in the studied core: a subglacial unit
(> 13 450 cal yr BP), a glacier-proximal unit
(13 450–11,500 cal yr BP), a glaciomarine unit I
(11 500–9200 cal yr BP), a glaciomarine unit II
(9200–3600 cal yr BP) and a glaciomarine unit III
(3600–1200 cal yr BP).
$\,13\,450\,cal\,yr\,BP)}?>Subglacial unit (> 13 450 cal yr BP)
The lowermost unit L1 (Fig. 4) was significantly coarser, more compacted and
devoid of foraminifera, which indicates that it is likely of subglacial
origin. During the late Weichselian Glacial Maximum, Storfjorden and
Storfjordrenna were covered by an ice stream that drained the
Svalbard–Barents Ice Sheet (SBIS; e.g. Ottesen et al., 2005). The SBIS
deglaciation occurred as a response to the sea-level rise and increased mean
annual temperature (Siegert and Dowdeswell, 2002). Rasmussen et al. (2007)
noted that the outer portion of Storfjordrenna (389 m depth; Fig. 1a) was
deglaciated prior to 19 700 cal yr BP. The bivalve shell fragment from
395.5 cm in our core suggests that the centre portion of Storfjordrenna was
ice-free before ∼ 13 950 cal yr BP. This observation indicates that
the ∼ 100 km long retreat of the grounding line from the shelf break
to the central portion of Storfjordrenna occurred over approximately
5700 years. The deglaciation of the inner Storfjorden basin occurred ca.
11 700 cal yr BP (Rasmussen and Thomsen, 2014), whereas the coasts of the
east Storfjorden islands, Barentsøya and Edgeøya, which are located
over 100 km north from the coring site, occurred some 500 years later, i.e.
11 200 cal yr BP (recalibrated after Landvik et al., 1995). Siegert and
Dowdeswell (2002) noted that during the Bølling–Allerød warming (ca.
14 700–12 700 cal yr BP), certain of the deeper bathymetric troughs
(e.g. Bjørnøyrenna) had deglaciated first, and large embayments of ice
formed around them. It is likely that Storfjordrenna was one of such
embayments at that time. Our data are in agreement with ice stream retreat
dynamics presented by Rüther et al. (2012) and refine the recent models
of the Barents Sea deglaciation (e.g. Winsborrow et al., 2010; Hormes et
al., 2013; Andreassen et al., 2014).
Glacier-proximal unit (13 450–11 500 cal yr BP)
The transition from a subglacial to glaciomarine setting is observed as a
distinct change in sediment colour, several peaks of IRD, a decreased amount
of clasts and the appearance of foraminifera. The sediment accumulation rate
(0.043 gcm-2yr-1) was of the same order of magnitude as
that of the modern proximal and central regions of the West Spitsbergen
fjords (see Szczuciński et al., 2009 for a review). Textural and
compositional analyses of L2 recorded a bimodal grain-size distribution and
low abundance of microfossils, suggesting that deposition during the
deglaciation occurred due to suspension settling from sediment-laden plumes
and ice rafting (Lucchi et al., 2013; Witus et al., 2014). This unit in our
core is limited to ∼ 60 cm and is characterised by a lack of
bioturbation in its lower portion.
The high flux of IRD is supported by the high Fe / Ca ratio and the
depleted δ18O values correlate well with the abundance of
C. lobatulus and A. gallowayi (Figs. 4 and 5), two species
connected with high-energy environments (Østby and Nagy, 1982), thus
indicating that the coring site was likely located proximal to one or several
ice fronts during the time of deposition of this unit.
During an early phase of the deglaciation of Storfjorden, the East
Spitsbergen Current was still not active because the ice sheet grounded
between Svalbardbanken and Storfjordbanken blocked the passage between
eastern and western Svalbard (Rasmussen et al., 2007; Hormes et al., 2013).
Thus, the first foraminiferal propagules (juvenile forms) were transported by
sea currents (Alve and Goldstein, 2010) from the south and west and settled
on the seafloor that was exposed after the retreat of grounded ice. The
proximal glaciomarine environment affected the foraminiferal assemblages and
resulted in low species richness, biodiversity and low foraminiferal
abundance. Consequently, foraminifera assemblages became dominated by fauna
typical of the glacier proximal settings: E. excavatum
f. clavata, C. reniforme and Islandiella spp.
(e.g. Vilks, 1981; Osterman and Nelson, 1989; Polyak and Mikhailov, 1996;
Hald and Korsun, 1997). The dominance of E. excavatum
f. clavata confirms the proximity to the ice sheet, decreased
salinity and high water turbidity (e.g. Steinsund, 1994; Korsun and Hald,
1998; Włodarska-Kowalczuk et al., 2013).
The upper portion of unit L2 (ca. 12 800–11 500 cal yr BP) spans the
Younger Dryas (YD) stadial. Records of marine sediments from Nordic and
Barents Seas (e.g. Rasmussen et al., 2007; Ślubowska-Woldengen et
al., 2007, 2008; Zamelczyk et al., 2012; Groot et al., 2014) as well as
δ18O records from Greenland ice cores (e.g. Dansgaard et al., 1993;
Grootes et al., 1993; Mayewski et al., 1993; Alley, 2000) show that the YD
was characterised by a rapid and short-term temperature decrease. This event
was likely driven by the weakened North Atlantic Meridional Overturning
Circulation, a result of the Lake Agassiz outburst (e.g. Gildor and
Tziperman, 2001; Jennings et al., 2006; Murton et al., 2010; Cronin et
al., 2012) or the interaction between the sea-ice and thermohaline water
circulation (Broecker, 2006), which led to a reduction of AW transport to the
north and a dominance of fresher Arctic Water. Our data show that the heavier
δ18O values recorded, e.g. 12 720 and 12 100 cal yr BP,
correlate with reduced to absent IRD fluxes, whereas the peaks of lighter
δ18O, e.g. 12 450, 12 150, and 11 780 cal yr BP, occurred
synchronously with significant enhanced IRD fluxes (Fig. 6). The absence of
IRD, occasionally for several decades, might reflect temporary polar
conditions (Dowdeswell et al., 1998; Gilbert, 2000) characterised by the
formation of perennial pack ice in Storfjorden that locked icebergs proximal
to their calving fronts and prevented their movement over the coring site
(Forwick and Vorren, 2009). Wollenburg et al. (2004) observed a decrease in
palaeoproductivity on the northern Barents Sea margin between 12 800 and
12 500 cal yr BP and the later palaeoproductivity peak at the termination
of YD; they concluded that permanent sea-ice cover causes the decrease in sea
productivity, whereas enhanced advection of Atlantic Water to the site might
result in palaeoproductivity increase. Those periods of accelerated AW inflow
resulted in massive iceberg rafting and delivery of IRD to Storfjordrenna,
thus reflecting more sub-polar conditions. Hydrological variability during
the Younger Dryas was previously noted in selected circum-North-Atlantic
deep-water records (Bakke et al., 2009; Elmore and Wright, 2011 and
references therein; Pearce et al., 2013). Moreover, oxygen stable isotope
records from an ice-core GISP2 show certain warmer spells during that time
(Stuiver et al., 1995), which coincides with higher ice rafting in
Storfjordrenna (Fig. 6). Bakke et al. (2009) noted that the earlier portion
of YD was colder and more stable, whereas the latter portion of this period
was characterised by alternations between sea-ice cover and an influx of
warmer and saltier North Atlantic waters. Our records show that during the
late YD, the δ18O data were slightly shifted towards lighter
values. Temporal resolution of our records does not allow for more detailed
comparison with available data; nevertheless, they clearly indicate that the
Younger Dryas was not uniformly cold and that at least a number of warmer
spells occurred on eastern Svalbard.
We also conclude that the data on δ18O presented in Fig. 6 reflect
temperature variations at the coring site according to the isotopically
lighter ArW palaeotemperature model (Duplessy et al., 2005). Another
explanation for the heavier δ18O periods during the YD could be the
intermittent inflow of warmer AW; however, this is unlikely to cause the
synchronous disappearance of IRD.
IRD flux (upper scale, grey shading) and oxygen stable isotopes
records (bottom scale, black line) compared with oxygen stable isotope
records from ice core GISP2 from Greenland during the Younger Dryas period
(12 800–11 500 cal yr BP; Stuiver et al., 1995).
Glaciomarine unit I (early Holocene; 11 500–9200 cal yr BP)
During the early Holocene, foraminiferal fauna, although low in abundance,
were dominated by species related to the glaciomarine environment
(E. excavatum and C. reniforme; Fig. 5). Increasing species
richness and biodiversity of foraminifera point to amelioration of
environmental conditions and a progressive increase in the distance to the
glacier front (Korsun and Hald, 2000; Włodarska-Kowalczuk et al., 2013).
The decrease of the Fe / Ca ratio is suggested to reflect increased the
marine productivity and a reduced supply of terrigenous material (Croudace et
al., 2006). The mean grain size (> 63 µm; Fig. 4) indicates
weaker bottom currents at the beginning of the early Holocene and stronger
bottom currents at the end of this period, which might be related to the
ongoing isostatic uplift of the land masses of Svalbard as well as the sea
level rise (e.g. Forman et al., 2004; Taldenkova et al., 2012).
Significant fluctuations of δ18O and δ13C and increasing
abundance of N. labradorica and Islandiella spp. suggest
that Storfjordrenna was under the influence of various water masses at this
time (Fig. 5). Comparison of our δ18O records with records from the
Storfjorden shelf (400 m depth; Rasmussen et al., 2007; Fig. 1a) and the
northern shelf of Svalbard (400 m depth; Ślubowska et al., 2005;
Fig. 1b) shows that all of the records are shifted towards lighter values in
the early Holocene (Fig. 7a), and the record from our core shows the most
depletion (from ca. 13 000 cal yr BP). We suggest that the records
located on the western and northern shelf of Svalbard directly mirror the
effect of warmer Atlantic Water inflow, whereas records from Storfjordrenna
were under the influence of isotopically lighter Arctic Water from the
Barents Sea (Duplessy et al., 2005). The shift from the Arctic Water domain
to the Atlantic Water domain during the end of the early Holocene is also
visible on a scatter plot of δ13C against δ18O (Fig. 7b).
The results grouped to the left indicate Arctic Water domination, whereas the
results grouped to the right show Atlantic Water domination.
(a) The comparison of δ18O records (corrected for
ice volume changes) between Łącka et al. (this study; black solid line)
and Ślubowska et al. (2005; grey solid line) and Rasmussen et al. (2007;
black dashed line) plotted versus thousands of calendar years. The δ18O records after Łącka et al. (this study) were measured on
E. excavatum f. clavata and the two latter ones
(Ślubowska et al., 2005 and Rasmussen et al., 2007) were measured on
M. barleeanum. (b) Scatter plot showing δ13C
versus δ18O values from core JM09-020-GC (this study).
According to Kaufman et al. (2004), the early Holocene is characterised by
higher summer solar insolation at 60∘ N (10 % higher than today),
leading to a reduction in sea-ice cover (Sarnthein et al., 2003). As ice
cover decreased, additional solar energy was stored in summer and
subsequently re-radiated during the winter (e.g. Gildor and Tziperman,
2001). This process accelerated the ice sheet melting, and eventually, its
retreat towards the fjord heads (Forwick and Vorren, 2009; Jessen et
al., 2010; Baeten et al., 2010). Our data suggest that the iceberg calving to
Storfjordrenna was significantly reduced or may have even disappeared at
approximately 10 800 cal yr BP. However, the supply of turbid meltwater
from land to the study area still resulted in a relatively high sediment
accumulation rate.
According to Risebrobakken et al. (2011) and Groot et al. (2014), the
presence of Arctic Water suppressed the warming signal in the western Barents
Sea. This observation is in agreement with our data on planktonic
foraminifera reappearing at the termination of the early Holocene (ca.
9600 cal yr BP; Fig. 5). During this period, N. pachyderma (sin.)
dominated, but certain peaks of N. pachyderma (dex.) and
T. quinqueloba were noted. The two latter species are treated as
subpolar species (Bé and Tolderlund, 1971), although
T. quinqueloba also could be related to oceanic frontal conditions
separating Atlantic and Arctic Water (Johannessen et al., 1994; Matthiessen
et al., 2001). The peaks of T. quinqueloba near 9600 cal yr BP
were noted previously in the western Barents Sea margin (e.g. Hald et
al., 2007; Risebrobakken et al., 2010).
Increasing foraminiferal biodiversity in Storfjordrenna (Fig. 5) as well as
the occurrence of the thermophilous mollusc Mytilus edulis on the
western Edgeøya (Salvigsen et al., 1992) suggest that the inflow of AW
crossed Storfjordrenna and continued northward to the inner fjord by
9600 cal yr BP.
Glaciomarine unit II (mid-Holocene; 9200–3600 cal yr BP)
The mid-Holocene was characterised by relatively stable environmental
conditions, low sediment accumulation rates
(0.002 gcm-2yr-1) and a minor delivery of IRD (Fig. 4),
resulting from rather limited ice rafting and a reduced supply of
fine-grained material to Storfjordrenna. Low sedimentation rates and the low
Fe / Ca ratio reflect the reduced glacial conditions on Svalbard during the
mid-Holocene (Elverhøi et al., 1995; Svendsen and Mangerud, 1997). In
contrast, Hald et al. (2004) noted that in the record from Van Mijenfjorden,
an enhanced tidewater glaciation occurred during this period; it was thus
argued that IRD is a more reliable indicator of glaciation than sedimentation
rates. However, ice rafting in Storfjordrenna was generally low.
Shifts between the dominant species C. reniforme and
E. excavatum f. clavata (Fig. 5) reflect
environmental/hydrological changes (Hald and Korsun, 1997). The decrease of
E. excavatum f. clavata (percentage and flux), which
prefers colder bottom waters (Sejrup et al., 2004; Saher et al., 2009) and
the increase of C. reniforme point to the constant inflow of less
modified AW and a reduction in sedimentation (e.g. Schröder-Adams et
al., 1990; Bergsten, 1994; Jennings and Helgadóttir, 1994; Hald and
Steinsund, 1996; Hald and Korsun, 1997). Furthermore, the relative abundance
of M. barleeanum (Fig. 5) indicates that environmental conditions in
Storfjordrenna were similar to those of contemporary Norwegian fjords that
are dominated by AW with a temperature of 6–8 ∘C and salinities of
34–35 (Husum and Hald, 2004). High total foraminiferal flux at the beginning
of this period as well as high foraminiferal species richness and
biodiversity clearly point to AW conditions at the bottom (Hald and Korsun,
1997; Majewski and Zajączkowski, 2007; Włodarska-Kowalczuk et
al., 2013). These conclusions are also supported by the heavier δ18O, which demonstrates AW dominance and a significant reduction in the
amount of freshwater and ArW in Storfjordrenna (Fig. 7). The reduced sea-ice
condition during the mid-Holocene was also observed on the northern Barents
Sea continental margin, seen as an increase in palaeoproductivity (Wollenburg et
al., 2004). The continuous presence of Mytilus edulis during the
entire mid-Holocene points to the reduced inflow of the East Spitsbergen
Current due to the AW inflow (Feyling-Hansen, 1955; Forman, 1990; Salvigsen
et al., 1992. The pathway and range of AW inflow to the western and
northeastern Svalbard during mid-Holocene were well described by
Ślubowska-Woldengen et al. (2008) and Groot et al. (2014). Taken together
with our results, these observations suggest that one of the main pathways of
AW inflow to the eastern Svalbard may have occurred through Storfjordrenna.
Although sediment accumulation rates were low and grain size and geochemical
proxies remained relatively constant during the mid-Holocene, the
foraminiferal flux (including planktonic foraminifera) increased in two
periods of 9000–8000 and 6000–5500 cal yr BP (Figs. 4 and 5,
respectively). In both cases, the increase in IRD and I. norcrossi
fluxes was followed by a slight depletion in δ18O and heavier
δ13C, suggesting minor cooling and likely seasonal sea-ice
formation leading to beach sediment transport by shore ice. Our observations
support earlier studies of the overall mid-Holocene shifts towards a colder
environment (Skirbekk et al., 2010; Rasmussen et al., 2012; Berben et
al., 2014; Groot et al., 2014; Sternal et al., 2014) and fluctuations in the
glacial activity in the Svalbard region (e.g. Forwick and Vorren, 2007, 2009;
Beaten et al., 2010; Ojala et al., 2014). Our data show an increased supply
of IRD fraction to the Storfjordrenna sediment followed by variation of
δ18O; however, the high flux of M. barleeanum associated
with Atlantic-derived waters (Steinsund, 1994; Jennings et al., 2006; Fig. 5)
indicates an AW condition in southern Storfjorden throughout the entire
mid-Holocene. A similar ameliorated condition with consistent AW inflow also
prevailed over the mid-Holocene in the Kveithola Trough south of
Storfjordrenna (Berben et al., 2014; Groot et al., 2014). To a lesser extent,
these two signals (AW inflow and higher IRD flux) are not necessarily
contradictory because snow accumulation on land and inconsiderable glacier
advance depend on humid air transport from the ocean. Thus, slight changes in
the atmospheric frontal zone over Svalbard could cause fluctuation of the
glacier range.
Glaciomarine unit III (late Holocene; 3600–1200 cal yr BP)
The late Holocene is characterised by a gradual increase in sediment
accumulation rates followed by numerous sharp peaks of sand content and minor
peaks of IRD flux as well as an increased Fe / Ca ratio, thus indicating
ice growth on land (compare with e.g. Svendsen and Mangerud, 1997; Hald et
al., 2004; Forwick and Vorren, 2009; Taldenkova et al., 2012; Kempf et
al., 2013) and slightly enhanced iceberg calving and/or ice rafting over the
core site. The IRD record shows few irregular small peaks in the late
Holocene (Fig. 6), which could be correlated with enhanced sea currents that
increase the drift of the icebergs, according to Hass (2002). Forwick et
al. (2010) suggested several glacier front fluctuations during the past two
millennia in Sassenfjorden and Tempelfjorden (W Spitsbergen), and hence we
assume that increased iceberg calving occurred at Storfjordrenna during this
time. However, increased IRD flux can also reflect deposition related to
enhanced shore ice rafting. The latter explanation is in agreement with the
heavier δ18O record (Fig. 5), indicating a minor cooling.
The mean grain size (> 63 µm) increases in the late Holocene
(Fig. 4) and may indicate stronger bottom current velocities and winnowing
of fine-grained sediments. Andruleit et al. (1996) observed similar
increased erosive activity of bottom currents during the late Holocene on
the SW Svalbard shelf. This sudden increase in current velocities might be
connected with (1) postglacial reorganisation of oceanographic conditions,
(2) relative lowering of the sea level during the postglacial isostatic
rebound and/or (3) more intensive sea-ice formation that enhanced the
formation of BSW, thus forming a seasonal near-bottom dense water mass
flowing over the coring site (Andruleit et al., 1996). Nevertheless, this
process is still not fully understood.
The sharp increase in the foraminiferal flux (Fig. 4) pointing to the
increased nutrient advection/upwelling and biological productivity at the
coring site during the late Holocene was likely caused by variable
hydrological conditions and most likely strong gradients leading to the
formation of hydrological fronts. In contrast, Wollenburg et al. (2004) noted
reduced palaeoproductivity in the northern Barents Sea over the entire late
Holocene, pointing to several events of heavy sea-ice cover. Our data show
increased fluxes of opportunistic species E. excavatum and
C. reniforme as well as an abundance of N. labradorica and
Islandiella spp. N. labradorica and Islandiella
spp. in areas with a high biological productivity in the upper surface waters
(e.g. Hald and Steinsund, 1996; Korsun and Hald, 2000; Knudsen et
al., 2012). Abundant though variable M. barleeanum is documented in
organic-rich mud within troughs of the Barents Sea (Hald and Steinsund, 1996)
and in temperate fjords of Norway (Husum and Hald, 2004), which points to
high productivity in the euphotic zone leading to enhanced export of organic
material/nutrients to the sea floor. Our data also show high
N. pachyderma flux throughout this unit, reflecting a significant
increase of euphotic productivity at the coring site. However, a low
percentage of dextral specimens and T. quinqueloba point to low
sea-surface temperatures (Fig. 5). This observation is in agreement with
Rasmussen et al. (2014), who noted that after ca. 3700 cal yr BP, Atlantic
Water was only sporadically present at the surface. Cooling at the sea
surface reflects the general trend in the Northern Hemisphere related to
orbital forcing and reduction of summer insolation at high latitudes over the
late Holocene (Wanner et al., 2008).
The last evidence of AW inflow to Edgøya area based on M. edulis
is dated to 5000 cal yr BP (Hjort et al., 1995). After that time,
M. edulis remained absent until the present time; however; its
disappearance could be related to the freshening of Surface Water (Berge at
al., 2006) and sea-ice forcing as opposed to the extinction of AW in
Storfjorden over the late Holocene (Rasmussen et al., 2007).