Quaternary Science Reviews 194 (2018) 128e142

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Quaternary Science Reviews
journal homepage: www.elsevier.com/locate/quascirev

Subpolar North Atlantic sea surface temperature since 6 ka BP:
Indications of anomalous ocean-atmosphere interactions at 4-2 ka BP
Lisa Claire Orme a, *, Arto Miettinen a, Dmitry Divine a, b, Katrine Husum a,
Christof Pearce c, Nicolas Van Nieuwenhove d, Andreas Born e, f, Rahul Mohan g,Marit-Solveig Seidenkrantz caNorwegian Polar Institute, Fram Centre, P.O. Box 6606 Langnes, 9296, Tromsø, Norway
Department of Mathematics and Statistics, Faculty of Science and Technology, The Arctic University of Norway, N-9037, Tromsø, Norway
Centre for Past Climate Studies, Arctic Research Centre and iClimate, Department of Geoscience, Aarhus University, Høegh-Guldbergs Gade 2, building
1672, 213, 8000, Aarhus C, Denmarkd
Department of Earth Sciences, University of New Brunswick, 2 Bailey Drive Fredericton, NB, E3B 5A3, CanadaeDepartment of Earth science, University of Bergen, Postboks 7803, 5020, Bergen, Norwayf
Bjerknes Centre for Climate Research, University of Bergen, Postboks 7803, 5020, Bergen, NorwaygNational Centre for Antarctic and Ocean Research, Ministry of Earth Sciences, Government of India, Headland Sada, Vasco-da-Gama, Goa, 403 804, Indiabca r t i c l e i n f o

a b s t r a c t

Article history:
Received 2 May 2018
Received in revised form
20 June 2018
Accepted 2 July 2018

Atmospheric circulation may change with future climate change in response to modiﬁcation of meridional temperature gradients, but the potential inﬂuence on ocean circulation is as yet unclear. Over the
mid-late Holocene, atmospheric circulation in the North Atlantic region has ﬂuctuated on millennial
timescales; therefore, the ocean response to these changes can be investigated using the paleoceanographic records that have been developed in the north-eastern subpolar North Atlantic. Here, we present
a diatom-based sea surface temperature reconstruction from the Iceland Basin, south of Iceland; thereconstruction shows the warmest temperatures of the record at 6.1e4 ka BP, cooler temperatures at 4-2
ka BP and warmer temperatures thereafter. Inter-record comparisons indicate that the cold period at c. 42 ka BP may have resulted from a strengthened East Greenland Current and/or melting of the Greenland
ice sheet, in response to a negative North Atlantic Oscillation. The ﬁndings highlight that atmosphericcirculation changes are likely to cause pronounced variations in the latitudinal exchange of heat, which
may have consequences for deep-water formation and global ocean circulation.
© 2018 Elsevier Ltd. All rights reserved.

Keywords:
Holocene
Paleoceanography
Paleoclimatology
North Atlantic
Micropaleontology
Diatoms
Sedimentology-marine cores
North Atlantic oscillation
East Greenland current
Sea surface temperature

1. Introduction
The dominant mode of variability in the atmospheric circulationof the North Atlantic region is the North Atlantic Oscillation (NAO).
This index, calculated as the pressure gradient between the Subpolar Low, centred over Iceland, and the Azores High, controls thelatitudinal position of storm tracks and wind strength in the North* Corresponding author.
E-mail addresses: Lisa.Orme@npolar.no (L.C. Orme), Arto.Miettinen@npolar.no(A. Miettinen), Dmitry.Divine@npolar.no (D. Divine), Katrine.Husum@npolar.no(K. Husum), christof.pearce@geo.au.dk (C. Pearce), nicolas.vannieuwenhove@unb.
ca (N. Van Nieuwenhove), Andreas.born@uib.no (A. Born), rahulmohangupta@
gmail.com (R. Mohan), mss@geo.au.dk (M.-S. Seidenkrantz).
https://doi.org/10.1016/j.quascirev.2018.07.007
0277-3791/© 2018 Elsevier Ltd. All rights reserved.

Atlantic region (e.g. Hurrell and Van Loon, 1997). During Northern
Hemisphere summer, the NAO generally weakens, with lower
amplitude variability and a northward shift of the pressure cells,
while they move southward and display a higher amplitude ofvariability in winter (Folland et al., 2009). A positive NAO (with thegreatest difference between the pressure dipoles) is associatedwith a northward-shifted storm track, driving a stronger andnorthward-shifted North Atlantic Current (NAC) and potentially agreater E-W extension of the subpolar gyre (SPG; Taylor andStephens, 1998; Curry and McCartney, 2001; Frantantoni, 2001),
although this latter response has been questioned (Foukal and
Lozier, 2017). Weaker and southward-shifted storm tracks occur
during negative NAO phases (characterised by a reduced pressure
difference), resulting in a weaker and southward-shifted NAC

L.C. Orme et al. / Quaternary Science Reviews 194 (2018) 128e142(Taylor and Stephens, 1998; Curry and McCartney, 2001).
During a positive (negative) NAO, over sub-annual timescales,
the response to air-sea heat ﬂuxes and wind-driven Ekman transport is cooling (warming) in a zonal band spanning the North
Atlantic north of 45 N (Kushnir, 1994; Seager et al., 2000; Visbeck
et al., 2003). However, over multi-annual to decadal/centennial
timescales it is suggested that a positive (negative) NAO causes
increased (decreased) convective activity in the Labrador Sea andstrengthening (weakening) of the SPG and meridional overturning
circulation, resulting in warming (cooling) north of 55 N (Eden and€kkinen and Rhines, 2004; Latif
Jung, 2001; Visbeck et al., 2003; Haet al., 2006). During a negative NAO strengthened northerly
winds to the east of Greenland can reinforce the East Greenland
Current (EGC) and increase the export of sea ice and freshwaterfrom the Arctic to the North Atlantic, a scenario which in thetwentieth century caused ‘Great Salinity Anomalies’ (Dickson et al.,
1996: Delworth et al., 1997; Belkin et al., 1998; Blindheim et al.,
2000; Ionita et al., 2016). Similar episodes have been identiﬁed
over decadal-centennial timescales in model and paleoclimate
analyses (Delworth et al., 1997; Renssen et al., 2005; Sicre et al.,2008; Ran et al., 2011). Nevertheless, the interactions between
the ocean and atmosphere are complex. Models indicate that, due
to the coupling of the ocean and atmosphere system, sea surface
temperature (SST) changes themselves may force NAO variability.
Here, potential feedbacks include changes in heat transport by thesubtropical gyre (STG) and the SPG (Bellucci et al., 2008) as well aschanges in evaporation, precipitation and atmospheric heatingprocesses (Rodwell et al., 1999).
The NAO is also negatively correlated with atmospheric temperatures over the Greenland ice sheet and is therefore a potentially important factor controlling past and future melting (Serreze
et al., 1997; Chylek et al., 2004; Fettweis et al., 2013). The melting of
the Laurentide ice sheet in the early Holocene is known to havelowered SSTs in this region (Andersen et al., 2004b; Berner et al.,2008; Blaschek and Renssen, 2013; Jiang et al., 2015; Sejrup et al.,2016), however there has been less research on the ocean temperature response to ﬂuctuations in the extent of the Greenland ice
sheet during the mid to late Holocene (Briner et al., 2016).
Compilations of SST reconstructions from the subpolar North
Atlantic show that during the mid-late Holocene the dominantfeature in the region was cooling (e.g. Calvo et al., 2002; Rimbu
et al., 2003; Sundqvist et al., 2014; Sejrup et al., 2016), which was
primarily thought to result from orbital forcing (e.g. Andersen et al.,2004a, 2004b; Renssen et al., 2005; Hald et al., 2007; Berner et al.,2008; Kissell et al., 2013; Jiang et al., 2015; Sejrup et al., 2016).
However, reconstructions indicate differences between regions, as
there is also some evidence for a warming after 2 ka BP in the
Iceland and Greenland Seas (e.g. Justwan et al., 2008; Miettinen
 ski et al., 2014; Moossen et al., 2015; Sejrup
et al., 2012; Telesin ttir et al., 2017). Thus, understanding theet al., 2016; Kristj
ansdo
past oceanographic changes is complex due to differences between
reconstructions from the same region and even the same site,which may result from real spatial patterns but also factors related
to the speciﬁc proxy used for the different SST reconstructions(climate-sensitivity, seasonal growth and habitat water-depth, aswell as the statistical methods used).
Our understanding of the NAO variability during the Holocene isbased on just a few records. Trouet et al. (2009) and Faust et al.
(2016) suggest that the Medieval Climate Anomaly (~950e1250
A.D.) had a persistent positive NAO and the Little Ice Age (1250/
1400e1850 A.D.) a negative NAO, and an imprint of these changes
has been found in some marine reconstructions (e.g. Abrantes et al.,2005; Seidenkrantz et al., 2007, 2008; Sicre et al., 2014). Longer
NAO reconstructions indicate that beginning at 6e4.5 ka BP until 2
ka BP the NAO was more neutral and negative, which was followed

129

by a more frequently positive NAO after 2 ka BP (Olsen et al., 2012;
Nesje et al., 2000; Faust et al., 2016), a trend supported by acompilation of SST reconstructions (Rimbu et al., 2003). The validityof basing NAO reconstructions on single sites has been questioned
due to non-stationarity in the proxy-NAO relationship through time(e.g. Lehner et al., 2012; Ortega et al., 2015). However, the NAO
reconstructions are supported by storm track reconstructions from
Europe, which show a southward storm track position (reﬂecting amore negative NAO) before 2 ka BP, followed by a northward storm
track shift (reﬂecting a more positive NAO pattern) after 2 ka BP(Bakke et al., 2008; Orme et al., 2017). These results therefore
suggest that through the Holocene there have been long-term
millennial shifts in atmospheric circulation.
The CMIP5 multi-model ensemble predicts that up until the endof the 21st century, storm tracks will stay within the range ofnatural variability in the northern hemisphere, although there issubstantial uncertainty in this prediction (Collins et al., 2013).
However, some studies based on modelling, observational andpalaeoclimate records (Francis and Vavrus, 2012; Yang andChristensen, 2012; Kim et al., 2014; Orme et al., 2017) suggestthat greater warming over the polar regions compared to that of the
mid-latitudes could weaken the temperature gradient, driving aweaker meridional circumpolar circulation and a negative NAO.
Such negative NAO circulation may result in changes in the subpolar North Atlantic, by reducing the amount of wind-driven
northward heat transport, increasing the southward export of
Arctic water and altering ocean-atmosphere heat ﬂuxes, which may
then alter rates of deep-water formation and the wider ocean circulation (e.g. Taylor and Stephens, 1998; Blindheim et al., 2000;
Curry and McCartney, 2001; Eden and Jung, 2001; Latif et al., 2006).
Thus, a central aim of this paper is to investigate the changes to theNorth Atlantic Ocean circulation that occurred as a result of thepossible long-term negative phase of the NAO at 4-2 ka BP (Olsenet al., 2012; Faust et al., 2016). As such, here we present a new
diatom-based SST reconstruction from the Iceland Basin for the last
6.1 ka BP and compare this with other SST reconstructions from the
subpolar North Atlantic to understand the spatial temperaturepatterns. We then compare these with past variations in atmospheric circulation, to better understand the long-term ocean-atmosphere interactions in a region of high importance for the global
climate system.
2. Oceanographic setting
The core site DA12-11/2-GC01 is located in the Iceland Basin
which is bounded by the Reykjanes Ridge to the west, Iceland andthe Iceland-Faroe ridge to the north and northwest and the Rockall
Plateau to the southwest (Fig. 1). Within the Iceland Basin sediment€rntransported by deep currents has accumulated to form the Bjo
and Gardar Drifts (Bianchi and McCave, 2000).
The NAC transports warm Atlantic water to the subpolar North
Atlantic along several pathways (Fig. 1). A western branch turns
northwest towards Iceland; part of the current passes over the
Iceland-Faroe Ridge into the Nordic Seas while the other part
travels as the Irminger Current (IC) ﬁrst to the south and then to thenorth around the Reykjanes Ridge (Jakobsen et al., 2003; Pollard
et al., 2004). The IC bifurcates in the Denmark Strait with a smallpart ﬂowing to the north of Iceland while a larger branch recirculates to the south and converges with the south ﬂowing EGC(Jakobsen et al., 2003; Pollard et al., 2004; Våge et al., 2011). The
eastern branch of the NAC remains close to the European continental slope as it heads northwards, passing through the Rockall
Trough to the Faroe-Shetland channel before entering the Nordic
Seas (Orvik and Niiler, 2002; Jakobsen et al., 2003). The two currents that entered the Nordic Seas head northwards along separate

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Fig. 1. Location of core DA12-11/2-GC01 in the subpolar North Atlantic. Left: Core location and main surface currents (continuous lines) and fronts (dashed lines). EGC: East
Greenland Current, EIC: East Icelandic Current, WGC: West Greenland Current, LC: Labrador Current, NAC: North Atlantic Current, NwAC: Norwegian Atlantic Current, IC: Irminger Current,
FC: Faroe Current, BC: Bafﬁn Current. Shaded regions represent approximate geographic distribution of the modern diatom assemblages (factors) in the subpolar North Atlantic
according to Andersen et al. (2004a) (updated by Oksman et al., 2017): light blue Arctic Greenland assemblage, orange North Atlantic assemblage, purple sub-Arctic assemblage,
green Norwegian Atlantic Current assemblage. Right: Bathymetric map of the Iceland Basin, adapted from Bianchi and McCave (2000). (For interpretation of the references to colourin this ﬁgure legend, the reader is referred to the Web version of this article.)paths, with an area of 200e600 km between them occupied by
Atlantic Water (Orvik and Niiler, 2002). The warm water of the NAC
is transported to Svalbard, where it forms the West Spitsbergen
Current (Aagaard and Greisman, 1975).
The northward transport of warmer Atlantic water is balancedby the southward transport of cold, low salinity water (polar surface water) and sea ice from the Arctic Ocean, carried by the EGC,
which passes through the Fram Strait and continues southwards
through the Denmark Strait (Aagaard and Carmack, 1989; Martin
and Wadhams, 1999; Woodgate et al., 1999; Rudels et al., 2002).
The Polar Surface Water of the EGC extends from the surface todepths ~150 m, with temperatures ~1.7  C and low salinity (34‰)(Aagaard and Coachman, 1968; Jeansson et al., 2008).
The SST distribution in the Iceland Basin is characterised by agradient, with warmer temperatures towards the southwest andcooler temperatures towards Iceland and the Reykjanes Ridge (Deson et al., 2010). The surface water masses are typically
Boisseassociated with Subpolar Mode Water, which has temperature andsalinity varying in the range of 7e12  C and 35e35.5‰ respectively
and results from mixing of subpolar and subtropical waters
advected to the Iceland Basin by the NAC (Wright and Worthington,
son et al., 2012).
1970; McCartney and Talley, 1982; De Boisse

3. Materials and methods
3.1. Core location
The 419.5 cm long marine sediment core DA12-11/2-GC01 wasrecovered in September 2012 by the Danish R/V Dana during theDA12-11/2 teaching cruise, using a gravity corer. The core was takenfrom a location at the northern edge of the Gardar Drift and the€rn Drift in the Iceland Basin (61.36.536  N,
eastern edge of the Bjo
20.42.164  W; Fig. 1) at a water depth of 2120 m. Due to bad
weather, it was not possible to retrieve a CTD depth proﬁle at thesite, but in the upper 50 m of the water column the average temperature at the site for the period 1955e2012 was 10.9 ± 1  C insummer and 8.1 ± 0.8  C in winter, while the salinity was
35.15 ± 0.08‰ in summer and 35.23 ± 0.04‰ in winter (WOA2013;
Locarnini et al., 2013; Zweng et al., 2013).

3.2. Age-depth model
The age-depth model of the full core is based on 10 AMS 14C
radiocarbon dates of either planktic foraminiferal tests (Globigerina
bulloides, Turborotalita quinqueloba) or shell fragments (holoplanktic thecosome pteropod Diacria trispinosa) (Van
Nieuwenhove et al., 2018). The results from the upper 2 m of core
presented here (Fig. 2) are constrained by six radiocarbon dates,
including date AAR24460 from 235 cm, which has a modelled ageof 7540 cal yrs BP (Van Nieuwenhove et al., 2018). Samples weredated at the Aarhus AMS Centre of the Department of Physics andAstronomy, Aarhus University. The radiocarbon ages were calibrated and the age-depth model developed with OxCal 4.2 (Bronk
Ramsey, 2009) using the Marine13 calibration curve (Reimer
et al., 2013), with a regional marine reservoir age offset DR ¼ 0
years (Van Nieuwenhove et al., 2018). Four pronounced tephra
layers could not be attributed to eruptions unambiguously so they
were not used to constrain the age-depth model, and the layers

L.C. Orme et al. / Quaternary Science Reviews 194 (2018) 128e142

Fig. 2. Age-depth model for the upper 210 cm of the DA12-11/2-GC01 core. The calibrated median age is shown as the dark blue curve, the blue shading shows the 1sigma probability interval. Radiocarbon dates are plotted as probability distributions
and the AMS laboratory codes are listed for each date. The dashed line around 35 cm
depth shows a tephra layer, which is assumed to represent an instantaneous deposit(Van Nieuwenhove et al., 2018). (For interpretation of the references to colour in thisﬁgure legend, the reader is referred to the Web version of this article.)

containing ash were excluded from the age model assuming an
instantaneous sedimentation. As the surface sediment was not
dated, the ages for the interval 0e21 cm depth were estimated byinterpolation between the upper radiocarbon date and the coring
year 2012 A.D.. The development of the age-depth model is
described in full by Van Nieuwenhove et al. (2018).
3.3. Diatom analysis and SST reconstruction
The uppermost 210 cm of the core representing the last 6.1 ka
cal BP (Fig. 2) was analysed for diatoms at 5 cm intervals giving an
average temporal resolution of 145 years. The diatoms were isolated from the sediment sample following the methods of Koç et al.
(1993). The process involves the removal of the organic and carbonate material from the sample through treatment with HCl andH2O2, removal of the clay fraction and mounting of the remaining
sample onto slides. Analysis of the diatom assemblages was carriedout on a Leica Orthoplan microscope using 1000 magniﬁcationand followed the counting protocol outlined by Schrader andGersonde (1978). A minimum of 300 diatom frustules were identiﬁed in each sample and the percentage abundance of each sample
calculated. The 300 counted frustules did not include Chaetoceros
spores, which are highly abundant and insensitive to SST (KoçKarpuz and Schrader, 1990).
Diatoms in the North Atlantic bloom during the spring-summer
season; therefore, the species variability represents summer conditions and particularly August temperature (Koç-Karpuz andSchrader, 1990; Birks and Koç, 2002; Berner et al., 2008). Past
SSTs are reconstructed using transfer functions, which establish therelationship between diatom species assemblages present in thesurface (modern) ocean sediments and the known, modern SST
from the sample locations. Individual diatom species have a
unimodal relationship with temperature (i.e. a maximum

131

abundance at an optimum temperature), but SST can be predictedbecause the optimum temperatures of different species are
distributed along a linear temperature gradient (Birks, 1998; Berner
et al., 2008). The direct relationship between the upper watertemperature and surface diatom community and the preserved
fossil community at a given location is complicated by the potential
for lateral transport of sinking diatoms. Individual diatoms or aggregates sink relatively slowly (0.1e100 m/day), so may be carried
long distances, while those within faecal pellets sink quickly(5e2700 m/day; Turner, 2002; Sarthou et al., 2005; Miklasz and
Denny, 2010), so deposition is more direct. However the contribution of each to the diatoms on the seaﬂoor is generally unknown
and spatially variable (Turner, 2002). Nevertheless, the wellestablished strong relationship between the summer temperatures and fossil diatom assemblages in surface sediments supportsthat in general diatom deposition in the subpolar North Atlantic
occurs locally.
The relative diatom abundances were converted to summer(August) SST using the North Atlantic diatom calibration data set
consisting of 183 surface sediment samples (Miettinen et al., 2015)
and a weighted averaging partial least squares (WA-PLS) transfer
function (ter Braak and Juggins, 1993). Using WA-PLS, the fossil
down-core diatom assemblages can then be used to estimate thepast SST (e.g. Birks and Koç, 2002). The WA-PLS transfer function
utilizes the unimodal species response model and attempts tomaximize the covariance between the modern diatom assemblages
and SST, in order to maximize the predictive power of the model(Birks and Koç, 2002) and this requires the number of componentsthat the model will use to be speciﬁed. The optimal number of WAPLS components was determined using a randomization t-test
applied to h-block prediction errors, which showed that the 2component model is the most reliable (Oksman et al., 2017). The
proxy error for the SST reconstruction was estimated using the hblock cross validation technique (Trachsel and Telford, 2016), amethod which accounts for spatial autocorrelation (and theresulting underestimation of error) by excluding from the calibration dataset samples closer than a certain distance from the coresite (Oksman et al., 2017). The root mean square prediction error(RMSEP) for the 2-component WA-PLS model with the h-blockcross validation is 1.1  C, R2 (the coefﬁcient of determination between observed and inferred values) is 0.92 and maximum bias is
2.8  C.
Factor analysis was conducted to characterise past oceanographic conditions at the core site in terms of variations inprevalent types of surface water masses. Q-mode factor analysis
on the modern diatom calibration dataset from the North Atlantic
has grouped diatom taxa into a number of assemblages (or factors) that are characteristic of different water masses (Imbrie andKipp, 1971; Andersen et al., 2004a; Oksman et al., 2017). These are
the Arctic Greenland Assemblage (factor 1), North Atlantic
Assemblage (factor 2), sub-Arctic Assemblage (factor 3), Norwegian Atlantic Current Assemblage (factor 4), Sea Ice Assemblage(factor 5), Arctic Assemblage (factor 6), East and West Greenland
Current Assemblage (factor 7), and the Mixed Water Masses
Assemblage (factor 8) (Andersen et al., 2004a). This classiﬁcationof assemblages has since been modiﬁed by Oksman et al. (2017)
who used the extended modern core top calibration data set(183 samples rather than 139) and renamed the Sea Ice assemblage as the Marginal Ice Zone assemblage. This modiﬁed classiﬁcation of assemblages was applied to the DA12-11/2-GC01
diatom record to show the relative importance of each factor, andtherefore each water mass, through time. The communality isused to assess the agreement between the factor analysis andfossil assemblages. It is deﬁned as a squared sum of factor loadings, which is actually a fraction of variance represented by the132

L.C. Orme et al. / Quaternary Science Reviews 194 (2018) 128e142

decomposition for every single sample analysed. Ideally it should
be 1, but in reality it is generally lower.
4. Results
The diatom species assemblage of DA12-11/2-GC01 is dominated by common North-Atlantic temperate species Thalassionema
nitzschioides, Rhizosolenia hebetata f. semispina, Rhizosolenia borealis, Proboscia alata, Thalassiosira oestrupii, Nitzschia bicapitata andSynedra sp. (each >5% abundance) (Fig. 3). The reconstruction
suggests an average summer SST of about 12  C (range 11.2e13.3  C)
(Fig. 4).
The reconstruction shows a weak negative SST trend throughthe whole record with a cooling of around 0.1  C/1000 years (Fig. 4).
This was likely associated with decreasing abundances of warm
species (Thalassiosira oestrupii, Thalassionema nitzschioides andperhaps Thalassiosira angulata, which have optimum temperaturesof 13.4, 11.1 and 9  C respectively), and increasing abundances ofcool species (Rhizosolenia hebetata f. semispina and Thalassiosira
gravida, which have optimum temperatures of 8.4 and 7.1  C
respectively) (Fig. 3).
The millennial-scale variations in temperature are pronounced
and show three distinct periods, with cooler SSTs at ~4-2 ka BP andwarmer SSTs before (~6.1e4 ka BP) and after (2-0 ka BP) this interval (Fig. 4). Given the age model uncertainties, the inferred cold
period could span up to 2.4 ka, covering the time interval of 4.2 to
1.8 ka BP. The earliest warm period at ~6.1e4 ka BP had averagereconstructed SSTs of 12e13.3  C, with the warmest temperaturesin the record occurring at ~6.1e5.5 ka BP (c. 13  C). The cooler
period ~4-2 ka BP had reconstructed SST that varied around 11.5  C,
with minima at 3.2 and 2.4 ka BP interrupted by a short warming at
2.7 ka BP. In the most recent period after 2 ka BP the SSTs again
increased peaking at 1.8 ka BP, yet SSTs did not attain values as highas those reconstructed for 6.1e4 ka BP. The low SSTs calculated for

4-2 ka BP are linked to a reduced abundance of the warm Synedrasp. (optimum temperature 12.7  C), and increased abundance ofspecies that prefer cooler waters (Thalassiotrix longissima, Thalassiosira nodulolineata, Rhizosolenia hebetata f. semispina andAsteromphalus robustus; optimum temperatures of 8.5, 7.0, 8.4 and5.1  C respectively) (Fig. 3). After 2 ka BP the increased abundanceof Synedra species and reduced abundances of those cooler water
species support a return to warm SSTs.
The signiﬁcance of the temperature change at 4-2 ka BP
compared to the rest of the record was tested statistically using a 2sample Students t-test. This tested the null hypothesis that theresults originated from two normal distributions with equal means
and non-equal unknown variances. Prior to testing, the reconstructed SST values grouped into these two subsets were perturbed
with an additive noise term of N (0,RMSEP). In total, 2000 perturbed reconstructions were generated. The test statistics for eachpair member were then collected and analysed. Results demonstrated a rejection rate for the null hypothesis of about 44%, suggesting that we can reject the hypothesis that the average SST at 4-2
ka BP was equal to the average SST over the rest of thereconstruction.
The factor analysis showed that three diatom assemblages were
dominant in the fossil data set; the Norwegian Atlantic Current
assemblage, the sub-Arctic assemblage and the North Atlantic
assemblage (Fig. 5). The down core distribution of the dominantfactors reﬂects the time variability of the reconstructed SST and
presumably the relative contributions of different water masses (asshown in Fig. 1) at the core location. The Norwegian Atlantic Current (factor 4) showed increasing values at 6.1e4.9 ka BP, constantvalues until 2 ka BP before slightly reduced and variable values after
2 ka BP, with a prominent minimum at 0.1 ka BP. The sub-Arctic
assemblage (factor 3) shows an increasing trend through theentire record from 6.1 ka BP to present, with elevated values at 4-2,
1 and 0.1 ka BP (Fig. 5). The North Atlantic Assemblage (factor 2)

Fig. 3. Diatom abundances as percentages of total sample. Note that the species are plotted on different x-axis scales. The shaded area highlights the reconstructed cold SST interval
at c.4-2 ka BP.

L.C. Orme et al. / Quaternary Science Reviews 194 (2018) 128e142

133

Fig. 4. Temperature reconstructions for DA12-11/2-GC01. Left: Diatom-based August SST reconstruction (this study). Right: Dinocyst-based summer (JulyeSeptember) SST
reconstruction, including raw data (dotted line) and smoothed data (ﬁve point weighted running average; solid line) (Van Nieuwenhove et al., 2018). The shaded area highlights thereconstructed cold SST interval at c. 4-2 ka BP of this study and the dashed line the modern summer SST at this site of 10.9  C (WOA2013; Locarnini et al., 2013).

shows a persistent negative trend between 6.1 and 2 ka BP, thereafter remaining constant (Fig. 5). High values of the communality
index before 2 ka BP support that the fossil assemblages weresimilar to the modern assemblages used for the factor analysis(Fig. 5). The decrease after this point indicates that during the last2000 years the fossil assemblages were not as strongly similar to
those found in the modern assemblages. We speculate that thismay be due to mixing of different water masses after 2 ka BP or due
to relatively few modern samples from the Iceland Basin in themodern calibration dataset.

5. Discussion
5.1. Sea surface temperature in the Iceland basin
The millennial trends in SST observed in the diatom-based
August SST reconstruction presented here are supported by asummer SST reconstruction based on dinoﬂagellate cysts from thesame core (Fig. 4; Van Nieuwenhove et al., 2018), which also showsa cold interval between 4 and 2 ka BP. In the diatom-based record
the mean reconstructed temperature between 4 and 2 ka BP is134

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Fig. 5. Results of Q-mode factor analysis showing the time variability of factor loadings for the three assemblages representing most of the variance in the observed fossil diatoms
sequence. From left: factor 4 (Norwegian Atlantic Current assemblage), factor 2 (North Atlantic assemblage), factor 3 (Sub-Arctic assemblage). Communality, or the fraction of total
variance accounted for by factor decomposition, is displayed in the right panel. See Fig. 1 for a summary of the modern spatial distribution of these assemblages. The shaded area
highlights the reconstructed cold SST interval at c. 4-2 ka BP of this study.

11.5  C compared with 12.5 and 12.1  C in the periods before 4 ka BP
and after 2 ka BP respectively, showing a reconstructed cooling of
0.6e1  C. In the dinocyst-based record the mean reconstructed
temperature between 4 and 2 ka BP is 10.3  C compared with 11.3
and 11.6  C in the periods before 4 ka BP and after 2 ka BP
respectively, showing a cooling of 1e1.3  C. Both records agree thatthe coldest conditions occurred around ~3.2 ka BP (Fig. 4), althougha second SST minimum at ~2.4 ka BP evident in the diatom
reconstruction is represented in the dinocyst-based reconstruction
only by a subtle interruption of a general warming trend.
The diatom-based reconstruction shows warmer reconstructed
temperatures than the dinocyst-based reconstruction and themodern measured summer SST (JuneeAugust) of 10.9  C. The
reconstructed temperatures are similar after 3 ka BP but differ by asmuch as 2  C before this. Given the diatom transfer function RMSEP
of 1.1  C and the dinocyst RMSE of 1.7  C (Van Nieuwenhove et al.,2018) the difference may have been caused by this uncertainty. A
further reason may be because the diatom transfer function reconstructs the peak summer warming (the strongest statistical
relationship is between diatom assemblages and August temperature; Koç-Karpuz and Schrader, 1990; Berner et al., 2008), whereas
the dinocyst transfer function used here reconstructs the averagesummer temperature from JulyeSeptember (Van Nieuwenhove
et al., 2018). Alternatively, there may be some differences in thewater depths that diatoms and dinoﬂagellates occupy.
The overall long-term cooling trend in the diatom-based SST
reconstruction for the last 6.1 ka ﬁts with the widely establishedcooling in the subpolar North Atlantic since the Holocene Thermal
Maximum, resulting from decreasing Northern Hemisphere summer insolation (e.g. Calvo et al., 2002; Marchal et al., 2002;
Andersen et al., 2004a, 2004b; Andersson et al., 2010; Jiang et al.,2015; Sejrup et al., 2016). Since 6.1 ka BP there was also a gradualincrease in the Sub-Arctic assemblage (typical of the central andwestern north Atlantic including the SPG) and a gradual decrease in
the North Atlantic assemblage (typical of the eastern North Atlantic
and the NAC). Given the modern spatial distribution of these water
masses (Fig. 1) this may support a south-eastward expansion ofcooler water masses. The SST cooling registered between 4 and 2 ka
BP represents a remarkable deviation from the overall negative
summer SST trend and therefore this interval will be the focus fordiscussion in the following sections. We ﬁrst summarise the spatial
pattern of SST in the subpolar North Atlantic and then discusspossible driving factors behind the observed variability at 4-2 ka BP.
5.2. Oceanographic changes at 4-2 ka BP
In the Iceland Basin (core Rapid 21-COM) and Reykjanes Ridge
area (core LO09-14, Figs. 6G and 7 (point 6)) SST reconstructions
based on diatoms and dinocysts show the cool event at 4-2 ka BP(Miettinen et al., 2012; Berner et al., 2008; Solignac et al., 2008),
although an alkenone-based SST reconstruction (core MD95-2015)shows only minimal cooling at this time (Fig. 7 (point 9); Marchal
et al., 2002). However, Mg/Ca reconstructions from the Iceland
Basin (MD99-2251 and RAPiD-12-1K), which reﬂect winterL.C. Orme et al. / Quaternary Science Reviews 194 (2018) 128e142

135

Fig. 6. Comparison between a selection of the considered subpolar North Atlantic temperature reconstructions from each region, with reconstructions from the west on the left and ski et al., 2014), B) planktic foraminifera
those from the east on the right. A) planktic foraminifera subsurface summer temperature reconstruction from core PS1878 (Telesin
subsurface summer temperature reconstruction from core MD99-2322 (Jennings et al., 2011), C) diatom-based August SST reconstruction from core MD99-2269 (Justwan et al.,nsdo
 ttir et al., 2017), E) alkenone-based summer SST reconstruction from core MD99-2266
2008), D) alkenone-based summer SST reconstruction from core MD99-2269 (Kristja(Moossen et al., 2015), F) diatom-based August SST reconstruction from core DA12-11/2-GC01 (this study), G) diatom-based August SST reconstruction from core LO09-14(Berner et al., 2008), H) surface spring-summer temperature reconstruction (based on paired Mg/Caed18O measurements on planktic foraminifera Globigerina bulloides) fromcore RAPiD-12-1K (Thornalley et al., 2009), I) foraminifera-based summer SST reconstruction from core M23258 (Sarnthein et al., 2003), J) foraminifera-based subsurface temperature reconstruction from cores T88-2/JM01-1199 (Hald et al., 2007), K) dinocyst-based summer (JulyeSeptember) SST reconstruction from core M23323 (Van Nieuwenhove
et al., 2016), L) alkenone-based annual SST reconstruction from core MD95-2011 (Calvo et al., 2002), M) foraminifera-based summer SST reconstruction from core MD95-2011(Andersson et al., 2010). The shaded area highlights the reconstructed cold SST interval at c. 4-2 ka BP of this study.

temperature, show contrasting results: both greater variability at 42 ka BP (Thornalley et al., 2009; Figs. 6H and 7 (point 8)) which may
reﬂect the cold event, and warming after 4 ka BP (Fig. 7 (points 7
and 8); Farmer et al., 2008). Sub-surface temperature reconstructions based on planktic foraminifera also do not have aclear cold event; instead they show two large oscillations after 4 ka
BP (core ODP-984; Fig. 7 (point 5); Came et al., 2007), a strong
warming at 4-2 ka BP (core MD95-2015; Marchal et al., 2002; Fig. 7(point 9)) and relatively stable temperatures (NA8722; Fig. 7 (point10); Duplessy et al., 1992). Therefore, between 4 and 2 ka BP thisregion had a summer cooling of the surface water, which does notappear to have inﬂuenced the subsurface.
On the northwest Iceland shelf (cores MD99-2269 and MD992266) the reconstructions of summer SST based on diatoms andalkenones show a cooler interval at 4-2 ka BP (Fig. 6C, D and 6E and
7 (point 1 & 2); Andersen et al., 2004a; Justwan et al., 2008;
nsdo
 ttir et al., 2017). On the northeast Iceland shelf, however,Kristja
diatom and alkenone-derived reconstructions (cores MD99-2275

and JR51-GC35) do not show this cooling; the different records
instead display either a shorter but high magnitude cold event at
, 2007), a3.2e2.6 ka BP (Fig. 7 (point 3); Bendell and Rosell-Mele
gradual cooling trend (Fig. 7 (point 4); Jiang et al., 2015) or warming(Fig. 7 (point 4); Sicre et al., 2008).
On the East Greenland shelf the occurrence of a cold interval at
4-2 ka BP is somewhat uncertain. It is shown by surface (diatombased) and subsurface (foraminifera-based) temperature reconstructions from core MD99-2322 with a cooling at c. 3.5e3 ka
BP followed by a warming trend (Figs. 6B and 7 (point 11); Jennings
et al., 2011; Miettinen et al., 2015), as well as in core Fox05R near
Sermilik (Andresen et al., 2013), where inﬂow of warm Atlantic
subsurface water reduced after 3.6 ka BP. However, a diatom studyof cores HU93030-19A and BS88-06-5A showed low temperatures
after 2 ka BP without warming (Andersen et al., 2004a).
In the Greenland Sea, a foraminifera-based subsurface temperature reconstruction shows pronounced subsurface cooling at
c.5.7e2 ka BP followed by warming (core PS1878; Figs. 6A and 7

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L.C. Orme et al. / Quaternary Science Reviews 194 (2018) 128e142

Fig. 7. Map of the spatial distribution of SST reconstructions and the occurrence of the cold interval at 4-2 ka BP. The points are colour-coded; blue signiﬁes that the record shows acold interval at 4-2 ka BP, green denotes that there is some suggestion of a cold interval but with greater uncertainty and orange means that there is no evidence for a cold interval.
Proxies indicative of surface (e.g. diatoms, alkenones, dinoﬂagellate cysts) and subsurface (planktic foraminifera) temperature variability are highlighted by circles and squares
ttir et al., 2017), 2d MD99-2266 (Moossen et al., 2015), 3d JR51-GC35 (Bendell and Rosell-Mele, 2007), 4d
respectively. Core sites: 1e MD99-2269 (Justwan et al., 2008; Kristjansdo
MD99-2275 (Sicre et al., 2008; Jiang et al., 2015), 5e ODP-984 (Came et al., 2007), 6e LO09-14 (Berner et al., 2008; Solignac et al., 2008), 7e MD99-2251 (Farmer et al., 2008), 8d
RAPiD-12-1K (Thornalley et al., 2009), 9- MD95-2015 (Marchal et al., 2002), 10- NA8722 (Duplessy et al., 1992), 11e MD99-2322 (Jennings et al., 2011), 12d CR 19/5 (Andersen et al.,2004a), 13d MSM5/5-723-2 (Werner et al., 2016), 14d MSM5/5-712-2 (Werner et al., 2016), 15dMD99-2304 (Hald et al., 2007; Ebbesen et al., 2007), 16e JM03-373PC2 (Rasmussen
et al., 2007, 2014), 17e M23258 (Marchal et al., 2002, orange; Sarnthein et al., 2003, blue), 18dT88-2 (Hald et al., 2007 and references therein), 19dPSh-5159N (Risebrobakken et al., ski et al., 2014), 22e M23323 (Van Nieuwenhove et al., 2016), 23d MD95-2011 (Marchal et al.,2010), 20- T79-51/2 (Hald et al., 2007 and references therein), 21e PS1878 (Telesin
2002; Calvo et al., 2002; Dolven et al., 2002; Andersson et al., 2003, 2010; Risebrobakken et al., 2003; Berner et al., 2011), 24e Troll 8903/28-03 (Klitgaard-Kristensen et al., 2001;
Hald et al., 2007), 25- PS21842 (Koç and Jansen, 1994). (For interpretation of the references to colour in this ﬁgure legend, the reader is referred to the Web version of this article.) ski et al., 2014), however surface temperatures(point 21); Telesin(diatom-based SST from core PS21842) appear to have decreased at
7-2 ka BP before they stabilised (Fig. 7 (point 25); Koç and Jansen,
1994). Further records from both the East Greenland Shelf and
the Greenland-Iceland Seas would be required to identify withgreater conﬁdence the temperature patterns for these regions.
In the region to the west and south of Svalbard the majority oftemperature reconstructions are based on foraminifera and therefore reﬂect subsurface temperatures. Many of these show thesubsurface temperatures cooled between 5 and 3-2 ka BP, followed
by warming (cores MSM5/5-723-2, MSM5/5-712-2, MD99-2304,
JM03-373PC2; Figs. 6 I and Fig. 7 (points 13, 14, 15, 16, 17);
Sarnthein et al., 2003; Ebbesen et al., 2007; Rasmussen et al., 2007,
2014; Werner et al., 2013, 2016). A transect of cores between Norway and Svalbard (including core T88-2; Figs. 6J and 7 (point 18))shows that the 3 northernmost records (north of 70 ) had cooler
subsurface temperatures at 5-3 ka BP (Hald et al., 2007). However,an alkenone-based SST record (core M23258) instead indicates acooling after 3 ka BP (Fig. 7 (point 17), Marchal et al., 2002).
Therefore, in this region it is clear that the subsurface signal is inline with the DA12-11/2-GC01 core reconstruction, however the
SST patterns remain unclear.
Within the Norwegian Sea reconstructions do not show a cold
event at 4-2 ka BP. Offshore northern Norway, an alkenone SST
reconstruction (PSh-5159N) showed cooling after 3 ka BP that lasted
until present (Fig. 7 (point 19); Risebrobakken et al., 2010). However,
IRD here increased between c. 5-2 ka BP, which is interpreted as
representing cooler summers when sea ice could reach the site
(Risebrobakken et al., 2010). On the Vøring plateau there are some
contrasting patterns: alkenone and diatom-based SST estimates

L.C. Orme et al. / Quaternary Science Reviews 194 (2018) 128e142

from core MD95-2011 (Fig. 7 (point 23)) indicate a cooling trend at
4.0e2.5 ka BP followed by stable temperatures (Fig. 6L; Marchal
et al., 2002; Calvo et al., 2002; Birks and Koç, 2002; Berner et al.,2011), a radiolarian-based reconstruction shows no changes(Dolven et al., 2002) and a warming trend is shown by a dinocystbased SST reconstruction (core M23323; Figs. 6K and 7 (point 22);
Van Nieuwenhove et al., 2016). Subsurface (foraminifera-based)
temperature reconstructions also do not show a cold event, instead
showing only little variation (Core Troll 8903; Fig. 7 (point 24);
Klitgaard-Kristensen et al., 2001; Hald et al., 2007) or a warming
trend (MD95-2011, Fig. 6M; Risebrobakken et al., 2003; Andersson
et al., 2003, 2010). Therefore, it can be concluded that a pronounced cold period did not occur at 4-2 ka BP in the Norwegian Sea.
In summary, the period between 4 and 2 ka BP that featured
colder summer SSTs in the Iceland Basin is manifested in some, but
not all, surface and subsurface temperature reconstructions south
and north of Iceland, from the East Greenland Shelf, the Greenland
Sea and to the south and west of Svalbard (Fig. 7). The identiﬁcationof this cold event across a range of sites supports that this was not a
result of local factors, such as glacier ﬂuctuations, or methodological factors, such as the long-distance transport of diatoms. In someof the reconstructions (such as LO09-14 and MD99-2269 from the
Iceland Basin and shelf, PS1878 from the Greenland Sea and T88e1
and M23258 south of Svalbard; Fig. 6) the cool period began earlier
at c. 5 ka BP. Some of the difference may be explained by age uncertainties from dating and marine reservoir ages. However, between 5 and 4 ka BP the DA12-11 core and others show gradualcooling, supporting that the onset of this cool event began around 5
ka BP. It is notable that the cold interval is not captured in reconstructions that are situated along the Norwegian coast but isprimarily identiﬁed in reconstructions from the sites directly
inﬂuenced by the EGC, or in regions proximal to this current.
Principal component analysis on surface and near-surface temperature reconstructions from the subpolar north AtlanticFennoscandian region by Sejrup et al. (2016) shows cold surface
temperatures during the Late Holocene. The subsurface temperature reconstructions however suggest a cooling centred at c. 2 ka BP
in the west followed by warming, with opposite patterns towards
Norway (Sejrup et al., 2016). Although there is some difference inthe timing of the cooling event, the contrasting patterns between
the east and west, and the spatial distribution of the subsurface
temperature, reﬂects the ﬁndings discussed here.
5.3. Potential causes of the cold period at 4-2 ka BP
5.3.1. Arctic water outﬂow
The cold interval at 4-2 ka BP coincides with a period that had
prevalent negative NAO conditions (Olsen et al., 2012, Fig. 8C) andsouthward-shifted storm-tracks (Bakke et al., 2008; Orme et al.,2017, Fig. 8A and B), supporting the possibility that atmosphericforcing caused the observed SST patterns. The SST decrease and theincrease in the Sub-Arctic assemblage at 4-2 ka BP, may have beencaused by a greater southward transport of polar water by the EGC
in response to a period with persistent negative NAO. Such phenomena were observed at sub-decadal scale in the 20th century
when strongly negative NAO anomalies and high pressure overGreenland caused strengthened northerly winds over east
Greenland, a strengthened EGC and increased polar water transport
from the Arctic Ocean (e.g. the Great Salinity Anomaly of the 1960's;
Blindheim et al., 2000). Given the complex forcings and feedbacks
between atmospheric and ocean temperatures and ocean circulation, it is conceivable that ocean surface water cooling may havebeen the cause (rather than the result) of the generally negative
NAO at 4-2 ka BP. However, we consider this unlikely as the SST
cooling appears to be regional (Fig. 7) and goes against the general137

trend of gradual cooling and increasing sea ice in the Arctic (Kimet al., 2004; Vare et al., 2009) that most likely would have had agreater inﬂuence on the atmosphere. Thus, the surface watercooling was more likely a result of a generally negative NAO scenario. Modelling studies have shown similar events at multidecadal
timescales; these feature an export of fresh, cold water and sea icefrom the Arctic into the Greenland Sea, which was then transported
south through the Denmark Strait into the North Atlantic (including
the Icelandic Basin) as a result of strengthened northerly winds(Delworth et al., 1997; Renssen et al., 2005). Interestingly, theRenssen et al. (2005) simulation indicates that these episodes
occurred particularly in the period 3.5e1.5 ka BP. Therefore, we
speculate that the period 4-2 ka BP had enhanced southwardtransport of the EGC, which allowed subarctic water into the Iceland Basin, either directly from the Denmark Strait and across the
Reykjanes Ridge with a south-eastward shift in the subarctic front,
or via recirculation within the SPG.
Various studies have found evidence that the EGC was stronger
at this time. On the East Greenland shelf, a stronger EGC at 4.5e1.4
ka BP has been inferred from variations in sea ice (Perner et al.,2015; Kolling et al., 2017) and after 3.6 ka BP foraminiferalchanges indicate an increased inﬂuence of polar water, although
this continued after 2 ka BP (Andresen et al., 2012). To the north,Justwan et al. (2008) interpreted cooler SSTs and an increase in an
Arctic diatom assemblage as reﬂecting a strengthened East Iceland
Current at 4.7e2.2 ka BP. In addition, to the west of Svalbard, cold
temperatures and enhanced sea ice at 5.2 to 3 ka BP have beeninterpreted as a spread of Arctic water through the Fram Strait
along a strengthened EGC (Werner et al., 2013). Finally, a study ofthe temporal variability of reservoir ages during the Late Holocene
from the north of Iceland shows higher reservoir ages at 4/3 - 2 ka
BP in three cores, which is interpreted as resulting from greater
input of ‘old’ carbon from the Arctic Ocean via a strengthened EGC(Eiriksson et al., 2004). The strengthened EGC may have allowed
more Arctic water to enter the Iceland Basin at around 4 ka BP, asevidenced by dinoﬂagellate cyst variability (Solignac et al., 2004;
Van Nieuwenhove et al., 2018) and ﬂuctuations in foraminifera anddiatoms on the Reykjanes Ridge at 3-2 ka BP (Rasmussen et al.,2002). Together, the evidence supports that the period 5/4 ka BP
until 2 ka BP may have had greater atmospherically-driven Arctic
water outﬂow along the EGC.
5.3.2. Greenland ice melt
A second hypothesis to explain the cooling at 4-2 ka BP is that ofincreased meltwater runoff from Greenland. Modern measurements show a strong negative correlation between the NAO andGreenland temperature (Serreze et al., 1997; Chylek et al., 2004);
therefore, the negative NAO at 4-2 ka BP (Olsen et al., 2012; Faust
et al., 2016) may have caused anomalous regional warming andincreased the amount of meltwater entering the North Atlantic andGreenland Sea. This is supported by evidence of warmer temperatures over the Greenland ice sheet between 4 and 2 ka BP (Alley,
2004; Briner et al., 2016) but contradicted by other reconstructions that show cooling trends or stable temperatures overGreenland during this interval (e.g. Johnsen et al., 2001; Kobashi
et al., 2017). During the mid-late Holocene, the period before
c.2e3 ka BP had a steep reduction in elevation at the GRIP drill site,
before the elevation slightly increased after 2 ka BP (Vinther et al.,2009), and the Greenland ice sheet reached its minimum extent at
5-3 ka BP (Briner et al., 2016). This latter study in particular supports that ice loss from Greenland may be a possible factor for
explaining the cooling of the subpolar North Atlantic. An input of
meltwater is supported by evidence of lower salinities south of
Iceland, on the Greenland Rise and the Rockall Plateau at c. 5-2 ka
BP (Duplessy et al., 1992; Solignac et al., 2004; Came et al., 2007;

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Fig. 8. Comparison between reconstructions reﬂecting atmospheric and ocean circulation changes over the last 6 ka BP. A) storm track reconstruction based on contrasting
reconstructed precipitation between southern and northern Norway (Bakke et al., 2008), B) storm track reconstruction based on contrasting patterns of reconstructed storminessfrom Scotland and Spain (Orme et al., 2017), C) NAO reconstruction based on hypolimnetic anoxia from a lake in Greenland (Olsen et al., 2012), D) DA12-11/2-GC01 SST reconstruction (this study), E) Atlantic Water Inﬂow reconstruction in the Nordic Seas (Giraudeau et al., 2010), F) upper water column stratiﬁcation reconstruction for the Iceland Basin(stratiﬁcation increases upwards) based on the d18O difference between Globigerina bulloides and Globorotalia inﬂata (Thornalley et al., 2009) and G) near-surface salinity reconstruction from the Iceland Basin derived from paired Mg/Caed18O measurements on planktic foraminifera Globigerina bulloides (Thornalley et al., 2009). The shaded area highlights
the reconstructed cold SST interval at c. 4-2 ka BP of this study.

Thornalley et al., 2009; Van Nieuwenhove et al., 2018). In the Iceland Basin (RAPiD-12-1K) the negative NAO events in particularcoincide with low salinity anomalies in the surface waters(Thornalley et al., 2009, Fig. 8G). However, this freshening may haveequally been caused by a strengthened EGC and greater export offreshwater from the Arctic.
A model forced by the Early Holocene melting of the Greenland
Ice Sheet showed the clearest SST response around southeastern
Greenland, extending into the Iceland Basin and through the
Denmark Strait (Blaschek and Renssen, 2013). This study ﬁnds thatclose to Greenland the freshening of surface water caused

stratiﬁcation and cooling because of reduced vertical mixing, while
cooling across the North Atlantic was the result of reduced AMOC
strength following meltwater input to the Labrador Sea (Blaschek
and Renssen, 2013). To know if similar changes occurred at 4-2 ka
BP, we require estimates of the volume of meltwater produced
during the period of reduced Greenland ice sheet extent at 5-2 ka
BP (Briner et al., 2016) and modelling of the effects of smaller
freshwater inputs. However, convection in this region is very sensitive to even small quantities of additional freshwater because ofstrong amplifying feedbacks related to the circulation of the subpolar gyre (Born et al., 2010; Born and Stocker, 2014).

L.C. Orme et al. / Quaternary Science Reviews 194 (2018) 128e142

5.3.3. Northward heat transport and the effects of SPG dynamics
The dynamics of the SPG and associated modulation of thenorthward heat transport is another potential cause of theobserved cooling at 4-2 ka BP. The increase of the sub-Arctic
assemblage at 4-2 ka BP may support this, as in addition to being
found to the north of Iceland it is also an assemblage characteristicof the SPG. The inﬂuence of the NAO on the SPG is uncertain. It has
been suggested that during the 1990's the positive to negative shift
in the NAO caused the SPG to change from an E-W to a N-S
conﬁguration, which resulted in warmer waters being transported
to the Iceland Basin (Hatún et al., 2005). However, subsequent
analysis would suggest that such longitudinal shifts of the eastern
SPG do not control the properties of the eastern subpolar North
Atlantic, and that the NAO does not cause changes in the conﬁguration of the SPG (Foukal and Lozier, 2017). Meanwhile, a model
suggests that a weaker SPG during negative NAO periods (with ablocking high pressure centred over the British Isles and eastern
North Atlantic) instead causes cooling in the Iceland Basin due toreduced advection (Moffa-Sanchez et al., 2014).
Reconstructions of SPG dynamics and northward heat transportshow steady or oscillating conditions rather than a clear change at
4-2 ka BP (Fig. 8F; Thornalley et al., 2009; Repschl€
ager et al., 2017;
Tegzes et al., 2017). Furthermore, in core DA12-11/2-GC01 thedinocyst Operculodinium centrocarpum, a species often foundwithin waters from the NAC, shows high values since 6.8 ka BP andparticularly at c. 4.4 ka BP, followed by declining values to present(Van Nieuwenhove et al., 2018), likely reﬂecting a continuous
reduction in the northward heat transport. This evidence suggeststhat at this time the SPG and NAC were not inﬂuenced by thenegative NAO and did not cause the SST cooling. The lack of evidence for a cold interval along the Norwegian coast (Fig. 7) may also
support that NAC heat transport was not reduced, although thisregion may not be sensitive to NAO changes if it was inﬂuenced by
NAC water during both positive and negative NAO phases(Blindheim et al., 2000). On the other hand, a record of Atlantic
Water inﬂow to the Nordic Seas does show a strong similarity with
the NAO and SST reconstructions, with reduced strength during theperiod 4-2 ka BP (Giraudeau et al., 2010, Fig. 8 E), which might
indicate that heat transport to the Iceland Basin was also reduced at
this time if the NAC weakened. In any case, the ambiguity in the
different proxy reconstructions outlined above precludes
concluding ﬁrmly that the millennial SST variations were caused byatmospherically-driven changes in SPG strength and northward
heat transport.
In summary, we consider that the SST cooling at 4-2 ka BP may
have been caused by either enhanced Arctic outﬂow via the EGC(section 5.3.1) or enhanced ice melt from East Greenland (section
5.3.2), both of which may be associated with negative NAO circulation. The currently available evidence of northward heat transport and SPG dynamics (section 5.3.3.) does not strongly support
that reduced heat transport caused the observed cooling.
5.4. Wider ocean changes
Lower salinity and colder water entering the subpolar northAtlantic (resulting from either Arctic outﬂow or melting of theGreenland ice sheet) may have inﬂuenced the wider ocean circulation, including the rates of deep water formation (i.e. Denmark
Strait Overﬂow Water (DSOW) and North Atlantic Deep Water) andas such the vigour of the AMOC. Observations indicate that during
the strongly negative NAO of the 1960's there was increased deepwater formation in the Greenland Sea due to stronger northerly
winds, strong cooling and increased sea ice formation (which led tobrine production) (Dickson et al., 1996). Therefore enhanced deepwater formation at 7-3 ka BP in the Greenland Sea (shown by high139

d13C values; Telesin ski et al., 2014) may have resulted from thisdirect atmospheric forcing. Furthermore, since DSOW is formed bywater transported southwards by the EGC (Mauritzen, 1996), thestrengthened outﬂow of DSOW at 6.5e3 ka BP southwest ofGreenland (MD99-2227) (Fagel et al., 2004) may have been the
result of the strengthened EGC. Increased freshwater input may
also have caused a strengthening of the Labrador Current at 5e3.1
ka BP, as it is known that varying input can modulate its vigour(Rashid et al., 2017). While Rashid et al. (2017) did not identify afreshwater input that could account for the high Labrador Current
strength, we suggest that it may have been supplied by a NAOdriven strengthening of the EGC or melting from Greenland. Thisevidence supports that ocean circulation beyond the northwest
Atlantic could have been inﬂuenced by the negative NAO circulation at 4-2 ka BP.
6. Conclusions
A diatom-based SST reconstruction from core DA12-11/2-GC01,
retrieved from the Iceland Basin, shows a pronounced SST coolingbetween 4 and 2 ka BP, with warmer temperatures prior to 4 ka BP
and after 2 ka BP. Comparison of the new reconstruction withothers from the subpolar North Atlantic indicates that this cold
period was also observed in some reconstructions to the west of thestudy region, in particular from locations under the inﬂuence of theEast Greenland Current.
We propose three hypotheses that might explain the observed
cooling and its spatial pattern. The ﬁrst, and perhaps most likely
cause of the cooling, is that a persistently negative NAO at this timecaused strengthening of northerly winds east of Greenland, whichstrengthened the EGC, bringing cool Arctic water as far south as the
Icelandic Basin. Alternatively, a greater input of meltwater fromGreenland caused by negative NAO-induced atmospheric warming
may have caused the SST cooling and freshening during this period.
A third hypothesis stipulates that the cooling was caused by
reduced northward heat transport via the NAC or SPG but this is notsupported by much of the available evidence. Although presented
here separately, we consider it likely that the three mechanisms
may have been operating concurrently.
It is predicted that negative NAO-like circulation may characterise future climate (e.g. Seierstad and Bader, 2009; Vihma, 2014)
and our results indicate that this may result in cooling in thewestern region of the subpolar North Atlantic. At 4-2 ka BP thisappears to have caused strengthened Greenland Sea deepwater ski et al., 2014), andformation and DSOW (Fagel et al., 2004; Telesina similar response may be expected with future climate change.
This may in turn inﬂuence the AMOC strength, but further researchwould be required to conﬁrm this for the period 4-2 ka BP and to
apply this feedback to future climate scenarios.
Declarations of interest
None.
Acknowledgements
We would like to thank the captain, crew and scientiﬁc party of
the DA12-11/2 expedition onboard R/V Dana for their work andefforts in recovering the core. The research cruise was funded by
the Danish Centre for Marine Sciences (2016-03/DCH) (Dansk
Center for Havforskning) as part of a PhD course “Benthic Researchat Sea” taking place in September 24e30, 2012. LO, AM, DD, KH, AB
and RM acknowledge funding from the Research Council of Norway(grant no. 248776/E10) and Ministry of Earth Science, Earth System
Science Organization (MoES/Indo-Nor/PS-2/2015). Additional

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funding was obtained from the Independent Research Fund
Denmark/Natural Science (G-Ice project no 7014-00113B; grant toMSS) and the Aarhus University Research Foundation (AUFF, grant
to CP). Many thanks to the helpful comments from two anonymous
reviewers. Additional thanks to Svetlana Divina for preparation of
the diatom slides.
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