# An ocean-ice coupled response during the last glacial: a view from a marine isotopic stage 3 record south of the Faeroe Shetland Gateway

J. Zumaque, F. Eynaud, S. Zaragosi, F. Marret, K. Matsuzaki, C. Kissel, Didier M. Roche, B. Malaize, E. Michel, I. Billy, et al.

► To cite this version:

J. Zumaque, F. Eynaud, S. Zaragosi, F. Marret, K. Matsuzaki, et al. An ocean-ice coupled response during the last glacial: a view from a marine isotopic stage 3 record south of the Faeroe Shetland Gateway. Climate of the Past, 2012, 8 (6), pp.1997-2017. 10.5194/cp-8-1997-2012. hal-02913959

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# An ocean-ice coupled response during the last glacial: a view from a marine isotopic stage 3 record south of the Faeroe Shetland Gateway

J. Zumaque¹*, F. Eynaud¹, S. Zaragosi¹, F. Marret², K. M. Matsuzaki¹,**, C. Kissel³, D. M. Roche³,⁴, B. Malaizé¹, E. Michel³, I. Billy¹, T. Richter⁵, and E. Palis¹

¹EPOC – Environnements et Paléoenvironnements Océaniques, UMR5805, Université Bordeaux I, Talence, France

²School of Environmental Sciences, University of Liverpool, Liverpool, L69 7ZT, UK

³Laboratoire des Sciences du Climat et de l'Environnement (LSCE), CEA/CNRS-INSU/UVSQ – UMR8212, 91191 Gif s/s Yvette, France

⁴Vrije Universiteit Amsterdam, de Boelelaan 1085, 1081 HV Amsterdam, The Netherlands

⁵NIOZ Royal Netherlands Institute for Sea Research, P.O. Box 59, 1790 AB Den Burg (Texel), The Netherlands

*present address: GEOTOP, Montréal, Canada

**present address: Institute of Geology and Paleontology, Graduate School of Science, Tohoku University, Aramaki 6-3, Sendai, 980-8578, Japan

*Correspondence to:* F. Eynaud (f.eynaud@epoc.u-bordeaux1.fr)

Received: 17 July 2012 – Published in Clim. Past Discuss.: 1 August 2012

Revised: 5 November 2012 – Accepted: 14 November 2012 – Published: 7 December 2012

**Abstract.** The rapid climatic variability characterising the Marine Isotopic Stage (MIS) 3 (~60–30 cal ka BP) provides key issues to understand the atmosphere-ocean-cryosphere dynamics. Here we investigate the response of sea-surface paleoenvironments to the MIS3 climatic variability through the study of a high resolution oceanic sedimentological archive (core MD99-2281, 60°21′N; 09°27′W; 1197 m water depth), retrieved during the MD114-IMAGES (International Marine Global Change Study) cruise from the southern part of the Faeroe Bank. This sector was under the proximal influence of European ice sheets (Fennoscandian Ice Sheet to the East, British Irish Ice Sheet to the South) during the last glacial and thus probably responded to the MIS3 pulsed climatic changes.

We conducted a multi-proxy analysis of core MD99-2281, including magnetic properties, x-ray fluorescence measurements, characterisation of the coarse (> 150 µm) lithic fraction (grain concentration) and the analysis of selected biogenic proxies (assemblages and stable isotope ratio of calcareous planktonic foraminifera, dinoflagellate cyst – e.g. dinocyst – assemblages). Results presented here are focussed on the dinocyst response, this proxy providing the reconstruction of past sea-surface hydrological conditions, qualitatively as well as quantitatively (e.g. transfer function

sensu lato). Our study documents a very coherent and sensitive oceanic response to the MIS3 rapid climatic variability: strong fluctuations, matching those ofstadial/interstadial climatic oscillations as depicted by Greenland ice cores, are recorded in the MD99-2281 archive. Proxies of terrige neous and detrital material suggest increases in continental advection during Greenland Stadials (including Heinrich events), the latter corresponding also to southward migra tions of polar waters. At the opposite, milder sea-surface con ditions seem to develop during Greenland Interstadials. After 30 ka, reconstructed paleohydrological conditions evidence strong shifts in SST: this increasing variability seems con sistent with the hypothesised coalescence of the British and Fennoscandian ice sheets at that time, which could have di rectly influenced sea-surface environments in the vicinity of core MD99-2281.

## 1 Introduction

The alternation of glacial/interglacial periods that defines the Quaternary is thought to be mainly triggered by changes in orbital parameters (e.g. Imbrie et al., 1993; Berger, 2006). However, a higher frequency in climatic variability is
detected during glacial periods and appears to have no direct
orbital forcing origin. This millennial-scale climatic variabil-
ity was observed for the first time in the oxygen isotope
records from Greenland ice cores (Dansgaard et al., 1993)
and is expressed through significant and rapid changes in at-
mospheric temperatures, especially during the Marine Iso-
tope Stage (MIS) 3 (e.g. Dansgaard et al., 1993; Kiefer et al.,
2002; Wolff et al., 2010; Austin and Hibbert, 2012). These
oscillations switching between warm (GIs for Greenland in-
terstadials) and cold (GSs for Greenland Stadials) climatic
phases are known as the Dansgaard-Oeschger (DO) cycles
(e.g. Bond et al., 1993; Dansgaard et al., 1993). These cy-
cles are punctuated by abrupt ice-sheet discharge events, the
so-called Heinrich Events (HEs), corresponding to some of
the DO coldest phases, and are illustrated in the North At-
lantic sediments by layers rich in silico-clastic coarse parti-
cles (ice-rafted debris – IRD; e.g. Bond et al., 1992). IRD
are transported by floating ice and preferentially deposited
in the Ruddiman Belt (Ruddiman, 1977), a latitudinal belt
(40 to 50° N) identified as the preferential area of melting for
icebergs during these discharge periods. Study of these IRD
particles allowed the characterisation of sources involved in
these events with great influence of the Laurentide ice sheet
during HE1, HE2, HE4 and HE5, and a greater influence
of European ice sheets during HE3 and HE6 (cf. Hemming,
2004 and references therein). At the same time, the North
Atlantic basin recorded large changes in hydrological condi-
tions, corresponding to temperature and salinity decreases in
sea-surface waters and changes in the meridional overturn-
ing circulation (MOC), with peculiar consequences on the
North Atlantic Drift (NAD) and on the formation of North
Atlantic Deep Waters (NADW) (e.g. Bond et al., 1992; Bond
and Lotti, 1995; Broecker, 1997; Kissel et al., 1999a; Alley
et al., 1999; Clark et al., 2002; Ramhstrof, 2002; Broecker,
2003; Hemming, 2004). Even though these phenomena are
now relatively well identified, the causes and consequences
of such climatic changes and dynamic interactions between
ice, ocean and atmosphere remain uncertain. For instance,
the role of ice shelves developed at the margins of Euro-
pean ice sheets during the last glacial period has probably
been underestimated, as suggested by recent works (Hulbe
et al., 2004; Alvarez-Solas et al., 2010). There is also a need
for a better definition of the geographical fingerprint of such
events to help identifying the likely source regions and per-
turbation mechanisms (Roche et al., 2010).

Here we propose to reconstruct the paleoclimatic and
paleoceanographic setting of the southern Faeroe Islands dur-
ing MIS3 in order to explore the evolution of the surface At-
lantic Ocean under the influence of growing European ice
sheets. To reach this goal, analysis of organic-walled di-
noflagellate cyst (dinocyst) assemblages was undertaken on
core MD99-2281, retrieved off SW Faeroe, and quantifica-
tions of surface hydrological parameters were obtained from
ecological transfer functions (the modern analogue technique
– MAT, e.g. Guiot and de Vernal, 2007) applied to these

assemblages. Dinocysts have long since been used as sea-
surface paleohydrological proxies, as their modern biogeо-
graphical distribution in surface sediments appears to be
tied to ecological parameters of the overlying surface wa-
ter masses (e.g. Harland, 1983; Rochon et al., 1999a,b, 2008;
Marret and Zonneveld, 2003; de Vernal and Marret, 2007).
Dinocysts are organic microfossils and derive from the sex-
ual reproduction of dinoflagellates. The wall of dinocysts,
comparable to the sporopollenin of pollen grains, has been
called dinosporin (e.g. De Leeuw et al., 2006) and confers to
dinocysts a strong resistance to dissolution, and a good fos-
silization potential (e.g. Evitt, 1985; Rochon et al., 1999a,b;
Marret and Zonneveld, 2003). Recent studies however sug-
gest that this compound is in fact closer to cellulose than to
sporopollenin (e.g. Versteegh et al., 2012), but their preser-
vation potential is still high, even if the assemblages can
be altered by oxidation processes in some specific environ-
ments (e.g. Zonneveld et al., 2007; Bogus et al., 2012). They
are among the rare proxies offering access to past hydrolog-
ical conditions of polar and/or glacial extreme environments
(e.g. de Vernal and Rochon, 2011).

In addition, our study compiles a set of multi-proxy
data (x-ray fluorescence – XRF – core scanning, magnetic
properties, abundances of the polar planktonic foraminifera
*Neogloboquadrina pachyderma* s and IRD concentrations)
that contributes to document and discuss the response of the
ocean during MIS3.

## 2 Environmental setting

2.1 Modern oceanic and sedimentological setting

2.1.1 Physiography

Core MD99-2281 (60°21′N, 09°27′W) was retrieved by the RV *Marion-Dufresne* during the IMAGES (International Marine Global Change Study) V - MD99-114 cruise (Labeyrie et al., 1999). It is made of 29 meters of hemipelagic silty clays (Labeyrie et al., 1999; Boulay, 2000). The core was collected at 1197 m deep at the foot of the Faeroe Bank located at the northern end of the Rockall Trough (Figs. 1 and 2). It has been retrieved at a location where seismic continuous parallel draping internal reflectors (see seismic section C in Boldreel et al., 1998) have been attributed to pelagic sediments deposited in a low-energy, deep-water environment, unaffected by the strong current activity (Boldreel et al., 1998).

Sediment accumulation in the Rockall Trough is primar-
ily constrained by bottom currents (Kuijpers et al., 2002).
This trough extends from 53° to 60° N west of the UK, with
water depths ranging from 4000 m (southern part) to 1000 m
(northern part). The topography is relatively complex (Mas-
son et al., 2002). The northwest margin is bordered by three
separate seamounts (George Bligh, Lousy and Bill Bailey's)
<figure><img src="image_4.png" /></figure>

**Fig. 1.** Location of the studied core (MD99-2281: this work) and other cores discussed in the paper (MD95-2009; MD95-2010 – Eynaud et al., 2002; ENAM93-21 – Rasmussen and Thomsen, 2004; core SO82-5 – Van Kreveld et al., 2000; MD04-2822; MD95-2006 – Knutz et al., 2001; Hibbert et al., 2010; MD04-2829CQ – Hall et al., 2011). The yellow arrows depict the pathway of the main modern surface currents (Fogelqvist et al., 2003). Dotted grey lines locate major modern hydrographic fronts: PF = Polar Front, AF = Arctic Front, SAF: Sub-Arctic Front (Dickson et al., 1988). The IF = Iceland Faeroe Front is indicated by the dark blue dotted line (Rasmussen et al., 2002b). The palaeogeography of the adjacent European continent during the Last Glacial Maximum with the hydrographic system/paleovalleys of the Manche River (Lericolais et al., 2003) and the paleo-coastline (bold blue line) at 21 ka BP (Bourillet et al., 2003) are also depicted. The full glacial extension of the North Hemispheric ice sheets (Grosswald and Hughes, 2002, and Ehlers and Gibbard, 2004) is schematized by the red dotted-line.

culminating at more than 500 m water depth. The merging of the different deep channels between these seamounts forms a connection between the Rockall Trough and the Iceland Basin (Fig. 2).

In the northeastern part, the Rockall Trough is bor-
dered by the Faeroe Bank (FB) and the Wyville-Thompson
ridge culminating at less than 100 and 500 m water depths,
respectively (Bett, 2001, Fig. 2). The barrier, which separates
the Faeroe Bank Channel and the Faeroe Shetland Channel,
plays an essential role in the distribution of water masses and
the dynamic of bottom currents (Kuijpers et al., 2002). These
landforms are the result of the opening of the North Atlantic
basin during the Lower Eocene, and post-rift tectonic events
that followed (Howe et al., 2006).

<figure><img src="image_5.png" /></figure>

**Fig. 2.** Detailed physiography of the studied area. (A) Isobath: the first depth contour represents the 200 m isobar on the shelf. The next depth contour is the 1000 m isobar. At water depths greater than 1000 m, the distance between the isobars is 1000 m. Bill Bailey's Bank (BBB), Faeroe Bank (FB), George Bligh Bank (GBB), Hatton Bank (HB), Hatton-Rockall Basin (HRB), Lousy Bank (LB), Rosemary Bank (RB) (Howe et al., 2006). Surface and subsurface Currents: North Atlantic Drift (NAD), East North Atlantic Water (ENAW). Bottom currents: Norwegian Sea Overflow Water (NSOW), North East Atlantic Deep Water, Wyville-Thompson Ridge Overflow Water (WTROW) (Kuijpers et al., 1998, 2002; Holiday et al., 2000; New and Smythe-Wright, 2001; Howe et al., 2006). (B) Bathymetric map (contour interval 100 m) zooming on the core site with location of the C seismic section documented by Boldreel et al. (1998), who identified pelagic deposits unaffected by currents in the coring area.

2.1.2 Hydrographic and oceanographic setting

The dynamic of surface currents in the Rockall Trough area is primarily driven by the influence of two major gyres: the subtropical anticyclonic and subpolar cyclonic North Atlantic gyres. The latter is characterised by numerous streams, including the well-known poleward current of the NAD, which defines a peculiar distribution pattern of fronts and water masses in the area. The topography of the region, and more specifically the presence of Iceland, forces the penetration of the NAD across the “Iceland–Faeroe–Scotland” ridge where our study site is located. The NAD splits into two branches in the vicinity of the Faeroe Islands (Figs. 1 and 2; Larsen et al., 2008, 2009).
Core MD99-2281 is presently located under the influence of the modified North Atlantic surface waters, between the Arctic Front (AF) to the north and the sub-Arctic Front (SAF) to the south (Fig. 1). The modern sea-surface temperature (SST) in the area varies from 8°C during winter to 12°C during summer (WOA, 1998). Below the surface, the water-mass distribution and deep circulation are governed by the bathymetric complexity of the sector (Fig. 2, Holliday et al., 2000).

The first 1000 m of the water column are occupied by a relatively homogeneous body of water whose temperature and salinity ranges characterise the East North Atlantic Water (ENAW, Holliday et al., 2000; New and Smythe-Wright, 2001). Below dominates the North East Atlantic Deep Water (NEADW). A core of this water enters the Trough from the south and, trapped by the topography to the north, makes a cyclonic gyre, joining the Wyville-Thompson Ridge Overflow Water (WTROW) to continue its trajectory into the Iceland Basin (Howe et al., 2006). North of the Wyville-Thompson Ridge, the Norwegian Sea Overflow Water (NSOW) migrates southward across the Faeroe Shetland Channel and then turns west to enter the Faeroe Bank channel (Kuijpers et al., 2002) to join the southern flank of the Iceland-Faeroe Ridge. The NSOW intermittently crosses the Wyville-Thompson Ridge to join the LSW, then reaching our core location (Holliday et al., 2000; New and Smythe-Wright, 2001; Kuijpers et al., 1998, 2002; Rasmussen et al., 2002b).

## 2.2 Paleoclimatic interest

Previous paleoceanographic studies have demonstrated the high sensitivity of the Faeroe area to millennial-scale climatic variability and its intimate relation to the MOC dynamics (Rasmussen et al., 1996a,b, 2002a,b, 2008; Rasmussen and Thomsen, 2004, 2008; Eynaud et al., 2002; Voelker, 2002). During the last glacial period, the southern part of the Faeroe Islands was under the direct influence of the proximal European ice sheets (Fennoscandian ice sheet: FIS and the British-Irish ice sheet: BIIS) (Fig. 1). Therefore, core MD99-2281 has likely recorded the history of these ice sheets at a millennial scale, especially episodes of decay (including those of HEs) which occurred during MIS3, together with changes in the penetration and vigour of the NAD current, and in the latitude of the Arctic Front. In addition, our site may have recorded instabilities of the marine extensions of the ice sheets, i.e. ice shelves that are up to now poorly considered in HE dynamics (Hulbe et al., 2004; Alvarez-Solas et al., 2010). Another point concerns the potential existence of an “ice bridge” in the North Sea joining the BIS to the FIS between 30 and 25 cal ka BP. The presence of this ice bridge is still strongly debated, even if this hypothesis has gained increasing acceptance in recent years (Bradwell et al., 2008; Scourse et al., 2009; Hibbert et al., 2010; Chiverrell and Thomas, 2010; Toucanne et al., 2010).

Instabilities of proximal glaciers have had a major impact on the surface hydrology of the region and have strongly influenced the NAD, and, consequently, the NADW overflow from the Nordic Seas (e.g. Bond et al., 1992; Bond and Lotti, 1995; Broecker, 1997, 2003; Hemming, 2004).

## 3 Material and methods

Core MD99-2281 has been sampled every 10 cm between 2090 and 1200 cm. The 90 samples were then wet sieved to separate fractions above and below 150 µm. Both fractions were analyzed separately. Planktonic foraminiferal assemblages and ice rafted detritus (IRD) concentrations were determined on the > 150 µm fraction, whereas palynological analyses (dinocyst assemblages, freshwater algae) were conducted on the < 150 µm one.

### 3.1 Dinoflagellate cysts (dinocysts)

#### 3.1.1 Dinocyst assemblage analysis

Palynological preparations follow the protocol described by de Vernal et al. (1996), slightly modified at EPOC laboratory (cf. http://www.epoc.u-bordeaux.fr/index.php?lang=fr&page=eq-paleo-protocoles for details). The dinocyst specific determination was done with a Leica Microscope at a X400 magnification. Around 300 dinocysts were counted in each sample to obtain a good statistical representation of the assemblage. Species identification follows Rochon et al. (1999b). Relative abundances were calculated relative to the total sum of Quaternary dinoflagellate cysts and absolute abundances (concentration cysts/cm³ of dried sed.) were obtained thanks to the marker grain method (Lycopodium tablets, de Vernal et al., 1996; Mertens et al., 2009). Counts of non-Quaternary reworked dinocysts were done in order to evaluate the ratio Rd/Md [reworked dinocysts/modern dinocysts] which along the western European margin could provide a proxy of terrigeneous advection (e.g. Zaragosi et al., 2001; Eynaud et al., 2007; Penaud et al., 2009).

#### 3.1.2 Transfer functions

An ecological transfer function has been applied to the raw data derived from the counts (relative abundances) to quantitatively reconstruct the sea-surface paleohydrological parameters. The transfer function was performed with the R software ("R" version 2.7.0, R Development Core Team, 2008) using a script developed by Guiot and Brewer (www.cerege.fr/IMG/pdf/formationR08.pdf) for the modern analogue technique (MAT, e.g. Guiot and de Vernal, 2007). This statistical tool principally uses the statistic distance between fossil (paleoceanographic record) and current (modern database) assemblages. Modern dinocyst assemblages are compiled in a geo-referenced database of marine surface sediment samples. The database we used here is that
of de Vernal et al. (2008) with 1189 stations from North Atlantic, Arctic and North Pacific oceans and their adjacent seas (see recurrent updates at http://www.geotop.ca/fr/bases-de-donnees/dinokystes.html). Calculations rely on a weighted average of SST values (compiled from the 2001 version of the World Ocean Atlas) from the best five modern analogues, with a maximum weight given for the closest analogue in terms of statistical distance, i.e. dissimilarity minimum (e.g. Kucera et al., 2005; Guiot and de Vernal, 2007). Among the sea-surface parameters that we can reconstruct on the basis of dinocysts are SSTs and sea-surface salinities (SSSs) for the warm season, i.e. mean summer (June-July-August average), the cold season, i.e. mean winter (January-February-March average), the warmest (August) and the coldest (February) months of the year, but also the sea-ice cover duration (months per year). The reader is referred to Guiot and de Vernal (2007, 2011a,b) for a review of the theory of transfer functions, and to de Vernal et al. (2001, 2005) for a step-by-step description of the application of transfer functions to dinocysts, including discussion about the degree of accuracy of the method.

Standard deviations must be taken into account in the re-
constituted hydrological parameters because of the accuracy
of the database calibration. They have been established by
comparing the predictions provided by the transfer function
and direct instrumental measurements made on surface wa-
ters. The root mean square errors of prediction (RMSEP) for
the parameters presented in this paper are ± 1.1 °C ± and
1.6 °C for SST of February and August, respectively, and
± 1.1 months per year for the ice cover duration (Penaud,
2009). They are ± 2.2 and 2.5 psu for February and August
SSS, and ± 2.3 and 2.4 psu for mean winter and mean sum-
mer SSSs, respectively. The relatively large error for salinity
prediction is due to the inclusion of Arctic sites characterised
by low and highly variably salinities (down to 11). In the
> 30 salinity domain, the error of prediction is 0.63 (Eynaud
et al., 2012). Associated and additional information is listed
on the GEOTOP website (http://www.geotop.ca/).

3.2 Magnetic properties

The core has been sub-sampled using u-channels to per-
form detailed magnetic analyses. In order to complement on-
board susceptibility measurements made with a large diam-
eter coil (low resolution), magnetic susceptibility has been
repeated at the LSCE with a 45-mm diameter MS2-C Bart-
ington coil. The data were generated every 2 cm on the u-
channels and they were normalized by the volume to get the
volume susceptibility (κ). The natural remanent magnetiza-
tion (NRM) was measured using a 755-R 2G cryogenic pass-
through magnetometer, equipped with a high resolution set of
coils and placed in the μ-metal shielded room of LSCE (see
Kissel et al., 1999b for detailed laboratory procedures). An
in-line alternating field (AF) demagnetization unit was used
for the stepwise demagnetization at 5, 10, 15, 20, 25, 30, 35,

40, 45, 50, 60 and 80 mT. After demagnetization at 80 mT, 87 to 97 % of the initial magnetization was removed. The an-hysteretic remanent (ARM) and isothermal remanent (IRM) magnetizations were acquired and demagnetized also on u-channels using the same method and instruments. ARM was acquired along the vertical geographical axis in a peak alternating field of 100 mT and a steady bias field of 50 µT. During acquisition, the samples were translated through the coils at a speed of about 1 cm s⁻¹ (Brachfeld et al., 2004). After acquisition, the ARM was then progressively demagnetized using 10 steps at 10, 15, 20, 25, 30, 35, 40, 50, 60, and 80 mT. Saturated IRM (SIRM) was then acquired, also along the vertical geographical axis in six steps (0.05, 0.1, 0.2, 0.3, 0.5 and 1 T) using a 2G 1.6 m long pulsed solenoid and stepwise demagnetized using the same 10 steps as for the ARM. During the demagnetization of NRM, ARM and IRM, the u-channels were translated at a speed of about 4 cm s⁻¹ through the demagnetization coils. All the data were acquired using software developed at LSCE, which allow to considerably reduce the translation speed during the measurements while the u-channel passes through the pick-up coil. This is particularly adapted to the measurements of very high magnetizations as the ones of core MD99-2281 because the reduced translation speed allows to reliably count the flux jumps. Previous works already conducted in the area established that the main magnetic mineral is magnetite (Kissel et al., 1999a; Ballini et al., 2006). The co-variations obtained between the low field magnetic susceptibility, the ARM and the IRM data therefore indicate that these parameters are only concentration dependent in this sediment (except during HE4, see below).

3.3 XRF (x-ray fluorescence) elementary analyses

A non-destructive X-ray fluorescence (XRF) scan of the core was done using the CORTEX scanner at a resolution of 2 cm (Richter et al., 2006). To complement these data, a high resolution scan (every 2 mm) was done at EPOC laboratory with the Avaatech XRF Core Scanner on the sections (990–2100 cm) investigated for dinocysts. This provides us with a semi-quantitative approach of the sediment chemical composition (for major and some minor chemical elements), and thus permits estimation of changes in the nature of sediments along time. Richter et al. (2006) demonstrated on the basis of the low resolution XRF scan of a proximal core MD99-2282 (59°29′N; 10°34′W) that the abundance of some elements, such as Ca, or the ratio K/Ti, provides information about glacial–interglacial alternances. It was also observed by Ballini et al. (2006) at the DO scale. Among other applications of the XRF tool, the ratio K/Ti could provide information about lithic grain sources, and the (Fe + Ti)/(Ca + Sr) ratio further gives access to the proportion of lithogenic versus biogenic components within the sediments. The variations in Fe, Mn and S abundances further help to identify early diagenesis processes (Mojtahid et al., 2005, Richter et al., 2006).
## 3.4 Analysis of the coarse fraction > 150 µm

### 3.4.1 Planktonic foraminifera

These calcareous-walled protists form part of the zoo-
plankton, and their geographical distribution covers al-
most all oceans (e.g. Kucera, 2007). There are currently
about forty species in the modern assemblage of planktonic
foraminifera. One of the mesopelagic species well known
and frequently used in paleoclimatology and biostratigraphy
is the polar taxon, *Neogloboquadrina pachyderma* sinistral
(NPS) coiled form. Quantification of this taxon abundance
(%) in the planktonic assemblages provides access to a qual-
itative estimate of surface water temperatures, and depend-
ing on the areas, directly follows the Polar Front migra-
tions (e.g. Ericson, 1959; Eynaud et al., 2009). Counts of
the relative abundances of this species have been made on
known aliquots of the dried > 150 µm residues (total sum of
planktonic foraminifera of minimum 300 specimens) from
the same samples studied for dinocyst assemblages.

In addition, stable isotope measurements were done on
NPS shells. Specimens were picked out from the 200-
250 µm fraction and analysed at EPOC using an Optima Mi-
cromass mass spectrometer. For each measurement, 80 µg
mean weight aliquots (5 to 6 specimens) were treated with
individual acid baths. The extracted CO₂ gas was analyzed
versus a laboratory reference gas calibrated with respect to
VPDB (Vienna Pee Dee Belemnite) using the international
standard NBS 19. For this paper, only δ¹⁸O data will be
discussed.

### 3.4.2 Ice-rafted debris

In hemipelagic cores, ice rafted detritus (IRD) are gener-
ally determined in the coarse lithic > 150 µm grain fraction.
These large lithic grains (LLG) are actually torn-off bedrock
under ice sheet dynamic processes. They were disseminated
by icebergs throughout the North Atlantic Ocean during ice-
sheet calving events, as for instance during Heinrich events
(Heinrich, 1988; Bond et al., 1992; Bond and Lotti, 1995;
Elliot et al., 2001; Broecker, 2003; Hemming, 2004). The
study of IRD concentrations (number of LLG > 150 µm g⁻¹
dry sediment) thus provides a way to document ice-sheet dy-
namics throughout time. The LLG were counted in the same
samples as those used for NPS relative abundance charac-
terisation. Their raw number per samples varies from 64 to
955 grains per g of dry sediment.

# 4 Stratigraphy of the core

Nine AMS ¹⁴C dates measured on planktonic foraminifera
monospecific samples (cf. Table 1) have been obtained on the
topmost eleven meters of core MD99-2281 (Table 1, Boulay,
2000), providing a first stratigraphic frame between 10 and
24 ka cal BP. For this work, four supplementary AMS ¹⁴C

dates were obtained between 12 and 19 m in the core. Re-
garding the time period we consider for this paper, radio-
carbon dates were calibrated using the “glacial polynomial”
from Bard (1998) and Bard et al. (1998), which provides cal-
ibration up to 36 ¹⁴C ka BP (Table 1).

Following several paleoceanographic studies in the area (Rasmussen et al., 1996b, 1997; Kissel et al., 1999a; Eynaud, 1999; Stoner et al., 2000; Laj et al., 2000), these age control points were combined to tie-points generated by the comparison of the record of magnetic concentration parameters to the NGRIP-GICC05 δ¹⁸O record (Svensson et al., 2008), i.e. the recommended North Atlantic regional stratotype (Austin and Hibbert, 2012). The rationale is that marine records of magnetic concentration parameters from MIS3 are consistent across the North Atlantic Basin on the path of different overflow branches of the NADW and could be tied on the high frequency climatic variability (DO cycles, e.g. Sanchez-Goñi and Harrison, 2010; Wolff et al., 2010) characteristic of this period (e.g. Kissel et al., 1999a). This approach is comparable to an event-based stratigraphy and has recently been acknowledged as one of the most robust ways to establish age models for marine sediments (Austin and Hibbert, 2012).

Our stratigraphy has thus been constructed by linking
the magnetic concentration increases to the beginning of
Greenland interstadial (GIs) warmings (Austin and Hibbert,
2012 recommended protocol) and conversely, the magnetic
concentration decreases to Greenland stadial (GS) coolings
(Fig. 3). The chronostratigraphic ages used for this work con-
form to the recent update done by Wolff et al. (2010) on DO
events. As a result, 18 stratigraphic control points have been
added to the AMS ¹⁴C dates (Table 1, Fig. 3). The age model
has been constructed on the basis of a linear interpolation
between tie points (Table 1, Fig. 4) giving an average sedi-
mentation rate of about 80 cm ka⁻¹ for the studied period.

Additional tie points, independent from climate, are also
retrieved from the record of the changes in the Earth's mag-
netic field intensity, namely the two prominent lows at-
tributed to the Mono Lake (MLE) and to the Laschamp (LE)
excursion (Laj et al., 2000). Due to the modulation of cos-
mogenic isotopes production by the dipolar Earth magnetic
field, these two excursions have their counterpart in the ¹⁰Be
and ³⁶Cl records from ice (Wagner et al., 2000). Therefore,
they allow a robust correlation between ice and sediment cli-
matic records. In both archives, the MLE occurred during
GS7 and the LE is centered on GI10. We observe that both
are at the right place in core MD99-2281 (Figs. 3 and 4),
giving confidence on the correlation discussed above.
<table><caption>**Table 1.** Radiocarbon dates for core MD99-2281 and age-control points derived from the correlation of the MS signal with the NGRIP GICC05 δ¹⁸O record (Svensson et al., 2008; Wolff et al., 2010). See also Fig. 3.</caption><thead><tr><th>Depth (cm) in core MD99-2281</th><th>LAB_ID</th><th>AMS <sup>14</sup>C age uncorrected (BP)</th><th>Calendar age corrected (cal BP) (Bard, 1998)</th><th>Calendar age corrected (cal BP) CALIB 6.0</th><th>σ ranges: [start]</th><th>σ ranges: [end]</th><th>Coherency of the two calibration method [Δ]</th><th>Tie points (age in cal BP)</th><th>Dated material/age control points/references</th></tr></thead><tbody><tr><td>40</td><td>GifA 100405</td><td>10 260 ± 100</td><td><strong>11 165</strong></td><td>11 261</td><td>11 137</td><td>11 384</td><td>95</td><td>11 165</td><td>G. bulloides, Boulay (2000)</td></tr><tr><td>170</td><td>GifA 100406</td><td>12 970 ± 110</td><td><strong>14 729*</strong></td><td>14 657</td><td>14 156</td><td>15 158</td><td>72</td><td></td><td>G. bulloides, Boulay (2000)</td></tr><tr><td>270</td><td>GifA 100407</td><td>10 800 ± 100</td><td><strong>12 081</strong></td><td>12 229</td><td>12 070</td><td>12 388</td><td>148</td><td>12 081</td><td>G. bulloides, Boulay (2000)</td></tr><tr><td>400</td><td>GifA 100408</td><td>11 060 ± 110</td><td><strong>12 400</strong></td><td>12 431</td><td>12 392</td><td>12 470</td><td>31</td><td>12 400</td><td>G. bulloides, Boulay (2000)</td></tr><tr><td>510</td><td>GifA 100409</td><td>15 290 ± 130</td><td><strong>17 529</strong></td><td>18 037</td><td>17 931</td><td>18 142</td><td>507</td><td>17 529</td><td>N. pachyderma s., Boulay (2000)</td></tr><tr><td>580</td><td>GifA 100410</td><td>15 760 ± 170</td><td><strong>18 092</strong></td><td>18 219</td><td>18 209</td><td>18 228</td><td>126</td><td>18 092</td><td>N. pachyderma s., Boulay (2000)</td></tr><tr><td>650</td><td>GifA 100411</td><td>16 040 ± 140</td><td><strong>18 427</strong></td><td>18 777</td><td>18 651</td><td>18 902</td><td>349</td><td>18 427</td><td>N. pachyderma s., Boulay (2000)</td></tr><tr><td>780</td><td>GifA 100412</td><td>18 060 ± 160</td><td><strong>20 830</strong></td><td>20 630</td><td>20 562</td><td>20 697</td><td>200</td><td>20 830</td><td>N. pachyderma s., Boulay (2000)</td></tr><tr><td>1180</td><td>GifA 100414</td><td>23 340 ± 240</td><td><strong>26 993</strong></td><td>27 144</td><td>27 037</td><td>27 251</td><td>151</td><td>26 993</td><td>N. pachyderma s., Boulay (2000)</td></tr><tr><td>1240</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>27 430</td><td>Correlation NGRIPGICC05, G13 termination, Wolff et al. (2010)</td></tr><tr><td>1275</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>27 730</td><td>Correlation NGRIPGICC05, G13 warming, Wolff et al. (2010)</td></tr><tr><td>1280</td><td>SacA 19117</td><td>24 210 ± 140</td><td><strong>27 992</strong></td><td>28 567</td><td>28 326</td><td>28 807</td><td>574</td><td>27 992</td><td>N. pachyderma s., ARTEMIS 2010</td></tr><tr><td>1300</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>28 550</td><td>Correlation NGRIPGICC05, G14 termination, Wolff et al. (2010)</td></tr><tr><td>1355</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>28 850</td><td>Correlation NGRIPGICC05, G14 warming, Wolff et al. (2010)</td></tr><tr><td>1430</td><td>SacA 19118</td><td>27 430 ± 140</td><td><strong>31 652</strong></td><td>31 306</td><td>31 211</td><td>31 400</td><td>346</td><td>31 652</td><td>N. pachyderma s., ARTEMIS 2010</td></tr><tr><td>1440</td><td>SacA 19119</td><td>27 940 ± 140</td><td><strong>32 226*</strong></td><td>31 576</td><td>31 411</td><td>31 740</td><td>650</td><td></td><td>N. pachyderma s., ARTEMIS 2010</td></tr><tr><td>1475</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>31 950</td><td>Correlation NGRIPGICC05, G15 termination, Wolff et al. (2010)</td></tr><tr><td>1523</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>32 450</td><td>Correlation NGRIPGICC05, G15 warming, Wolff et al. (2010)</td></tr><tr><td>1548</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>33 290</td><td>Correlation NGRIPGICC05, G16 termination, Wolff et al. (2010)</td></tr><tr><td>1605</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>33 690</td><td>Correlation NGRIPGICC05, G16 warming, Wolff et al. (2010)</td></tr><tr><td>1650</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>34 730</td><td>Correlation NGRIPGICC05, G17 termination, Wolff et al. (2010)</td></tr><tr><td>1711</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>35 430</td><td>Correlation NGRIPGICC05, G17 warming, Wolff et al. (2010)</td></tr><tr><td>1755</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>36 570</td><td>Correlation NGRIPGICC05, G18 termination, Wolff et al. (2010)</td></tr><tr><td>1820</td><td>SacA 19121</td><td>34 610 ± 290</td><td><strong>39 586*</strong></td><td>39 105</td><td>38 722</td><td>39 488</td><td>481</td><td></td><td>N. pachyderma s., ARTEMIS 2010</td></tr><tr><td>1880</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>38 170</td><td>Correlation NGRIPGICC05, G18 warming, Wolff et al. (2010)</td></tr><tr><td>1930</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>39 810</td><td>Correlation NGRIPGICC05, G19 termination, Wolff et al. (2010)</td></tr><tr><td>1960</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>40 110</td><td>Correlation NGRIPGICC05, G19 warming, Wolff et al. (2010)</td></tr><tr><td>2030</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>40 710</td><td>Correlation NGRIPGICC05, G110 termination, Wolff et al. (2010)</td></tr><tr><td>2090</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>41 410</td><td>Correlation NGRIPGICC05, G110 warming, Wolff et al. (2010)</td></tr><tr><td>2110</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>42 290</td><td>Correlation NGRIPGICC05, G111 termination, Wolff et al. (2010)</td></tr><tr><td>2170</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td><strong>43 290</strong></td><td>Correlation NGRIPGICC05, G111 warming, Wolff et al. (2010)</td></tr></tbody></table>

Bold characters: age values used for the interpolation. * Age reversals removed.
<figure><img src="image_6.png" /></figure>

**Fig. 3.** Magnetic susceptibility (MS) record from core MD99-2281 (dark line = onboard data versus red line = low field magnetic susceptibility) and its correlation to the NGRIP GICC05 $\delta^{18}$O record (Svensson et al., 2008) to define supplementary age control-points over the MIS3 period. See also Table 1. Pink bars underline GIs, labels and limits after Wolff et al. (2010). The horizons having recorded the geomagnetic excursions (MLE for Mono Lake excursion and Laschamp) are also indicated.

# 5 Results

## 5.1 Dinocyst data

Thirty dinocyst taxa were identified in the MIS3 assemblages. Their absolute abundances in the sediment (concentrations) are high and vary between 1000 and 90 000 cysts/cm³ along the studied section, with a mean value around 10 000 cysts/cm³ (Fig. 5). As a whole the dinocyst concentrations follow the MS properties data, thus indicating maximum dinocyst production and/or deposition and/or preservation to the studied site during warm GIs, and conversely during cold GSs. Some strong excursions were recorded over the studied interval, with especially two peaks in concentrations of about 90 000 and 57 000 cysts/cm³ at 1660 and 1950 cm depths, respectively.

The assemblages are dominated by five species (Fig. 5, Table 2), which are: *Bitectatodinium tepikiense* (largely dominates the assemblages), *Operculodinium centrocarpum*, *Brigantedinium spp.*, *Islandinium minutum* (upper part of the section), and *Pentapharsodinium dalei* (lower part of the section). Table 2 summarizes the most salient features recorded for the encountered species both at present and along the studied section of core MD99-2281. The

<figure><img src="image_7.png" /></figure>

**Fig. 4.** Age model and sedimentation rates for the last 50 ka BP in core MD99-2281 (see also Table 1). Dates in red are those deduced from the correlation (see also Fig. 3) of the MS record and the NGRIP GICC05 $\delta^{18}$O (Wolff et al., 2010). A comparison of the relative paleointensity (RPI) record from core MD99-2281 with the reference RPI curve of GLOPIS-75 is shown for the 32–42 ka interval.

continuous presence along the record of specimens from the genus *Brigantedinium*, with occurrence of other peridinioid-species, indicate that the alteration of the assemblage by oxygenation processes (e.g. Zonneveld et al., 2007) was not intense. This is also supported by high dinocyst concentrations in the sediments during GIs which record active deep-sea circulation (e.g. Kissel et al., 1999a), and thus better ventilation.

## 5.2 Sea-surface paleohydrology (from dinocyst transfer function) during MIS3

The sea-surface paleohydrological conditions quantified on the basis of the dinocyst MAT transfer function (database 1189) are shown in Fig. 5 (right panel).

In the studied section, August SSTs vary between 8.5 and 17 °C, with a mean around 16 °C, and February SST between 6 and -1 °C. The reconstructed February SSTs are considerably lower than the modern ones over the area, which fluctuate around an average of 8.7 °C (WOA, 1998). In contrast, the reconstructed August SSTs are well above the modern averages of 12.1 °C (WOA, 1998). Our reconstructions indicate that a positive anomaly of nearly 4 °C (in comparison to the modern conditions) was recorded over the area during MIS3 in summer.

Sea-ice cover duration varies between 0 and 2 months/year along the main part of the section, except in the topmost part of the studied section where it reaches values up to 5 months per year. These values are above the modern ones, the area being presently free of sea-ice even in winter, but are still
<figure><img src="image_8.png" /><figcaption>**Fig. 5.** Results from dinocyst analyses (see also Table 2) over the section 1200–2100 cm in core MD99-2881, with (a) dinocyst total concentrations in the sediment (nb. of cyst/cm³ of dry sed.); (b) relative abundances of some selected species; (c) Dinocyst derived reconstructions (modern analogue technique) of sea-surface conditions (sea-ice cover duration ± 1.1 months per year; February SST ± 1.1 °C; August SST ± 1.6 °C; February SSS ± 2.2; August SSS ± 2.5; mean winter SSS ± 2.3; mean summer ± 2.4 psu). The dotted lines surrounding each reconstructed parameter correspond to minima and maxima values found in the set of the 5 selected analogues. Modern values of SSTs after WOA (1998). These results are compared to (d) the ARM (for anhysteretic remanent magnetization) data and (e) the sedimentation rate calculated from our age model (see Fig. 4).</figcaption></figure>

very low. They should be considered rather as an index of
the intermittent presence of ice, regarding the RMSEP of
1.1 months per year.

Paleosalinity reconstructions show variations within a
range of 4 psu, from 30 to 34 in summer (during mean
summer months, i.e. July-August-September, and mean Au-
gust only) and 31 to 35 in winter (during mean winter
months, i.e. January-February-March as well as February
only). SSS oscillations are recorded synchronously during
both seasons (same trends along our record), even if sum-
mer SSSs tend to be less saline than during winters. This
is probably due to the influence of the summer melting of
proximal ice sheets. Mean annual modern values over the
area are of 35.3, with a positive Δ 0.07 psu difference be-
tween winter and summer seasons (WOA 2009, http://www.
nodc.noaa.gov/OC5/WOA09/netcdf_data.html). Sea-surface
waters were thus sensibly more brackish in the area during
MIS3.

5.3 Multiproxy compilation

Figure 6 compiles the multiproxy dataset we obtained on core MD99-2281 between 1100 and 2100 cm, i.e. 26.5 and 41.5 ka cal BP (± 1.6 ka considering the GICC05 chronological uncertainty, Svensson et al., 2008). High-resolution XRF data are compared to MS data, planktonic foraminifera data and to large lithic grain (LLG > 150 µm) concentration in the sediment, this latter indicating IRD arrivals on site.

The XRF data show a high variability in correlation with
changes in magnetic records (Fig. 6) with each minimum in
magnetic concentration corresponding to a minimum in the
XRF calcium detection curve (XRFCa), and conversely max-
ima in magnetic concentration matching maxima in XRFCa.
This indicates that the magnetic concentration record does
not result from a dilution effect by calcium carbonate and
really illustrates changes in the input of magnetites by over-
flow water, as for the other cores distributed along the path of
the NADW. This correlation also suggests relatively low bio-
genic calcareous inputs (Richter et al., 2006) during stadial
periods, alternating with higher inputs during interstadials.
At the opposite, interstadials record increases in the XRFCa
<figure><img src="image_9.png" /></figure>

**Fig. 6.** Multiproxy results obtained on the MIS3 section of core MD99-2281 compared to the (a) δ¹⁸O NGRIP GICC05 record (Svensson et al., 2008), with: (b) Magnetic properties (volume susceptibility – *Kappa* – ARM; Isothermal Remanent – IRM – magnetizations); XRF data with (c) Ca raw counts, (d) ratio (Fe + Ti)/(Ca + Sr) and (e) ratio K/Ti (pink curve: smoothed 3 points running average); (f) planktonic foraminifera absolute concentration in the sediment ([PF] = Nb. of shells > 150 µm per g of dry sed.); (g) relative abundances (%) of the polar taxon *N. pachyderma sin.* (NPS); (h) δ¹⁸O measurements on monospecific samples of NPS shells (200–250 µm); (i) Large lithic grain concentration in the sediment (LLG = Nb. of grain > 150 µm per g of dry sed.); (j) Reworked dinocyst versus modern dinocyst concentrations, i.e. Rd/Md ratio; (k)–(n) dinocyst derived reconstructions of sea-surface conditions with (k) February SST ± 1.1 °C; (l) August SST ± 1.6 °C; (m) sea-ice cover duration ± 1.1 months per year, (n) sea-surface salinities ± 2.5 psu. Greenland stadials (light grey bands), including Heinrich events (dark grey bands) are defined in agreement with the stratigraphic synthesis of Wolff et al. (2010). Within HE4, the 39 ka event (see text) is highlighted.

content occurring in concordance with higher concentration of planktonic foraminifera shells in the sediment and lower NPS δ¹⁸O values. It is worth noting that these conditions are particularly characteristic of the interstadial terminal parts in our record.

The XRF Ca is corroborated by the ratios (Ti + Fe)/(Ca + Sr), K/Ti and the LLG concentration in the sediment, which together indicate strong terrigeneous input preferentially during stadials, but poor in magnetic content. The K/Ti XRF intensity ratio reflects changing provenance of the terrigenous fraction, with an enhanced ice-rafted continental crust (high K) contribution during stadials vs. relatively more bottom current derived basaltic input (high Ti) during interstadials (cf. Ballini et al., 2006; Richter et al., 2006; Grützner and Higgins, 2010). Considerable short-term variability is superimposed upon the general stadial-interstadial pattern.

## 6 Towards an integrated view of the response of sea-surface paleoenvironments south of the Faeroe Islands during MIS3

The Marine Isotopic Stage 3 covers more than 30 000 yr. Its lower stratigraphic limit is located at around 59 ka BP (e.g. Martinson et al., 1987) and its upper stratigraphic boundary has recently been revised by Wolff et al. (2010) and positioned at the transition between GS4 and GI3 around 27.73 ka cal BP. Other previous work considered MIS3 termination at the end of HE3 after 29 ka cal BP (Voelker, 2002; Bradwell et al., 2008; Hibbert et al., 2010).

Figure 6, integrating the most significant results of our study, summarizes salient features detected during MIS3 in core MD99-2281. If the DO structure (cold GSs versus warm GIs) is clearly recognizable in this multiproxy dataset, it is worth noting the good consistency between the different studied paleoenvironmental proxies. The highlighted
<table><caption>**Table 2.** Summary of the ecology of the main species recorded over the studied section of core MD99-2281.</caption><thead><tr><th>Dinocyst specie</th><th>Modern bioprovinces (Rochon et al., 1999b)</th><th>Known ecological preferences (Rochon et al., 1999b; Marret and Zonneveld, 2003; de Vernal and Marret, 2007)</th><th>Presence in core MD99-2281</th></tr></thead><tbody><tr><td><i>B. tepikiense</i></td><td>– temperate to subarctic (high abundance spots at the St. Laurence and Baltic Sea outlets)</td><td>– tolerates a wide range of SST and salinities, but most represented in areas when summer SST exceed 10°C<br/>– strong affinity for stratified surface waters characterised by a large seasonality (Rochon et al., 1999b)</td><td>– dominates the assemblages along most of the MIS3 section with percentages sometimes reaching 95%. These atypical high values are not observed at present in the North Atlantic ocean surface sediments (max. of 30%).</td></tr><tr><td><i>O. centrocarpum</i></td><td>– observed from mid to high latitudes<br/>– associated with the North Atlantic Drift (NAD) pathway</td><td>– tolerates a wide range of temperatures and salinities</td><td>– second most abundant species encountered, especially at the base of the section</td></tr><tr><td><i>Brigantedinium spp.</i></td><td>– ubiquitous, often dominant in upwelling zones<br/>– also associated with sub-arctic and arctic water masses</td><td>– tolerates a wide range of temperatures, salinities and sea-ice cover duration<br/>– heterotrophic taxa, has been related to its main food source, diatoms.</td><td>– well represented at the top of the section<br/>– Relatively difficult to identify at the species level due to the crumbled aspect of the cyst</td></tr><tr><td><i>I. minutum</i></td><td>– sub-polar to polar</td><td>– associated with long sea-ice cover duration (from 8 to 12 months)</td><td>– poorly represented in our sequence, except in the upper part of the section</td></tr><tr><td><i>P. dalei</i></td><td>– sub-polar to polar environments in North Atlantic basins, north of the NAD</td><td>– tolerates a wide range of SSSs, SSTs, and sea-ice cover duration<br/>– requires a 4°C SST minimum during summer<br/>– abundant in regions with a large seasonal gradient of SSTs</td><td>– low abundances in the sequence; presence most noticeable at the base of the section</td></tr></tbody></table>

variability is well defined through qualitative data as well as through quantitative ones, with a pattern which reveals the sequencing of MIS3 millennial scale events over the area. The following discussion has been structured on this basis, considering separately GSs and GIs in order to better discriminate their specific paleoenvironments.

## 6.1 Interstadials (GIs)

The pattern detected for GIs on the basis of our study (Fig. 6) appears to be reproducible for almost all the recovered

GIs and reveals milder climatic and sea-surface conditions. XRF Ca, δ¹⁸O low values (signing a warm and/or meltwater advection) and the absolute abundance of planktonic foraminifera, appear together in close agreement with the reconstructed February SSTs. They all converge to indicate mean sea-surface conditions which were probably nearly as warm as the modern ones. However, detailing the internal structure of the GIs reveals pulsed warm excursions (especially well expressed in February SSTs) that are detected late during GIs, at the transition towards GSs. This is quite
atypical considering the abruptness classically attributed to
DO atmospheric warmings (Severinghaus and Brook, 1999;
Lang et al., 1999; Schulz, 2002). These late events also coin-
cide with brief decreases in summer SSTs and positive ex-
cursions in salinity, together expressing conditions closely
similar to the modern ones. Other proxies reveal GIs marked
by lower IRD inputs, together occurring with reduced ad-
vection of terrigeneous material (i.e. low values of the ratios
$[\text{Ti} + \text{Fe}]/[\text{Ca} + \text{Sr}]$ and K/Ti, also probably linked to a dilu-
tion by an enhanced carbonate biogenic productivity), and a
ratio of Rd/Md close to zero. This suggests reduced calving
from the surrounding ice sheets. The magnetic content is high
at the same time as bottom currents are enhanced, transport-
ing magnetite-rich sediments from the basaltic sill. The re-
constructed sea-ice cover duration indicates GIs mostly sea-
ice free. This is coherent with warmer SSTs recorded during
winter (February) and summer (August) for most of the dura-
tion of these warm events (except within events of larger sea-
ice duration excursions, see below). However, during GIs,
our reconstructions argue for very different sea-surface con-
ditions than the modern ones, especially because strong SST
anomalies (modern versus GIs) are detected (Fig. 5): win-
ters appear to be colder compared to modern ones (−6 °C)
whereas summers record positive anomalies of nearly 4 °C.
These anomalies are strongly reduced during warm winter
excursions that mark the end of the GIs; SSTs thus tend very
briefly to mirror modern conditions. Transient sea-ice cover
development accompanies these peculiar events.

High relative abundances of NPS ($> 80\%$) in the North Atlantic modern sediments mark the presence of the polar front (PF) (e.g. Eynaud et al., 2009). A shift is observed concerning this species in the vicinity of Faeroes with abundances sharply increasing north of these islands and reaching monospecific values in the Greenland Sea. Percentages up to 40% can be recorded in modern sediment of the area. The signal reconstructed in MD99-2281 during MIS3 could thus provide a record of the evolution of the PF position over the area. The pattern identified within this proxy follows the DO oscillations with GIs recording percentages close to their modern values, i.e. conditions which characterise a northern position of the PF. Along the studied section, relative abundances of NPS mirror the evolution of the $^{XRF}$Ca sig- nal (Fig. 6).

GI8

The GI8 records atypical conditions. It is considered as one
of the warmer MIS3 interstadials (Sanchez-Goñi et al., 2008)
and is marked by a gradual cooling in ice records; its occur-
rence in our archive evidences at least two distinct phases.
The reconstructed sea-surface conditions characterise a GI8
recording warm August SST of nearly 16 °C during the first
centuries, followed by a sharp cooling at about 37.4 ka BP,
which instead corresponds to warmer February SSTs and to
an abrupt increase (the largest in our record) in the detrital

material as documented by the LLG concentrations (Fig. 6).
However, it should be noted that this transition is not accom-
panied by any increase in the $[\text{Ti} + \text{Fe}]/[\text{Ca} + \text{Sr}]$ or K/Ti ra-
tios or any change in the $^{Xrf}$Ca or magnetic concentration.
$\delta^{18}$O values indicate rapid minor oscillations within GI8,
with a trend toward higher values recorded precisely when
the LLG peak occurred. We also note that percentages of
NPS are then relatively high compared to the other GIs, with
also high absolute abundances of planktonic foraminifera
shells in the sediment. This terminal phase probably relates
to milder conditions (especially during the winter months, as
detected by dinocysts).

6.2 Stadials (GSs and Heinrich Events HEs)

In our record, all GSs (underlined by grey bands on Fig. 6), except HE3, are characterised by high LLG concentrations and high abundances of NPS, coupled with high values of the ratios Rd/Md, $[\text{Ti} + \text{Fe}]/[\text{Ca} + \text{Sr}]$ and K/Ti. Conversely, both $^{XRF}$Ca content and planktonic foraminiferal abundances significantly decrease. This reflects important terrigeneous advections (Matthiessen et al., 2000; Zaragosi et al., 2001; Richter et al., 2006; Ménot et al., 2006; Penaud et al., 2009). However, there are less magnetic particles than during GIs. This is related to the calving of the proximal ice sheets that does not bring magnetic particles at that time, and is combined with a reduced bottom current, while, during GIs, stronger deep currents bring magnetic particles from the Ice-land plateau. The high percentages of NPS (between 80 and 100%), accompanied by high $\delta^{18}$O values, further indicate a marked southward migration of polar waters.

These discharges are accompanied by SST approaching
0 °C in February. However, a strong seasonality is observed
during these periods as SSTs are still high in August. Such
a strong seasonality occurring during GSs is at the oppo-
site of what is at present characteristic of the area. Modern
sea-surface environments are indeed strongly influenced by
the penetration of the NAD, reducing the SST difference be-
tween winter and summer (3 °C only). Conditions recorded
during MIS3 could then be due to a globally reduced NAD
influence over the site of study. The increase in seasonal-
ity we find could also be related to sea-surface conditions
marked by the presence of proximal ice sheets that may have
maintained cold atmospheric situations during winters. Addi-
tionally, another scenario could imply the development dur-
ing summers of a thin, low salinity surface layer in rela-
tion to the seasonal melting of local ice sheets, which could
be rapidly warmed up during the warm season. This atyp-
cal warm summer signal could also be related to the forma-
tion of polynias, which could have constituted patchy "hot
spots" free of ice. Actually, such phenomena are observed at
present in the modern Arctic Ocean, where strong winds as-
sociated or not to the upwelling of warm waters, create holes
in the sea-ice cover (e.g. Platov, 2011).
## 6.2.1 Heinrich event 4 (HE4)

Known as the largest amplitude cooling event over HEs (e.g. Cortijo et al., 1997), with in particular a very strong Laurentide ice-sheet (LIS) signature, HE4 can be divided into three phases in some parts of the Northeast Atlantic (e.g. Hemming, 2004), with a first European phase, a second longest Canadian phase, and then a terminal European phase (Snoeckx et al., 1999; Grousset et al., 2000). Multiphasing within HE4, with especially a mid-event warming, was recently reported from the central North Atlantic (Reykjanes Ridge, Jonkers et al., 2010a).

In our multiproxy dataset, a typical GSs signature is recorded for HE4, with indices suggesting ice-sheet collapse (high LLG concentrations, low $\delta^{18}$O values). Interestingly, a major peak in the $^{XRF}$Ca content is recorded in the middle of HE4 (around 39 ka). This peak is quite atypical regarding other GSs that otherwise record low $^{XRF}$Ca values. It also differs from interstadial $^{XRF}$Ca maxima with respect to shape, amplitude and duration, as well as in its relationship to proxies of terrigenous input. The $^{XRF}$Ca peak displays a sharp onset and termination but lacks an intermittent “plateau” of roughly constant or slightly variable values; it is inverse V-shaped rather than inverse U-shaped, corresponding to a pulse rather than a period of high $^{XRF}$Ca. The Ca peak during HE4 coincides with a maximum of the K/Ti intensity ratio (highest values of the entire record), while there is little if any relationship between both XRF parameters for the remainder of MIS3. Regarding data of magnetic properties, it finds echoes only in the low field magnetic susceptibility ($\kappa$, Fig. 6), suggesting advection of large-sized magnetic grains. The pulse-like structure combined with the bulk terrigenous and magnetic signatures suggest that this anomaly may reflect an ice-rafting event, hence a detrital origin of carbonates. Relatively low $^{XRF}$Sr with respect to $^{XRF}$Ca provides further circumstantial evidence for this hypothesis, as detrital carbonates are generally depleted in Sr compared to biogenic ones (Hodell et al., 2008). According to Hemming (2004), high proportions of detrital carbonates characterise IRD sourced from Hudson Strait and Laurentide ice sheet. This signal might then have overlapped the European ice sheets one. In this case, changes due to local ice sheets instabilities have probably occurred before the signal that characterises the destabilization of the LIS, as featured in our record. However, this is not sufficient to suggest a precursor role of the European system (e.g. Scourse et al., 2000) for HE4. It must indeed be noted that European ice sheets finally have had the same behaviour for each GSs throughout MIS3, which attests to their sensitivity, with no evidence of asynchronicity with the LIS. Taking into account the distance between the core MD99-2281 and the Laurentide source, few IRD may have reached the area after crossing the North Atlantic, and are therefore only recorded as exceptional outburst events. This is also observed in the proximal core MD04-2829CQ further south (Fig. 1) that only

records LIS-sourced IRD during a 200–300 yr period within HE4 (for the section studied in this work), which might argue for this first hypothesis (Hall et al., 2011). Another possible explanation for the peak of $^{XRF}$Ca could also be related to calcite secondary precipitation under anoxic bottom conditions, as suggested by some authors in the atypical sedimentological facies of HEs (Auffret et al., 1996; Zaragosi et al., 2001).

An interesting feature during HE4, but also during the others GSs, is the high SSTs during summer (August), up to 4 °C more than modern values, contrasting with the percentages of NPS that then reach nearly 100 % (Fig. 6). This could suggest a strong stratification of the water masses but also emphasizes ecological strategy differences between planktonic foraminifera and dinoflagellates, the latter living in the 50 upper meters of the water column while NPS can be mesopelagic (de Vernal et al., 2006; Jonkers et al., 2010b).

## 6.2.2 Heinrich event 3 (HE3)

HE3 is one of the most contentious HEs, especially from a stratigraphic point of view (e.g. Sanchez-Goni and Harrison, 2010), but also because of its atypical low IRD rate (e.g. Bond and Lotti, 1995; Grousset et al., 2001) and its strong European geochemical signature (Snoeckx et al., 1999; Grousset et al., 2000).

Our data reveal a low concentration of IRD during HE3 with two LLG peaks bracketing the event (Fig. 6), at its onset (32 ka BP) and at its termination (29 ka BP). These two low amplitude peaks correspond also to increases in the Rd/Md and $[Ti + Fe]/[Ca + Sr]$ ratios and to minima in the $^{XRF}$Ca, thus reflecting advection of terrigenous material probably related to proximal ice-sheet dynamics. Abundances of NPS follow this bimodal pattern, with two periods marked by quasi-monospecific values. Conversely, dinocyst derived sea-surface conditions show no characteristic excursions at this time, but rather out of phase: sea-ice cover duration increases briefly before the first IRD pulse, and just after the last one, i.e. respectively within GI5 and 4.

A clear rise in the $^{XRF}$Ca content is noted in the second part of HE3, coeval with two minor peaks in magnetic concentration. However, $^{XRF}$Ca values remain under previous GI values. As already introduced for HE4, it is possible that such a rise characterises LIS preferentially detrital-sourced material advection. Bond et al. (1992) hypothesised warmer SSTs in the North Atlantic during HE3, leading to an accelerated melting of icebergs, which thus cannot fully reach the eastern basins of the North Atlantic. This longitudinal “barrier” was also mentioned by Hemming (2004). If so, the late arrival of this collapse from the North American ice sheets could have led to further destabilization of the European ice sheets (second peak recorded by proxies at the transition to GI4).

From the end of HE3, some changes are taking place. Sea-surface conditions seem to record higher amplitude changes between GIs and GSs, as shown for example in the
<figure><img src="image_10.png" /><figcaption>**Fig. 7.** Compilation of comparable MIS3 marine records from the subpolar North Atlantic (see Fig. 1 for the location of the cores), including core SO82-5 (Van Kreveld et al., 2000), core MD99-2281 (this work), core MD95-2009 (Eynaud et al., 2002) and core ENAM93-21 (Rasmussen and Thomsen, 2004). Proxies from the same nature are shown with the same colour. A reconstruction of the MIS3 sea-level (m) is also shown on the left (Rohling et al., 2009; Siddall et al., 2006). Grey bands: same legend as Fig. 6.</figcaption></figure>

percentages of NPS or in the MS record. This is also attested by the reconstructed SSTs with extreme oscillations recorded in August and in the sea-ice duration as well.

## 7 Heterogeneity of the climatic response of the North Atlantic basin during MIS3

Our dataset constitutes one of the first multiproxy datasets detailing MIS3 sea-surface paleoenvironments south of the Norwegian Sea. Some detailed reconstructions of MIS3 exist further north or further south (e.g. Rasmussen et al., 1996a,b, 1997, 2002a,b; Rasmussen and Thomsen, 2004, 2008) but only a few have been using the same multiproxy approach. Figure 7 tentatively compiles high resolution MIS3 data comparable to the dataset generated for MD99-2281. Despite some differences in the chronology (compiled records other than MD99-2281 derive from the GISP2 chronology) and resolution, it illustrates hydrographical gradients, and thus zonal responses within the same subboreal latitudinal band.

MD99-2281 reconstructions could be compared to those previously produced by Eynaud et al. (2002), who used the same set of proxies in two marine cores retrieved from the southern Norwegian Sea (north of the Faeroe Islands – MD95-2009 and on the Voring plateau – MD95-2010, Fig. 1). Intriguingly, Eynaud et al. (2002) observed an antiphase between sea-surface paleoenvironments (sea ice cover duration, February SST) of the Norwegian Sea and the GIs/GSs atmospheric conditions registered above Greenland (e.g. Dansgaard et al., 1993). Their records demonstrated the establishment of low February SSTs accompanied by a longer sea-ice cover duration in the Norwegian Sea during GIs, whereas the Greenland ice-records testify warm interstadial atmospheric conditions (and vice-versa during GSs, see Fig. 7). According to Eynaud et al. (2002), this paradox pointed to a decoupling between the North Atlantic eastern and western parts during MIS3, involving past climatic phenomena comparable to the modern North Atlantic Oscillation (“Greenland above/Oslo below normal”, e.g. Bromwich et al., 1999).
Low concentrations of IRD were observed by these authors during HE4, in sharp contrast with what was observed in the temperate North Atlantic (e.g. Cortijo et al., 1997; Hemming, 2004). It further suggests isolation of the Norwegian Sea in terms of response to the climatic variability.

Similar results were obtained by de Vernal et al. (2000, 2001, 2005, 2006) for the Nordic Seas with positive SST anomalies (SST above the modern ones) detected during the Last Glacial Maximum. The reconstructed August SSTs on core MD99-2281 during MIS3 show comparable positive anomalies (up to 4°C) dominating the whole studied period (Fig. 5). These observations could be due to peculiar oceanic conditions developing during MIS3 and strongly related to the dynamics of the proximal ice-sheets. A strong stratification of the water column could have been favoured by partial melting of the ice sheets and thus advection of fresh waters, preferentially during summers. The existence of a strong stratification could justify a rapid warming of “a brackish superficial water lens” generating such an anomaly.

The anomalous August SSTs testified a “Norwegian Sea type response” for the summer months (see Eynaud et al., 2002). Late increases of February SST during each GIs furthermore underline some kind of inertia in the sea-surface response of the southern Faeroe domain, thus arguing for a potential influence of the Norwegian sector. However, the high concentrations of IRD observed in core MD99-2281 during HE4 (Fig. 6) suggest that our site also responds synchronously to the temperate North Atlantic, and is thus outside the paradoxical “Nordic Seas” type response.

GI6, 5 and 4 are also paradoxically marked by a higher sea-ice cover duration roughly in antiphase with the percentages of NPS. The same anomaly is observed in the northern core MD95-2010 where interstadial phases over Greenland correspond to a longer sea-ice cover development (Eynaud et al., 2002). We can suggest that from GI6 to GI3, the conditions are so drastic, with proximal European ice sheets reaching their critical maximum size that “our region adopts the same behaviour as the Nordic seas do”. The IRD signal recorded in core MD99-2281 during these last GIs could thus correspond to the one also recorded in core MD95-2010 and could therefore be related to the destabilization of the FIS. It is also worth noting that the ice cover in core MD95-2010 varies between 5 and 8 months/year for the GIs, while our core does not record an ice cover exceeding 2 months/year during the whole MIS3. It is conceivable however that from GI6, a change in the local sea-surface conditions may have been initiated due to the gradual development of the proximal ice-sheets.

# 8 The role of European ice sheets on proximal sea-surface paleoenvironments during MIS3

## 8.1 Ice-sheet dynamics: the MIS3 marine record

In the Faeroe Island area, arrivals of detrital material (IRD) to the ocean were interpreted as proximal European ice sheets sourced: their record throughout MIS3 is continuous and highly sensitive to millennial climatic variability as shown by our data and also by previous studies in the same sector (e.g. Scourse et al., 2009; Hibbert et al., 2010; Chiverell and Thomas, 2010; Hall et al., 2011). Continental data from the proximal British Isles have shown “ice-free” conditions during MIS3, as well as in central Scotland, Northern Ireland and the Isle of Man. The BIIS extension was then reduced in latitude at that time and mainly located in the north of Scotland, probably with an advance northward and northwestward (Scourse et al., 2009; Chiverell and Thomas, 2010). This configuration could have promoted the development of an ice shelf, feeding a quite constant icebergs calving and contributing to the observed IRD input record in the area. It is not possible from our data to pinpoint the exact source of IRD; however, the large size of the FIS compared to the British one suggests a preferential Fennoscandian source for IRD. Nevertheless, Hall et al. (2011) pointed to a major BIIS contribution in IRD fluxes recorded in core MD04-2829CQ. This latter being located only 100 km (nearly one degree in latitude) southward of core MD99-2281 (Figs. 1 and 8), and furthermore responding in the same way during GSs, i.e. recording higher IRD concentrations, it is thus likely that BIIS derived products may also have reached the site of core MD99-2281.

## 8.2 The MIS3/MIS2 transition

The MIS3 termination and the beginning of MIS2 are worth discussion considering our results. Indeed, changes are taking place at the transition between HE3 and GI4 at 28.9 cal ka BP where alternating GIs/GSs sea-surface conditions seem to be more extreme, with marked shifts in SSTs (August especially) and in the sea-ice cover duration. Generally, the MIS3 to MIS2 transition marks conditions that seem progressively more severe in the area, also registering higher instability. Even if only observed in the last 2000 yr of the sequence, it is clearly expressed throughout all our proxies and could be due to growing European ice sheets, with maybe the development of ice shelves on the continental shelf. In fact, the maximum ice volume, which marks the Last Glacial Maximum between 19 and 23 cal ka BP (sensu Mix et al., 2001), does not necessarily correspond to the maximum geographical extension of each ice sheet, which also evolved independently (e.g. Chiverrell and Thomas, 2010). Indeed, some previous studies have suggested that the BIS advanced between 30 and 25 cal ka BP. Hibbert et al. (2010) show that the core MD04-2822 (Fig. 8), located
<figure><img src="image_11.png" /><figcaption>**Fig. 8.** BIIS-FIS ice sheets and their associated paleo-ice-stream main directions (black solid line, Bradwell et al., 2008). The full glacial extension of the North Hemispheric ice sheets is schematized by the red dotted-line (Grosswald and Hughes, 2002); Ehlers and Gibbard, 2004). Palaeo-coastline at 21 ka BP in blue (Bourillet et al., 2003). The studied core (MD99-2281: this work) and other cores discussed in the paper (MD95-2009; Eynaud et al., 2002; MD04-2822; MD95-2006; Knutz et al., 2001; Hibbert et al., 2010; MD04-2829CQ; Hall et al., 2011) are also located.</figcaption></figure>

west of the Barra-Donegal Fan, has recorded an increase in IRD at 27.4 cal ka BP. The same is observed by Hall et al. (2011) in core MD04-2829CQ, located southeast of the Rosemary Bank, pointing to an increase of the BIIS size after ~28 ka BP. Similarly, Knutz et al. (2001) observed in core MD95-2006 (Fig. 8), located in the heart of Barra-Donegal Fan, increases of IRD concentrations (coming from the Tertiary basaltic provinces involving a British source) at about 30 ka, which coincide with the MIS3/MIS2 transition (Hibbert et al., 2010). Scourse et al. (2009) also illustrate an advance in the west of the British ice sheet (BIS) at the Bench level and Rosemary Barra-Donegal Fan just after HE3, at about 29 ka. IRD concentration increases at 27 ka on the Irish western margin and at 25 ka on the Goban Spur, characterising a progressive southward ice-sheet development (Scourse et al., 2009). This progression also affected the northeastern British ice margin. Actually, geomorphological studies conducted in the North Sea Basin and Norwegian Channel (Bradwell et al., 2007, 2008) have suggested confluent British and Fennoscandian ice sheets during this period. This coalescence associated to a major paleo-ice-stream system purging the “ice-bridge” (Fig. 8, Bradwell et al., 2008) might have alimented one of the most important ice shelves of the European margin. Core MD99-2281, which is located in the axis of two of these ice streams (Fig. 8, Bradwell et al.,

2008), the Minch paleo-ice stream and the North Sea paleo-ice stream, may have thus recorded destabilizations of this system.

## 9 Conclusions

The present multiproxy study, conducted on the MIS3 section of core MD99-2281 located off SW Faeroes, provides a new and unique set of data, expanding our knowledge about the evolution of NW Europe oceanic palaeoenvironments during the last glacial. Our results demonstrate that the marine reservoir echoes directly to the rapid climate changes detected in the atmosphere (Greenland $\delta^{18}$O records). Deep-sea parameters (deduced from magnetic properties) show the same general pattern as along the main path of the NADW in the subpolar North Atlantic (e.g. Kissel et al., 1999a). These parameters were therefore used to tie the surface multi-proxy records on the millennial climatic variability that characterises the MIS3. Results demonstrate that the surface oceanic response of the area was strongly influenced by the dynamics of the proximal ice sheets. Data identify a much more important calving of European ice sheets during GSs and milder conditions during GIs, with however late and terminal warmings (rather occurring at the transition towards GSs). Large amounts of detrital material (including IRD) are recovered for each GSs and are likely to have a European origin, and more precisely a BIS source. Sea-surface paleoenvironmental conditions during HEs are quite similar to those recorded during other GSs, except for some indices of a probable larger Laurentide influence (both within HE4 and HE3).

Marked paleoenvironmental changes seem to occur after 28.9 cal ka BP (post-HE3 record). This reinforces previous findings which hypothesise the existence of an ice bridge between the British and the Fennoscandian ice sheets, which would have potentially directly influenced our site of study after HE3. The coalescence of the FIS and BIS could have fed an ice shelf proximal to the site of core MD99-2281; the high variability of sea-surface palaeoenvironments detected in our data could support the presence of such a structure. This implies that the glacial advance of the British ice sheet toward its maximum extension was probably reached before the Last Glacial Maximum. Further studies are needed to better constrain this late history of the British ice sheet and its relation to the dynamics of its surrounding ice shelves.

*Acknowledgements.* The authors are grateful to IPEV, the captain and the crew of the *Marion-Dufresne* and the scientific team of the GINNA cruise (Labeyrie et al., 1999). We wish to thank Y. Balut for his assistance at sea. J. St Paul, O. Ther, M.-H. Castera and M. Georget provided invaluable technical assistance. We acknowledge financial support by the ARTEMIS $^{14}$C Accelerator Mass Spectrometry French project. Part of the analyses conducted on MD99-2281 was supported by the French INSU (Institut
National des Sciences de l'Univers) program LEFE/EVE (Les enveloppes fluides et l'environnement/Evolution et variabilité du climat à l'échelle globale) "RISCC: Rôle des Ice-Shelves dans le Changement Climatique". The magnetic analyses at LSCE were made with the help of Camille Wandres from laboratory funds from the French Atomic Energy Commission and from the CNRS. This is an U.M.R./EPOC 5805 (Université Bordeaux I – C.N.R.S.) contribution.

Edited by: M. Siddall

<figure><img src="image_12.png" /></figure>

INSU
Institut national des sciences de l'Univers

The publication of this article is financed by CNRS-INSU.

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