Evaluation of past stratification changes in the Nordic Seas by comparing planktonic foraminiferal δ18O with a solar-forced model

Marine Micropaleontology 94–95 (2012) 91–96

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Evaluation of past stratification changes in the Nordic Seas by comparing planktonic foraminiferal δ 18O with a solar-forced model Johannes Simstich a,⁎, Stephan J. Lorenz b, Henning A. Bauch c a b c

Institute for Relevant Research (IRR), 58313 Herdecke, Germany Max-Planck-Institute for Meteorology, 20146 Hamburg, Germany Academy of Sciences, Humanities, and Literature, Mainz; c/o GEOMAR, Helmholtz Center for Ocean Research, 24148 Kiel, Germany

a r t i c l e

i n f o

Article history: Received 19 March 2011 Accepted 22 June 2012 Keywords: Planktonic foraminifera Nordic Seas Stable isotopes Stratification

a b s t r a c t Density changes in the upper water column of the northern North Atlantic may enhance or reduce vertical convection of surface water with profound effects on meridional overturning and climate in the wider region. This study tests the capability of paired δ18O values of two planktonic foraminiferal species – Neogloboquadrina pachyderma (s) and Turborotalita quinqueloba – for the reconstruction of near-surface density stratification in high latitudes or the glacial ocean. Foraminiferal data from two sediment cores of crucial areas of the Nordic Seas were compared with insolation-induced thermal stratification changes as obtained by simulations with the general circulation model ECHO-G. The comparison suggests that insolation was the chief mechanism to change thermocline strength during most of the Holocene. Prior to that, stratification depended by and large on the varying amounts of meltwater injected at the sea surface. Similar to the modern central Arctic Ocean, a pronounced and thick halocline prevented surface waters from deep convection in the central Nordic Seas. Parts of the Norwegian Sea, however, were also stratified but more analogous to the modern Greenland Sea, where deep convection can occur in late winter as a result of the density increase upon a combination of cold temperatures and wind stress. Our findings thus support previous results of an active meridional overturning also in a glacial ocean. © 2012 Elsevier B.V. All rights reserved.

1. Introduction The Nordic Seas, i.e. the Greenland, Iceland, and Norwegian seas (Fig. 1) define one of the key areas in the global ocean, where convective sinking of surface waters contributes to the renewal of deep and intermediate waters. For deep convection to occur, the thermohaline stratification of the upper 50–100 m of the water column must allow the waters at the surface to become more dense than the underlying waters. Density reductions in the surface layer of the Nordic Seas are regarded to lead to reduced deep water formation and disturbance of the global thermohaline circulation system and climate in the North Atlantic region (Rahmstorf et al., 2005). Most paleoceanographic reconstructions of such circulation changes in the deep ocean focus on one layer. For instance, benthic records show variations in the intensity of deep water formation in the Nordic Seas since the last glaciation (Bauch et al., 2001). Planktonic records from the Last Glacial Maximum (LGM) indicate highly variable temperatures and salinities in the near surface layer (de Vernal et al., 2006), and in parts densities suitable for deep water formation (Weinelt et al., 1996). However, there is a lack of knowledge about density differences between the sea surface and the layer directly underneath the pycnocline. It is ⁎ Corresponding author at: 58313 Herdecke, Germany. Tel./fax: +49 2330 807190. E-mail address: [email protected] (J. Simstich). 0377-8398/$ – see front matter © 2012 Elsevier B.V. All rights reserved. doi:10.1016/j.marmicro.2012.06.006

therefore unclear whether the stratification provided conditions for deep convection or not. In our study, we use the difference between the oxygen isotopic composition of calcite tests of two planktonic foraminifera – Neogloboquadrina pachyderma (sinistral coiling) and Turborotalita quinqueloba – as proxy for temperature and salinity gradients in the upper 100 m of the water column through the last 30,000 years for two sites representing different hydrographical regimes in the Nordic Seas (Fig. 1). As an indicator for the presence of meltwater we employ ice rafted debris (IRD). We then compare the proxy data with data from a general circulation model for atmosphere and ocean, which is forced by solar insolation in order to test the influence of insolation for ocean stratification at high-northern latitudes. 1.1. Rationale of isotope differences in the foraminifera Despite some differences in water mass preference, N. pachyderma (s) and T. quinqueloba are ubiquitous in the modern Nordic Seas (Bé, 1977), and even co-occur below sea ice (e.g. Carstens and Wefer, 1992). Sediment trap studies show that – in high latitudes – both species cover the same season, i.e. intermediate and high abundances occur from June to October (e.g. Donner and Wefer, 1994; von Gyldenfeldt et al., 2000; Schroeder-Ritzrau et al., 2001; Kuroyanagi et al., 2002; Jonkers et al., 2010). However, deep convection occurs

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After reproduction, all dead foraminifera setting to the sea floor bear a cumulative isotope signal from their habitat depth range of calcite growth. Isotope analysis of a sediment sample merely provides a value, which reflects a single distinct, but only ‘apparent’ calcification depth. Modern apparent calcification depths reconstructed from dated sediment surface samples on an E–W profile across the Nordic Seas near 67–70°N on average amount to 117 m (1σ: ±73 m, max.= 258 m, min.= 17 m, n = 54) for N. pachyderma (s) and 47 m (1σ: ±18 m, max.= 94 m, min.=27 m, n = 44) for T. quinqueloba (Simstich et al., 2003).

Fram Strait

Greenland

70

o

N

er

Greenland Sea

m

m

Su

2.1. The sediment records

PS1243 Iceland Sea

M23323 Norwegian Sea

0 oE

60 o N

20 o W

Iceland

2. Material and methods

1 0 oW

65 o N

Winter <0.1‰ Δδ18O

Sea ice

Norway

Fig. 1. The Nordic Seas, core positions (asterisks), surface currents (arrows), and average sea ice margins in summer and winter (stippled lines). Area of low Δδ18Oforam b0.1‰ in sediment surface samples to the northwest of the thick black line (0.2‰ added to δ18O of T. quinqueloba to account for vital effects, Simstich et al., 2003).

mainly in winter (Killworth, 1983). Therefore, our approach does not allow us to reconstruct deep convection directly, but to distinguish between different hydrographic situations, which may, or may not, be suitable for deep convection to occur. Plankton tows show that N. pachyderma (s) and T. quinqueloba inhabit the same depth intervals between sea surface and ~ 250 m, with preference to the mixed layer above the seasonal thermocline b25– 50 m (Arikawa, 1983; Bock, 1990; Simstich et al., 2003; Kuroyanagi and Kawahata, 2004). In areas with a cold halocline on top, like in the western Greenland Sea and Fram Strait, or the southern Nansen Basin (Arctic Ocean), both species prefer the warmer and more salty Atlantic-derived water > 50–100 m (Carstens and Wefer, 1992; Carstens et al., 1997; Volkmann, 2000). However, both species are restricted in their maximum habitat depth and, therefore, live within the halocline if its thickness exceeds 150 m as in the central Arctic Ocean (Carstens and Wefer, 1992). For reproduction, these foraminifera settle in the colder, subthermocline waters. Here, N. pachyderma (s) adds a thick secondary calcite crust to its primary test, as indicated by abundance maxima of encrusted N. pachyderma (s) below the thermocline (Arikawa, 1983; Berberich, 1996; Kohfeld et al., 1996; Simstich et al., 2003) or in the deeper halocline (Carstens and Wefer, 1992; Volkmann, 2000). Based on shell thickness (Arikawa, 1983) and shell weights (Kohfeld et al., 1996) the crust is assumed to make up roughly 50% of the total shell mass. T. quinqueloba, in contrast, is hardly affected by encrustation (Arikawa, 1983). The differences in calcification behavior introduce differences in the isotopic composition of the foraminiferal tests. The 18O/ 16O ratios (versus a standard as δ 18O, with δ = [(Rsample/Rstandard) − 1] * 1000, and R = 18O/ 16O) in the CaCO3 tests (δ 18Oforam) mainly depend on temperature and δ 18O of the ambient water (δ 18Owater) during test formation. Crust formation of N. pachyderma (s) in colder and more saline water (i.e. higher δ 18Owater) below the thermocline or in the deep halocline leads to enrichment of the heavier isotope 18O in the final, encrusted tests (Simstich et al., 2003). Additional species specific biological fractionation can be corrected for by adding 0.2‰ to δ 18O of T. quinqueloba (Simstich et al., 2003), assuming that this factor remained unchanged within the time frame of our study.

Core PS1243 (69.38°N, 6.51°W, 2721 m water depth) is located in the northeastern Iceland Sea (Fig. 1) close to the modern Arctic Front. The core was sampled in 1 and 2 cm intervals, respectively, and as 1 cm-thick slices (further information in Bauch et al., 2001). In the Norwegian Sea, core M23323 (67.77°N, 5.92°E, 1286 m) was obtained from the top of a small mound on the Vøring Plateau, underneath the Norwegian Atlantic Current. This core was sampled in 2 cm thick slices taken in 2–8 cm intervals. All samples were wet sieved (>63 μm) and rinsed with de-ionized water. Approximately 30 specimens of N. pachyderma (s) and 60 T. quinqueloba were picked from the size fraction 125– 250 μm. They were cracked open, cleaned with methanol in an ultrasonic bath and dried at 40 °C. Oxygen isotope compositions of the tests were measured at the Leibniz-Laboratory at University of Kiel on a Finnigan MAT 251 mass spectrometer with an automated Carbo-Kiel preparation line (Kiel Device I) versus the Pedee Belemnite (PDB) scale established via the NBS 20 (National Bureau of Standards) carbonate stable isotope standard. The analytical error for δ18O is ±0.07‰. Data is available via the PANGAEA data base. Stratigraphy of core PS1243 is based on 13 radiocarbon dates and the Vedde ash, which is chemically identified by Wallrabe-Adams & Lackschewitz (Wallrabe-Adams and Lackschewitz, 2003). Stratigraphic details, sedimentation rates and age/depth plot are published previously (Bauch et al., 2001). The age model of core M23323 is based on seven radiocarbon dates, a light ash, assumed to be the Vedde ash, and linear interpolation between dated fix points (Table 1, Fig. 2). For dating, foraminifera samples were taken from the abundance maxima of N. pachyderma (s) and analysed with the accelerator mass spectrometry (AMS) facility at Kiel University. Conversion from radiocarbon years before present ( 14C yr BP) to calendar years before present (cal. yr BP) was performed with CALIB 5.0.1 (Stuiver and Reimer, 1993) using marine04 (Hughen et al., 2004) as calibration data set for foraminifera and intcal04.14c (Reimer et al., 2004) for the Vedde ash (Birks et al., 1996). The global average was taken for reservoir age correction, as given by the programme (Stuiver et al., 1998). Radiocarbon ages are also available from the PANGAEA data base. Ages are reported as thousand calendar years before present (cal. ka BP; Table 1). 2.2. The model ECHO-G In order to understand the physical mechanisms responsible for the isotope variations, we utilized simulations with the coupled atmosphere-ocean general circulation model ECHO-G (Lorenz et al., 2006). The atmospheric part of ECHO-G is the general circulation model ECHAM4 with a resolution of ~2.8° × 2.8°. ECHAM4 is coupled to the HOPE ocean model including a dynamic–thermodynamic sea-ice model. ECHO-G was adapted to account for variations in the annual distribution of solar radiation resulting from the varying orbital parameters (Fig. 3A, Berger, 1978). The timescale of the model was shortened by an acceleration factor of 100, thus, the last 30,000 years are presented in only 300 integration years. Due to the acceleration

J. Simstich et al. / Marine Micropaleontology 94–95 (2012) 91–96

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Table 1 Radiocarbon dates from core M23323 and the Vedde ash (Birks et al., 1996). Mean ages printed in bold letters were chosen as fix points for linear age interpolation. Depth (cm)

Species

Laboratory code

Age ±1σ (14C yr BP)

53

N. pachyderma (s)

KIA2517

3550 ± 50

161

N. pachyderma (s)

KIA2518

7970 ± 50

181

Vedde ash

211

N. pachyderma (s)

KIA2519

12,440 ± 60

223

N. pachyderma (s)

KIA2520

14,540 ± 70

233

N. pachyderma (s)

KIA2521

15,800 ± 80

263

N. pachyderma (s)

KIA2522

16,950 ± 90

295

N. pachyderma (s)

KIA2523

18,850 ± 145

10,310 ± 50

technique, the simulation is too fast for the deep ocean to respond. However, the uppermost 100 m – the focus of our study – are well represented in the model (Lorenz and Lohmann, 2004). Three experiments for the last 30 ka were performed. Throughout the simulations, the greenhouse gas concentrations were fixed; for example, carbon dioxide is set to 280 ppm. Moreover, modern values for vegetation, sea level, and continental ice were used, i.e. no glacial characteristics for continental or sea ice were applied. The only driving mechanism for the model was the seasonally changing solar insolation due to the moving parameters of the Earth's orbit in space (Berger, 1978; Lorenz et al., 2006). The June–October mean of the daily insolation was at a minimum at ~ 21 ka during the LGM, and reached the Holocene maximum between 8 and 9 ka (Fig. 3A). The resulting temperature and salinity curves (Fig. 3B, C), therefore, show the influence insolation had on the thermohaline stratification of the upper water column near the core positions during the last 30 ka. 3. Results

1σ 2σ 1σ 2σ 1σ 1σ 1σ 2σ 2σ 2σ 1σ 2σ 1σ 2σ 1σ 2σ 1σ 1σ 2σ 1σ 2σ

Age ranges (cal. yr BP)

Relative area under distribution

Mean age (cal. yr BP)

3369–3495 3327–3568 8373–8482 8336–8542 11,993–12,172 12,206–12,234 12,325–12,339 11,840–11,862 11,968–12,254 12,256–12,381 13,816–13,958 13,764–14,027 16,653–17,069 16,423–17,265 18,684–18,816 18,615–18,873 19,563–19,730 19,741–19,807 19,496–19,874 21,776–22,246 21,391–22,328

1 1 1 1 0.85 0.10 0.05 0.02 0.80 0.18 1 1 1 1 1 1 0.7 0.3 1 1 1

3430 3450 8430 8440 12,080 12,220 12,332 11,850 12,110 12,320 13,890 13,900 16,860 16,840 18,750 18,740 19,650 19,770 19,680 22,010 21,860

PS1243 (Iceland Sea) varies between 0.5 and 1.1‰. In contrast, the Δδ 18O of core M23323 (Norwegian Sea) remains close to 0‰ in the early LGM, not exceeding 0.5‰.

3.2. Modelling results The model results show a partly inverse relationship between temperatures at sea surface and 100 m water depth (Fig. 3B). Due to insolation, sea surface temperatures increased by 1–2 °C from glacial values to a maximum around 11 ka and decreased by 0.5–1 °C during the mid Holocene, remaining on that level to the present. Conversely, in 100 m water depth, temperatures decreased by 1–1.5 °C from glacial values to lower values during the deglaciation and the early Holocene, and increased by 0.5–1 °C afterwards. Thus, the thermal gradients were lowest during the glaciation before 20 ka with b1 °C/100 m in the Iceland Sea and ~ 3 °C/100 m in the Norwegian Sea (Fig. 3B). The gradients were highest (> 2 °C/100 m and ~ 5 °C/100 m, respectively) around 11 ka and decreased to 2 °C/100 m and 4 °C/100 m, respectively, during the Holocene.

3.1. Foraminiferal oxygen isotopes

35

300

25 200 20 150 15 100 10 50

0 0

Sedimentation rate (cm/kyr)

30

250

Depth in core (cm)

Except for depth intervals in the Younger Dryas at 12–13 ka and at the beginning of Heinrich event 1 (H1) around 17–18.5 ka in core PS1243, and in the Bølling-Allerød and H1 between 13 and 18 ka in M23323, there were sufficient T. quinqueloba tests available in the sediment samples for isotope analyses. In both cores, nearly all the δ 18O values for N. pachyderma (s) are higher than for T. quinqueloba (Fig. 4A, D). As expected, both species typically show higher glacial than Holocene values. Glacial to interglacial changes were lower in the Iceland Sea than in the Norwegian Sea due to higher Holocene δ 18O values found in the latter. However, even though the trends in the δ 18O records of N. pachyderma (s) and T. quinqueloba are the same, the differences between both isotope signals (Δδ18Oforam =δ18ON. pachyderma (s) – (δ18OT. quinqueloba + 0.2‰)) are not constant throughout the cores (Fig. 4B, E). Both cores show highest Δδ 18O values of ~ 1.5‰ around the Younger Dryas interval (11.6–12.9 ka). During the Holocene, the Δδ 18O isotope differences decrease to ~ 0.5‰. In the Last Glacial Maximum (LGM, 18.0–21.5 ka), the δ 18O differences between N. pachyderma (s) and T. quinqueloba increase in the Iceland Sea, but nearly disappear in the Norwegian Sea – i.e., the Δδ 18O in core

5

5

10

15

20

0 25

Age (cal. kyr BP) Fig. 2. Age–depth relation (black line) and linear sedimentation rates (stippled line) of core M23323.

J. Simstich et al. / Marine Micropaleontology 94–95 (2012) 91–96

Insolation (W/m2)

94

340

Salinity at sea surface ran parallel to salinity in 100 m water depth (Fig. 3 C), thus indicating that insolation had a small effect on salinity, but did not change its gradients.

A

320 300

4. Discussion

280

4.1. Comparing measured and model estimated changes in δ 18O

B

8

Temperature and salinity obtained by the models can be translated in δ 18O values as theoretically expected in foraminiferal calcite precipitating in equilibrium with temperature (T) and δ 18Owater. The latter is estimated from salinity (S) using the regional salinity: δ18Owater relationship for the central and eastern Nordic Seas (Simstich et al., 2003, based on unpublished data of H. Erlenkeuser):

Temperature (oC)

NS 10m

6

IS 10m

4


 2 r ¼ 0:72; n ¼ 346

δ Owater ¼ −12:17 þ 0:36S 18

NS 100m

2

ð1Þ

IS 100m

Equilibrium calcite (δ 18Oexp) is then calculated from the empirical equation for the 18O/ 16O fractionation during inorganic calcite precipitation (Shackleton, 1974):

0

IS 10m IS 100m

5

10

15

20

25

30

18 δw ¼ 0:99973δ Owater  0:27‰

Age (cal. ka BP) Fig. 3. A) Mean daily insolation June–October at 68°N and resulting B) temperature and C) salinity changes in the model (11-year moving averages) at sea surface (10 m resp.) and 100 m water depth. Gray shading circumscribes maxima and minima of the three simulated time series; black lines show means. Every time series averages data from June to October within rectangles of 4 × 4 data points around the core positions (IS: Iceland Sea — PS1243: 11.25–2.81°W, 66.97–71.16°N; NS: Norwegian Sea — M23323: 2.81–11.25°E, 65.58–69.76°N).

The differences between δ 18Oexp in 100 m and at sea surface (Δδ18Oexp =δ18O100m −δ18Osea surface) mirror the temperature gradients as described above. Δδ18Oexp indicate how much of the Δδ18Oforam changes can be explained by insolation changes directly. Our strategy of comparing δ18O differences in data and model avoids complications

Age (cal. ka BP) 0

δ18O (‰ PDB)

1

10

15 YD

H1

20

Age (cal. ka BP) 25

30

0

5

LGM

10

15 YD

H1

20

25

LGM

30

D

A

2

3

3

1.5 1

PS1243 Iceland Sea

M23323 Norwegian Sea

4

B

E

0.5

0

300

1.5 1

0.5

400 IRD (grains/g)

1

2

4

Δδ18O (‰ PDB)

5

ð3Þ

0 15

C

F 10

200 5

100 0

δ18O (‰ PDB)

0

with (Hut, 1987):

18

35

ð2Þ

Δδ O (‰ PDB)

35.5

h i 18 0:5 δ Oexp ¼ 21:9  3:16ð31:061 þ TÞ þ δw

NS 10m NS 100m

0

5

10 15 20 Age (cal. ka BP)

25

30

0

5

10 15 20 Age (cal. ka BP)

25

Grains >2mm % >125 μm

Salinity

C

0 30

Fig. 4. Data of core PS1243: A) δ18O of T. quinqueloba (asterisks) and N. pachyderma (s) (squares); B) Δδ18Oforam (triangles) and its cubic spline (thick gray line), and Δδ18Oexp calculated from results of model ECHO-G (thick dashed line); C) ice rafted debris (IRD) as number of grains >0.250 mm per gramme dry sediment. D)–F) As in A)–C) but for core M23323; F) IRD as number of grains >2 mm (black line) and sand >0.125 mm in % of dry weight (dashed line). Gray bars indicate time intervals: YD: Younger Dryas; H1: Heinrich event 1; LGM: Last Glacial Maximum. 0.2‰ added to δ18O of T. quinqueloba to account for vital effects (Simstich et al., 2003).

J. Simstich et al. / Marine Micropaleontology 94–95 (2012) 91–96

due to large scale or whole ocean δ18O changes such as those associated with sea level changes.

95

surface. Such a strong halocline would have very likely suppressed deeper-going (down to sea bottom) convection in the glacial Iceland Sea.

4.2. Insolation as driving force for stratification? 4.4. Glacial stratification in the Norwegian Sea In the Holocene, after 7 ka in the Iceland Sea and after 10 ka in the Norwegian Sea, the spline Δδ 18Oforam curves run roughly parallel and below the Δδ 18Oexp curves from ECHO-G (Fig. 4B, E), thus indicating that changes of the δ 18O differences between the apparent calcification depths of N. pachyderma (s) and T. quinqueloba can be explained by changes of the temperature gradient due to insolation. In the period before 7 ka in the Iceland Sea, the spline Δδ 18Oforam curves exceeded Δδ 18Oexp (Fig. 4B), thus showing that the density gradients across the pycnocline were stronger than induced by insolation alone. If explained only by temperature, the 1‰ excess of Δδ 18Oforam over Δδ 18Oexp prior to 23 ka in the Iceland Sea would require an additional ~ 4 °C difference across the seasonal thermocline (Eqs. (1)–(3)). Cooling of subthermocline water in that magnitude would bring it below freezing (Fig. 3B). Exclusive heating of the upper water column due to lateral advection of warm Atlantic Water also seems unlikely because the Atlantic Water layer is thick enough (600 m today, i.e. sill depth of Iceland–Scotland Ridge) to affect the habitats of both foraminifera above and below the seasonal thermocline simultaneously. 4.3. Meltwater influence and glacial stratification in the Iceland Sea Sediments deposited before 10–11 ka are rich in ice-rafted debris (Fig. 4C, F), which indicates the former presence of melting icebergs and/or sea ice on the sea surface. Thus, the Δδ 18Oforam excess probably resulted from a lowering of δ 18Owater due to the supply of 18Odepleted meltwater. Disregarding potential nonlinearities (Rohling and Bigg, 1998), quantifications of the salinity reductions connected with the δ 18Owater lowering relate to the slope of the relationship between salinity and δ 18Owater. Today, S:δ 18Owater equals 1:0.36 in most areas of the Nordic Seas (Eq. (1)), and 1:0.5 near the Greenland coast (Simstich et al., 2003) and in the North Atlantic as revealed by GEOSECS data; in the Fram Strait and the Polynya off NE Greenland, the slope is 1:1.1 (Kohfeld et al., 1996). The true slope is not yet determined for the last glaciation and deglaciation (Schaefer-Neth and Paul, 2001). On the basis of those three slopes, the 0.3‰ excess of average Δδ 18Oforam over Δδ 18Oexp in the LGM of the Iceland Sea (Fig. 4B) could result from a salinity difference of 0.3–1 between the apparent calcification depths of N. pachyderma (s) and T. quinqueloba. This exceeds by far its modern equivalent (0.05, World Ocean Atlas 2001), and indicates that the stratification structure during the LGM was largely different from today's. At present, warm Atlantic Water with salinities of 35.1–35.5 plays a major role in the central and eastern Nordic Seas. Although Atlantic water mass inflow was still active during the LGM (e.g. Bauch, 1994; Hebbeln et al., 1994), its salinity was probably higher by 1 unit compared with today (Fairbanks, 1989). In contrast, reconstructions based on foraminiferal δ 18O (~35; Duplessy et al., 1991) and dinocyst transfer functions (~34; de Vernal et al., 2005) suggest that sea surface salinities in the Iceland Sea were lower than today. Thus, the salinity difference between the Atlantic derived water and the meltwater on top may have been between ~1 and 2.5. This implies that the sea surface salinity reduction was 1–8 times higher than estimated above from insolation induced stratification changes in the model and the Δδ 18Oforam. This finding also suggests that the upper water column of the glacial Iceland Sea was similarly stratified as the modern central Arctic Ocean (c.f. Bauch et al., 2001; de Vernal et al., 2006; Rasmussen and Thomsen, 2008; Rasmussen and Thomsen, 2009; see also Rasmussen et al., 2007). Under such conditions water of Atlantic origin is covered by Polar Water forming a massive halocline of ~200 m thickness in which salinity decreases continuously, sometimes exponentially, toward sea

Meltwater also dominated the surface of the Norwegian Sea during LGM, as shown by coarse grains in core M23323 (Fig. 4F). However, low Δδ 18Oforam values (Fig. 4E) indicate a different stratification structure than in the Iceland Sea. Modern analogues with similarly low Δδ 18Oforam are found in sediment surface samples from the northwestern Iceland and Greenland seas, and western Fram Strait (Fig. 1). Here, the Polar Water on top of the warmer, more saline water of Atlantic origin is thinner (50–100 m) than in the central Arctic Ocean. In addition, enhanced zooplankton abundances are found in the highly productive marginal ice zone (Ramseier et al., 2001), which retreats during spring and summer from its maximum to its minimum extent near the Greenland coast (Fig. 1). Plankton tows show that living and encrustation of N. pachyderma (s) in such areas happen in the same depth range within the Atlantic Water where also highest abundances of living T. quinqueloba are found (e.g. Carstens et al., 1997; Volkmann, 2000) – the two species therefore also adopt similar δ 18O values here. As a result, the area seasonally traversed by the ice margin is also the area of low Δδ 18Oforam (Fig. 1). The finding shows that the Norwegian Sea was at least partly stratified like parts of the modern Greenland Sea, where deep convection can occur in late winter. The requirements for deep convection (Killworth, 1983) may have been met in the eastern Norwegian Sea in the early LGM where katabatic winds from the Fennoscandian ice sheet could have maintained polynyas, thus facilitating surface cooling, sea ice formation with brine release, and mixing with the dense Atlantic water from below the meltwater layer (see also Bauch et al., 2001). Deep water formation in the Nordic Seas during the LGM cannot be excluded as proposed earlier (see Table 1 in Sarnthein et al., 2001). 5. Conclusions Paired δ 18O values of N. pachyderma (s) and T. quinqueloba were tested in their capability for the reconstruction of glacial–interglacial hydrographic changes in high-northern latitudes. We compare foraminiferal data with insolation induced thermal stratification changes as obtained by simulations with the general circulation model ECHO-G. The comparison suggests a linkage between oceanic stratification and insolation during the Holocene, but a completely different – meltwater driven – stratification during the last glacial time. Similar to the modern central Arctic Ocean, a strong halocline prevented surface waters from deep convection in the central Nordic Seas. Parts of the Norwegian Sea, however, were probably stratified like parts of the modern Greenland Sea, where deep convection can occur in late winter. Acknowledgments We thank H. Erlenkeuser for the unpublished data and generous advice, H.-H. Cordt and H. Heckt for isotope measuring, J. Rumohr and F. Blaume for samples and unpublished data. We thank M. Sarnthein for his initiative, U. Ninneman and an anonymous reviewer for their helpful suggestions. References Arikawa, R., 1983. Distribution and taxonomy of Globigerina pachyderma (Ehrenberg) off the Sanriku Coast, Northeast Honshu, Japan. Tohoku University Scientific Reports, Serie 2 (Geology) 53 (2), 103–157. Bauch, H., 1994. Significance of variability in Turborotalita quinqueloba (Natland) test size and and abundance for paleoceanographic interpretations in the Norwegian– Greenland Sea. Marine Geology 121, 129–141.

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