1

Dansgaard-Oeschger cycles: interactions between ocean and 1 sea ice intrinsic to the Nordic Seas 2

Trond M. Dokken1,3, Kerim H. Nisancioglu2,1,3, Camille Li4,3, David S. Battisti5,3, 3

Catherine Kissel6 4

1. UNI Research AS, Allegaten 55, 5007 Bergen, Norway; 2. Department of Earth Science, University of 5 Bergen, 5007 Bergen, Norway; 3. Bjerknes Centre for Climate Research, 5007 Bergen, Norway; 4. 6 Geophysical Institute, University of Bergen, 5007 Bergen, Norway; 5. Department of Atmospheric 7 Sciences, University of Washington, Seattle WA 98195, USA; 6. Laboratoire des Sciences du Climat et de 8 l’Envionnement, CEA/CNRS/UVSQ, 91198 Gif-sur-Yvette Cedex, France. 9

ABSTRACT 10

Dansgaard-Oeschger (D-O) cycles are the most dramatic, frequent, and wide reaching 11

abrupt climate changes in the geologic record (Voelker, 202). On Greenland, D-O 12

cycles are characterized by an abrupt warming of up to 15°C from a cold stadial to a 13

warm interstadial phase (Dansgaard et al., 1993; Severinghaus & Brook, 1999; NGRIP, 14

2003), followed by gradual cooling before a rapid return to stadial conditions. The 15

mechanisms responsible for these millennial cycles are not fully understood, but are 16

widely thought to involve abrupt changes in Atlantic Meridional Overturning 17

Circulation (AMOC) due to freshwater perturbations (Broecker et al., 1985; Broecker et 18

al., 1990; Birchfield & Broecker, 1990). Here we present a new, high resolution multi-19

proxy marine sediment core monitoring changes in the warm Atlantic inflow to the 20

Nordic Seas as well as in local sea ice cover and influx of Ice Rafted Debris (IRD). In 21

contrast to previous studies, the freshwater input is found to be coincident with warm 22

interstadials on Greenland and has a Fennoscandian rather than Laurentide source. 23

Furthermore, the data suggest a different thermohaline structure for the Nordic Seas 24

during cold stadials, in which there is a continuous circulation of relatively warm 25

Atlantic water beneath a fresh surface layer topped by sea ice. This implies a delicate 26

2

balance between the warm sub-surface Atlantic water and fresh surface layer, with the 27

possibility of abrupt changes in sea ice cover, and suggests a novel mechanism for the 28

abrupt D-O events observed in Greenland ice cores. 29

1. INTRODUCTION 30

There is a wealth of proxy data showing that the climate system underwent large, abrupt 31

changes throughout the last ice age. Among these are about one score large, abrupt 32

climate changes known as Dansgaard-Oeschger events that occurred every 1-2 kyr. 33

These events were first identified in ice cores taken from the summit of Greenland 34

(Dansgaard et al. 1993; Bond et al., 1993), and are characterized by an abrupt warming 35

of up to 15C in annual average temperature (Severinghaus & Brook, 1999; Lang et al. 36

1999; Landais et al. 2004; Huber et al. 2006). Subsequent observational studies 37

demonstrated the D-O events represent a near-hemispheric scale climate shift (see e.g. 38

Rahmstorf, 2002 for an overview). Warming in Greenland is coincident with warmer, 39

wetter conditions in Europe (Genty et al., 2003), an enhanced summer monsoon in the 40

northwest Indian Ocean (Schulz et al., 1998; Pausata et al., 2011), a northward shift of 41

precipitation belts in the Cariaco Basin (Peterson et al., 2000), aridity in the 42

southwestern United States (Wagner et al., 2010), and changes in ocean ventilation off 43

the shore of Santa Barbara, California (Hendy et al., 2002). Each abrupt warming 44

shares a qualitatively similar temperature evolution: about 1000 years of relatively 45

stable cold conditions, terminated by an abrupt (less than 10 years) jump to much 46

warmer conditions that persist for 200-400 years, followed by a more gradual transition 47

(~50 to 200 years) back to the cold conditions that precede the warming event. This 48

sequence of climate changes is often referred to as a D-O cycle. 49

3

The leading hypothesis for D-O cycles attributes them to an oscillation of the Atlantic 50

Meridional Overturning Circulation (AMOC) (Broecker et al., 1985), with "on" and 51

"off" (or "weak") modes of the AMOC (Stommel, 1961; Manabe & Stouffer, 1985) 52

producing warm interstadials and cold stadials, respectively, through changes in the 53

meridional ocean heat transport. Indeed, interstadial and stadial conditions on 54

Greenland summit are associated with markedly different water mass properties 55

(temperature and salinity) in the subtropical Atlantic (Sachs & Lehman, 1999), North 56

Atlantic (Curry & Oppo, 1997) and Nordic Seas (Dokken & Jansen, 1990). Transitions 57

between the two phases are thought to arise from changes in the freshwater budget of 58

the North Atlantic (Broecker et al., 1990). In particular, Birchfield and Broecker (1990) 59

proposed that the continental ice sheets covering North America and Eurasia were the 60

most likely sources for the freshwater. 61

A number of studies with climate models of intermediate complexity (EMICs) have 62

produced two quasi-stable phases in the AMOC by imposing ad hoc cyclic freshwater 63

forcing (Marotzke & Willebrand, 1991; Ganopolski & Rahmstorf, 2001; Knutti et al., 64

2004). Although these models simulate a temperature evolution on Greenland Summit 65

that is qualitatively similar to observations, the changes in the simulated local and far-66

field atmospheric response are small compared to observations. Similar experiments 67

with state-of-the-art climate models produce local and far-field responses that are in 68

better agreement with the proxy data (e.g. Vellinga & Wood, 2002). However, these 69

models require unrealistically large freshwater perturbations, and on Greenland, the 70

simulated stadial to interstadial transitions are not as abrupt as observed. In both cases, 71

the prescribed external forcing is ad hoc, and yet is entirely responsible for the existence 72

and the longevity of the two phases (i.e., the cold phase exists as long as the freshwater 73

4

forcing is applied). 74

Despite the unresolved questions surrounding D-O cycles, there has been progress. The 75

direct, causal agent for the abrupt climate changes associated with D-O warming events 76

is thought to be abrupt reductions in North Atlantic sea ice extent (Broecker 2000; 77

Gildor and Tziperman 2003; Masson-Delmotte et al. 2005; Jouzel et al. 2005). Studies 78

using atmospheric general circulation models (AGCMs) show that the local 79

precipitation and temperature shifts (inferred from oxygen isotopes and accumulation 80

changes) on Greenland summit associated with a typical D-O event are consistent with 81

the response of the atmosphere to a reduction in winter sea ice extent, in particular in 82

the Nordic Seas region (e.g. Figure 1; see also Renssen and Isarin 2001; Li et al 2005; 83

Li et al. 2010). Away from the North Atlantic, the northward shift of tropical 84

precipitation belts associated with D-O warming events mirrors the southward shift 85

simulated by coupled climate models in response to freshwater hosing and expansion of 86

sea ice (Vellinga and Wood 2002; Otto-Bliesner & Brady, 2010; Lewis et al., 2010), 87

and by AGCMs coupled to slab ocean models in response to a prescribed expansion of 88

North Atlantic sea ice (Chiang et al., 2003). Thus supporting the existence of a robust 89

link between North Atlantic sea ice cover and far-field climate. 90

More recently, a number of studies have shown that cold stadial conditions are 91

associated with subsurface and intermediate depth warming in the North Atlantic 92

(Rasmussen & Thomsen 2004; Marcott et al. 2011). This subsurface warming may have 93

bearing on many aspects of D-O cycles, from the abruptness of the warming (stadial-to-94

interstadial) transitions (Mignot et al., 2007) to the possibility of ice shelf collapses 95

during D-O cycles (Alvarez Solas et al., 2011; Marcott et al., 2011; Petersen et al., 96

5

2013). 97

In this study, we focus on the subsurface warming rather than ice-rafting to build 98

physical arguments for the mechanisms behind D-O cycles. Changes in sea ice and 99

subsurface ocean temperatures during D-O transitions are linked to each other by the 100

ocean circulation, and to changes in Greenland by the atmospheric circulation. To 101

elucidate these links, requires marine records that (i) are situated in key regions of the 102

North Atlantic, (ii) are representative of conditions through the depth of the water 103

column, and (iii) can be compared to records of atmospheric conditions (as measured 104

from the Greenland summit) with a high degree of certainty in terms of timing. In 105

addition, because of the abruptness of the transitions (on the order of decades or less), 106

the marine records must be of exceptionally high resolution. 107

Here we present new data from a Nordic Seas core that meet these criteria. The core site 108

is strategically located in the Atlantic inflow to the Nordic Seas, where extremely high 109

sedimentation rates allow for decadal scale resolution of both surface and deep-water 110

proxies. In addition to an age model based on calibrated C-14 ages and magnetic 111

properties, ash layers provide an independent chronological framework to synchronize 112

signals in the marine core with signals in an ice core from the summit of Greenland. 113

The new marine proxy data reveal systematic changes in the hydrography of the Nordic 114

Seas as Greenland swings from stadial (cold) to interstadial (warm) conditions and back 115

through several D-O cycles, with details of the transitions better resolved than in any 116

previously studied core. The interstadial phase in the Nordic Seas is shown to resemble 117

current conditions, while the stadial phase has greatly enhanced sea ice coverage and a 118

hydrographic structure similar to that in the Arctic Ocean today. The variations recorded 119

6

by the marine proxies near the transitions are used to determine the possible 120

mechanisms that cause the North Atlantic climate system to jump abruptly from stadials 121

to interstadials and to transition slowly back from interstadials to stadials. In a departure 122

from most existing hypotheses for the D-O cycles, the conceptual model presented here 123

does not require freshwater fluxes from ice sheets to enter the ocean at specific times 124

during the D-O cycle. Rather, it focuses on the role of ocean-sea ice interactions within 125

the Nordic Seas in triggering stadial-interstadial transitions. 126

The paper proceeds as follows. Section 2 contains a description of the marine sediment 127

core, analysis methods, the construction of the age model for the core, and the 128

construction of a common chronology with an ice core from the summit of Greenland. 129

Section 3 describes the cold stadial and warm interstadial climate of the Nordic Seas as 130

seen in the marine sediment data. Section 4 analyzes the relevance of the new proxy 131

data for the D-O cycles as observed on Greenland, and in particular the mechanisms 132

behind the abrupt transitions. Section 5 presents a summary of the results and a 133

discussion of the implications. 134

2. MATERIAL AND METHODS 135

Core MD992284 was collected in the Nordic Seas during the MD114/IMAGES V 136

cruise aboard R.V. Marion Dufresne (IPEV) at a water depth of 1500 m on the 137

northeastern flank of the Faeroe-Shetland channel. The core is strategically located in 138

the Atlantic inflow to the Nordic Seas and has an exceptionally high sedimentation rate, 139

allowing for decadal scale resolution of both surface and deep water proxies. The 140

unique, high-resolution marine core with a well constrained age model (Figure 2) shows 141

evidence for systematic differences in the Nordic Seas during Greenland interstadials 142

7

versus Greenland stadials. These new data are used to characterize the important 143

hydrographic features of the Nordic Seas in each of these two phases of the D-O cycle 144

as well as the transition between the phases. 145

Figure 2a shows a segment of the NorthGRIP ice core (NGRIP) from Greenland 146

(NGRIP, 2004) during late Marine Isotope Stage 3 (MIS3) containing several typical D-147

O cycles. To compare the marine record to the ice core record, we have constructed an 148

age model by matching rapid transitions in MD992284 to NGRIP using the Greenland 149

Ice Core Chronology 2005 (GICC05; Svensson et al., 2006) and ash layers found in 150

both cores. The common chronology allows the Nordic Seas data to be linked directly to 151

Greenland ice core signals. 152

Age model for marine sediment core 153

Age models based on the ice core and calibrated 14C dates can be used to estimate “true 154

ages” for the marine core, but will be subject to errors due to missing or misinterpreted 155

ice layers, uncertainty in reservoir ages and uncertainty in sedimentation rates. Our 156

approach here is thus not to rely on “true ages”, but rather on a stratigraphic tuning of 157

the marine record to the ice core record. We use the high frequency variations in 158

measured Anhysteretic Remanent Magnetization (ARM) during Marine Isotope Stage 3 159

(MIS3) in the marine core (Figure 2b). The fast oscillations in magnetic properties 160

during MIS3 in the North Atlantic/Nordic Seas have been shown to be in phase with 161

changes in d18O of Greenland ice cores (Kissel et al., 1999). We use the ARM record 162

for tuning only, with no further subsequent paleo-proxy interpretation of this parameter. 163

The tuning has been performed using the tuning points as indicated in Figure 2 and the 164

software package AnalySeries 2.0 (Paillard et al., 1996). 165

8

A final step in the synchronization between the ice core and the marine core is done 166

with tephra layers. Two ash zones identified in the marine record have also been 167

identified in the NGRIP ice core (Svensson et al., 2008), the Fugloyarbanki Tephra 168

(FMAZII) (Davies et al., 2008) and Fareo Marine Ash Zone III (FMAZIII) 169

(Wastegaard et al., 2006). These ashes allow for a direct, independent comparison 170

between the marine and ice core records. Our chronology places the ash layer FMAZIII 171

less than 100 years after the onset of IS8 in NGRIP, and thus provides a verification of 172

the initial tuning of the ARM record in MD992284 to the d18O record in NGRIP. These 173

two ash layers are also used as tuning points in the final age model. The final age model 174

gives the best possible synchronization of the marine record to the NGRIP ice core 175

record (NGRIP, 2004) making it possible to determine the relative timing of observed 176

changes in the marine proxies compared to temperature changes on Greenland. 177

The final tuned age model presented for MIS3 (Figure 3) only uses the ages given by 178

the tuning to NGRIP and the ash layers. As an independent check, we note that the new 179

age model for the marine sediment core is consistent with calibrated 14C dates (Table 1). 180

These are calculated using the calibration software CALIB 6.0 (Stuiver & Reimer, 181

1993) and an ocean reservoir age of 400 years, applying the “Marine09” calibration 182

curve (Hughen et al., 2004) for ages younger than 25 kyr BP (14C age, kilo years before 183

present) and the “Fairbanks0107” calibration curve for AMS 14C samples of ages older 184

than 25 kyr BP (Fairbanks et al., 2005). Offsets of at least a few thousand years from 185

the “true age” are expected due to the relatively large error in 14C-measurements in the 186

interval 30 - 40 kyr BP, as well as errors in transformation from 14C-ages to calibrated 187

ages. Additional error is introduced by the assumption that sedimentation rates are 188

constant between the dated levels. 189

9

Near surface ocean temperature estimates 190

For the planktonic foraminifera, raw census data of the planktonic assemblages were 191

analyzed and the transfer function technique (Maximum Likelihood - ML) was used to 192

calculate near surface temperature. The ML method operates with a statistical error 193

close to ±1ºC, and is found to give the least autocorrelation compared to other statistical 194

methods (Telford & Birks, 2005). 195

Transforming the δ18O into a measure of sea water isotopic composition 196

The ratio of oxygen-18 to oxygen-16 as they are incorporated into foraminiferal calcite 197

(δ18Ocalcite) is dependent on both seawater δ18O (expressed as per mil deviations from 198

Vienna Standard Mean Ocean Water - SMOW) and calcification temperature. 199

Foraminiferal based SST estimates provide an independent estimate of the calcification 200

temperature, and therefore the temperature effects on calcite δ18O. The temperature 201

to δ18O calibration used is based on Kim and O´Neil (1997): 202

T = 16.1 - 4.64 (δ18Ocalcite - δ18OSMOW) + 0.09 (δ18Ocalcite - δ18OSMOW)2 203

Knowing the calcification temperature (T) from the foram transfer function and the 204

measured δ18Ocalcite we can extract the isotopic composition of the ambient sea water 205

(δ18OSMOW). The relationship between salinity and δ18OSMOW in the ocean depends on 206

the spatial variation of δ18O associated with the fresh water flux to the ocean, the 207

amount of sea water trapped in ice caps, as well as ocean circulation (Bigg & Rohling, 208

2000). 209

Before translating glacial δ18O data into a measure of δ18OSMOW, the effect of global ice 210

volume must be accounted for. Fractionation processes of oxygen isotopes when water 211

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freezes into ice produce an ocean-wide δ18O shift of about 0.1 ‰ per 10 m of sea level 212

stored in continental ice sheets (Fairbanks, 1989). A review of recent advances in the 213

estimates of sea level history during MIS3 is given by Siddall et al. (2008). The 214

estimated sea level changes during MIS3 vary in amplitude and timing between the 215

different reconstructions. In our study, we use the sea level reconstruction of 216

Waelbroeck et al. (2002), which fits well with available coral based estimates for MIS3, 217

to remove the effects of changing ice volume on the δ18O of ocean water. 218

3. Interstadial and Stadial phases of a D-O cycle 219

Positioned in the path of inflowing warm Atlantic water to the Nordic Seas and close to 220

the routing of freshwater from the Fennoscandian ice sheet, our marine sediment core is 221

ideally suited to better understanding the changes to the hydrography and sea ice cover 222

of the Nordic Seas during MIS3. Here we will describe the warm and cold phases of D-223

O cycles in the Nordic Seas as suggested by the new sediment core data, before 224

examining the transitions in section 4. 225

The Nordic Seas during Greenland stadials 226

Foraminiferal assemblages used to reconstruct near surface ocean temperature show that 227

water masses at the core site in the Nordic Seas are ~2°C warmer during cold Greenland 228

stadials than during warm interstadials (Figure 3b). Further, there is a warm overshoot 229

at the onset of each interstadial. Despite difficulties in acquiring reliable low-230

temperature reconstructions in polar and Arctic water masses (e.g., Pflaumann et al. 231

2003), the estimated higher temperatures during stadials compared to interstadials is a 232

robust result. This is based on higher percentages of “Atlantic species” associated with 233

relatively warm conditions (such as Globigerina bulloides, Neogloboquadrina 234

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pachyderma (dextral) and Turborotalita quinqueloba) during stadials. Interstadials are 235

instead dominated by Neogloboquadrina pachyderma (sinistral). 236

Planktonic foraminifera are known to change their depth habitat in Arctic environments 237

to avoid the low salinity surface, which is separated from the subsurface warm Atlantic 238

layer by a halocline (strong vertical salinity gradient). This results in different depth 239

distributions of foraminifera under different hydrographic conditions (Carstens et al. 240

1997). Therefore a reasonable interpretation of these data is that, during cold stadials in 241

the Nordic Seas, the planktonic assemblages are located below the fresh surface layer 242

and thus record the relatively warm temperatures just below the halocline (Figure 4a). 243

At these depths, the foraminifera are in contact with relatively warm Atlantic water that 244

is isolated from the atmosphere by the halocline and sea ice cover, both of which are 245

features of the stadial phase of the D-O cycle. Note, that at the end of each stadial 246

phase, there is a warm overshoot that provides a clue to understanding the stadial-to-247

interstadial transition (discussed in section 4). 248

Combining stable isotope (δ18Ocalcite) measurements on the planktonic foraminifera N. 249

pachyderma sin. with the ocean temperature reconstruction produces an estimate of the 250

isotopic composition of sea water (δ18OSMOW). δ18OSMOW is partly related to salinity, but 251

it also reflects spatial shifts in water masses. The isotopic composition of polar water 252

may be different compared to water of subtropical origin, but because the planktonic 253

foraminifera prefer to stay within the warm Atlantic layer, we do not expect that the 254

mean δ18OSMOW (salinity) at the core site will be substantially different during stadials 255

compared to interstadials. Indeed, we see in Figure 3c that there is no systematic 256

difference in the mean δ18OSMOW during stadial and interstadial periods. The common 257

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spikes at the transitions between stadial and interstadial phases and the consistent trends 258

within the stadial and interstadial periods provide insight into the transition mechanisms 259

and are discussed in section 4. 260

The benthic δ18O record measured on Cassidulina teretis (Figure 3d) is remarkably 261

similar to the NGRIP record, exhibiting the characteristic shape and abrupt transitions 262

of D-O cycles. One explanation for the link between the benthic δ18O and NGRIP is 263

provided by Dokken and Jansen (1999), who presented evidence that there are two 264

different forms of deep water production in the Nordic Seas during the last glacial 265

period, depending on the sea ice coverage. As in the modern climate, there is deep open 266

ocean convection in the Nordic Seas during the last glacial period when the surface is 267

free of sea ice. When Greenland is cold (the stadial phase of the D-O cycle), however, 268

they argued that the benthic δ18O becomes lighter in the deep Nordic Seas because there 269

is sea ice formation along the Norwegian continental shelf, which creates dense and 270

istopically light brine water that is subsequently transported down the continental slope 271

and injected into the deeper water. In this scenario, the close link between the benthic 272

δ18O record and NGRIP is due to the deep water being affected by continuous sea ice 273

formation throughout the stadial phase and by open ocean convection during the 274

interstadial phase of the D-O cycle. This interpretation of the benthic record is 275

commensurate with all of the other proxy data in Figure 3, which indicates that during 276

the stadial phase of the D-O cycle, the Nordic Seas was covered with sea ice, the 277

surface waters were fresh (Figure 3c) and at the freezing point. Further evidence for the 278

existence of extensive sea ice stems from Figure 3a and modeling work that links the 279

Greenland δ18O record to sea ice extent in the Nordic seas (see Section 1). 280

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An alternative interpretation of benthic δ 18O at nearby core sites attributes the changes 281

to bottom water temperature rather than brine (Rasmussen & Thomsen, 2004). The 282

temperature change that would be required to explain the 1-1.5 per mil shifts seen in 283

benthic δ18O at our site implies a 4-6°C warming at a depth of 1500 m during stadials, 284

and almost twice as much during Heinrich events. We believe that enhanced brine 285

rejection at the surface and moderately increased temperatures at depth act together to 286

create the light benthic δ18O values during stadials, and that both processes are 287

necessary for maintaining stadial conditions. 288

The record of Ice Rafted Debris (IRD) indicates that there is less frequent ice rafting 289

during cold stadial periods on Greenland (Figure 3e). In this core, the IRD is of 290

Scandinavian origin, and the record implies that there is less calving from the 291

Fennoscandian ice sheet during the cold stadial, and therefore less meltwater derived 292

from ice bergs. Note that there are two Heinrich events during the time interval shown 293

in Figure 3 (denoted H3 and H4 in Figure 3), both of which occurred during the latter 294

part of a cold stadial phase of a D-O cycle. In this study, we are not concerned with the 295

infrequent, large Heinrich ice rafting events and do not consider Heinrich events to be 296

fundamental to the physics of a D-O cycle. Instead, the reader is referred to the study of 297

Alvarez-Solas et al. (2011) for a discussion of the possible relationship between these 298

two phenomena. 299

In general, the planktonic δ13C exhibits more negative values during stadials (Figure 300

3f), suggesting that the water at the preferred depth for the planktonic foraminifera is 301

isolated from the atmosphere. Thus, benthic and planktonic data collectively suggest 302

that during the stadial phase of the D-O cycle, the eastern Nordic Seas is characterized 303

14

by extensive sea ice cover, a surface fresh layer separated by a halocline from warmer, 304

saltier sub-surface waters, and by deep water production by brine rejection along the 305

Norwegian coast. 306

The Nordic Seas during Greenland interstadials 307

In the interstadial phase of the D-O cycle, when Greenland is warm (Figure 3a), the 308

Nordic Seas are thought to be mostly free of sea ice. Without the presence of sea ice and 309

an associated fresh surface layer, the water column is weakly stratified and the 310

incoming Atlantic water efficiently releases heat to the atmosphere, as is the case in 311

today’s Nordic Seas (Figure 4b). As a consequence, subsurface temperatures in the 312

Nordic Seas are found to be relatively cold (Figure 3b), and any fresh water that may be 313

deposited in the surface layer is easily removed by strong wind mixing, particularly 314

during the winter. 315

As shown by Li et al. (2005), reduced North Atlantic sea ice cover can account for the 316

warm temperatures observed in the Greenland ice cores during interstadials of the last 317

glacial cycle (Figure 1). Due to the exceptional chronology of the high resolution 318

marine sediment record, we can ubiquitously link the timing of warm interstadial 319

conditions on Greenland to periods in the marine record with high benthic δ18O (Figure 320

3d) and well ventilated conditions in the Nordic Seas as evidenced by planktonic δ13C 321

(Figure 3f). The high benthic δ18O values indicate less efficient brine formation 322

combined with lower temperatures, supporting the presence of reduced sea ice cover 323

during interstadials in the Nordic Seas (Dokken & Jansen, 1999). Note that the benthic 324

δ13C record (Figure 3g) is not as straightforward to interpret as the planktonic δ13C 325

15

record: the strong overshoot seen during the transitions between stadial and interstadial 326

conditions is discussed in detail in section 4. 327

4. IMPLICATIONS FOR D-O CYCLES 328

Transition between stadial and interstadial phases 329

The records presented thus far support the idea of a reorganization of the vertical 330

thermohaline structure of the Nordic Seas between the stadial and interstadials phases of 331

D-O cycles. In the interstadial phase, conditions in the Nordic Seas resemble those 332

existing today: the heat transported into the region by the ocean and stored seasonally in 333

the mixed layer is released to the atmosphere, resulting in a moderate, seasonal sea ice 334

cover. In the stadial phase, however, conditions in the Nordic Seas resemble those 335

existing in the Arctic Ocean today: there is a fresh surface layer buffered from the warm 336

Atlantic water below by a halocline. In essence, there is an expansion of sea ice 337

southwards from the Arctic into the Nordic Seas. Indeed, the stratification and vertical 338

structure of the Nordic Seas inferred from our proxy data (Figure 3 & 4) are similar to 339

those of today’s Arctic, where water is several (~3°C) degrees warmer at a depth of 340

400—500 m than at the surface in winter (e.g. Rudels et al., 2004). 341

During the stadial phase, the planktonic foraminfera are mainly recording the 342

temperature of water within, or just below, the halocline (Figure 4a). The sea ice cover 343

prevents wind mixing, and sea ice formation along the Norwegian coast creates dense 344

brines that are rejected from the surface to deeper in the water column, thus helping to 345

maintain the halocline. As the stadial phase progresses, the planktonic foraminfera show 346

an increase in temperature (Figure 3b) consistent with their depth habitat being 347

continuously fed by the northward advection of relatively warm and salty Atlantic 348

16

water. With no possibility of venting heat to the atmosphere above, because of extensive 349

sea ice cover, the warming due to the inflow of warm Atlantic water gradually reduces 350

the density of the subsurface waters, and weakens the stratification that allows the 351

halocline and sea ice cover to exist. The transition to warm interstadial conditions on 352

Greenland occurs when the stratification and halocline weaken to the point of collapse, 353

at which point heat from the subsurface layer is rapidly mixed up to the surface, melting 354

back the sea ice. 355

The removal of sea ice in the Nordic Seas allows efficient exchange of heat between the 356

ocean and the atmosphere, and is seen as an abrupt warming (D-O event) of up to 15°C 357

in Greenland at the start of each interstadial (NGRIP, 2003; Figure 1). There is a distinct 358

overshoot (< 100 years long) in foraminifera-derived subsurface temperatures right at 359

the transition (Figure 3b). The planktonic foraminifera live at the interface between the 360

surface and subsurface layers, and immediately feel the temperature effect as warm 361

subsurface water is mixed up to the surface. Once the fresh surface layer has been 362

mixed away, the ocean is able to vent heat to the atmosphere. The foraminifera now find 363

themselves living in a new habitat, an unstratified surface layer, which is colder than 364

their subsurface stadial habitat. 365

Once in the interstadial (Figure 4b), conditions are comparable to those of today: the 366

Nordic Seas are relatively ice free; inflowing warm Atlantic water is in contact with the 367

surface; and there is an efficient release of heat to the atmosphere. The planktonic-based 368

temperatures show little change (Figure 3b), but there is a consistent reduction in 369

salinity, in particular towards the end of each interstadial (Figure 3c). This freshening of 370

the upper part of the Nordic Seas water column is most likely due to melting and 371

17

calving of the nearby Fennoscandian ice sheet in response to the warm climate 372

conditions, as is supported by the increased frequency of IRD deposits at the site 373

(Figure 3e). 374

Supporting evidence from δ13C 375

The carbon isotope data from planktonic (N. Pachyderma sin.) and benthic (C. teretis) 376

foraminifera at the site (Figure 3f and g) give additional clues about the transition 377

between stadial and interstadial phases in the Nordic Seas. The isotopic signature of 378

carbon (δ13C) preserved in the calcium carbonate shells of foraminfera is related to 379

ventilation and water mass age: at the surface CO2 exchange with the atmosphere and 380

marine photosynthesis preferentially extracts 12C from seawater, causing enrichment of 381

surface water in dissolved inorganic 13C. When the water mass is isolated from the 382

surface mixed layer, its δ13C value decreases with age due to mixing with different 383

water masses and gradual decomposition of low δ13C organic matter. Although the 384

benthic foraminifera C. teretis do not live directly on the surface of the marine 385

sediments, evidence from studies in the Nordic Seas show that the δ13C of C. teretis 386

tracks the δ13C recorded by epibenthic species (benthic fauna living on top of the 387

sediment surface at the seafloor). In the absence of epibenthic foraminifera in high-388

deposition regions such as the core site, C. teretis can be used with caution to infer past 389

changes in deep water δ13C (Jansen et al., 1989). 390

In the stadial phase, with extensive sea ice, relatively young water of Atlantic origin 391

with high δ13C (Figure 3f) enters the Nordic Seas below the fresh surface layer and 392

halocline. At the onset of the interstadial there is a brief period with extremely light 393

benthic δ13C (Figure 3f). This is consistent with a mixing in of old, poorly ventilated 394

18

deep water masses of Arctic origin from below the Atlantic layer in the Nordic Seas 395

(e.g. Thornally et al., 2011). Enhanced mixing at the transition will also bring well-396

ventilated waters with high δ13C down from the surface; however, this signal is 397

overwhelmed by the mixing in of the old water from below with extremely light δ13C. 398

Early in the interstadial phase, following the abrupt transition, the core site is 399

dominated by high δ13C as seen in both the benthic (Figure 3f) and planktonic (Figure 400

3e) records. At this point, the Nordic Seas are free of sea ice, well ventilated, and the 401

stratification is weak due to efficient winter mixing and absence of a halocline. 402

Continuous input of terrestrial freshwater from melting ice sheets, such as the 403

Fennoscandian, also contribute to the high planktonic δ13C. However, towards the end 404

of the interstadial the influence of sea ice is increasing in the Nordic Seas (Figure 3d) 405

and the stratification is increasing, as evidenced by decreasing surface salinity (Figure 406

3b). As a consequence, benthic δ13C is reduced (Figure 3f) indicating increased 407

stratification and less ventilation of waters at the depth of the core site. 408

Returning to the stadial phase (Figure 5I), with an extensive sea ice cover and return of 409

the halocline, the surface and intermediate waters are now isolated from the atmosphere, 410

as evidenced by relatively low planktonic δ13C values (Figure 3e). Although the 411

qualitative features we have described are robust for each D-O cycle, there are 412

quantitative differences between each cycle. For example: convection may reach deeper 413

during some interstadials than others, making the light spike more pronounced; or the 414

ocean may re-stratify more quickly during some interstadials, resulting in a more 415

immediate recovery to heavier δ13C values. Such differences will also be reflected in 416

the other proxy records. 417

19

5. DISCUSSION 418

The conceptual model for D-O cycles presented in the previous section, and 419

summarized in figures 5 and 6, is based on the proxy records measured in MD992284, 420

and specifically, on the relationship between signals in these marine records and signals 421

in the Greenland oxygen isotope record. The marine records themselves provide direct 422

information on the water column in the Atlantic inflow region of the Nordic Seas. 423

However, the expected changes in sea ice cover illustrated in the conceptual model have 424

implications for the response and role of other parts of the climate system in D-O 425

cycles. We have focused on the Greenland oxygen isotope record as a reference time 426

series for D-O cycles, but additional information is provided by the Greenland 427

deuterium excess record (Jouzel et al., 2005; Masson-Delmotte et al., 2005; Thomas et 428

al., 2009). Deuterium excess shows a rapid decrease in the temperature of source waters 429

for Greenland precipitation at stadial to interstadial transitions. Assuming that the 430

abrupt transition from stadial to interstadial conditions is mediated by the removal of 431

sea ice in the Nordic Seas, this would provide a new, local source of moisture for 432

precipitation on the Greenland ice sheet during warm interstadials. The ocean 433

temperature of this source would be significantly lower than sources further to the south 434

(Sachs & Lehman, 1999), thus explaining the observed 3°C rapid decrease in the source 435

temperature for precipitation on Greenland during phases of rapid increase in surface 436

air temperature of D-O cycles (Steffensen et al., 2008). 437

On a larger scale, changes in sea ice cover are expected to be associated with a 438

reorganization of atmospheric circulation, as is suggested by changes in dust delivered 439

to Greenland (Mayewski et al., 1997). We hypothesize that during interstadials, the 440

20

storm track extends northeastwards into the Nordic Seas, transporting heat into the 441

region and inhibiting formation of sea ice by mechanical mixing. During stadials, the 442

storm track is more zonal and positioned south of the Fennoscandian ice sheet, allowing 443

for the quiescent conditions required to maintain strong katabatic winds and sea ice 444

production in leads and polynyas on the continental shelf. Shifts in the orientation of the 445

Atlantic storm track could help explain the dust record in Greenland, as well as far field 446

monsoon and precipitation changes associated with D-O cycles (Pausata et al., 2011). 447

Many previous models of D-O cycles require, or assume, a flux of freshwater to the 448

North Atlantic mainly from the Laurentide ice sheet in order to maintain extensive sea 449

ice cover over the Nordic Seas during stadials. In contrast, the provenance of IRD in 450

MD992284 indicates that the freshwater forcing originates from the Fennoscandian ice 451

sheet, and is associated with interstadials as much as, if not more than, stadials. This 452

provides a local freshwater source in the path of the Atlantic inflow in a highly sensitive 453

area of the Nordic Seas, which can aid more directly in creating a halocline and 454

facilitating expansion of sea ice compared to freshwater sources on the western rim of 455

the North Atlantic. 456

We further expect that the presence of a halocline and extensive sea ice during stadials 457

greatly reduces the ventilation of the deep ocean as it prevents contact between 458

inflowing Atlantic water and the atmosphere in the Nordic Seas. This will impact the 459

properties of the return flow of Atlantic water leaving the Nordic Seas, and should be 460

recorded by intermediate and deep water proxies throughout MIS3. The marine proxy 461

data presented in this study support the existence of an active circulation of relatively 462

warm Atlantic water into the Nordic Seas during stadials, but do not give a direct 463

21

measure of the rate of Atlantic meridional overturning circulation. Evaluating changes 464

in the AMOC as a response to enhanced stratification in the Nordic Seas during stadials 465

is beyond the scope of this study. However, such changes are expected, as the returning 466

outflow of Atlantic water from the Nordic Seas would not be sufficiently dense to sink 467

into the deep abyss during stadials. As discussed in several studies (Crowley, 1992; 468

Stocker & Johnsen, 2003), a reduction in the AMOC could lead to warming of the 469

Southern Ocean and possibly explain the “seesaw” pattern with warm temperatures on 470

Antarctica when Greenland is cold (EPICA , 2006). 471

MIS3 records from other Nordic Seas marine cores share many consistent features with 472

the records presented in this study. However, different interpretations of these features 473

lead to hypothesized mechanisms for D-O cycles that are different from ours in 474

important ways. 475

For example, Rasmussen & Thomsen (2004) suggest that there is warming at depth 476

during stadials, based on the depleted benthic δ18O from a sediment core obtained from 477

the central Nordic Seas basin at 1000 m depth. We also observe depleted benthic δ18O 478

during stadials at our core site, but we interpret this primarily to be due to brine 479

production. Given the position of our core on the Norwegian slope, interpreting the 480

depleted benthic δ18O as a temperature signal would imply temperatures 4-6°C warmer 481

during stadial periods at a depth of 1500 m in the Nordic Seas, which we find unlikely. 482

Another example, Petersen et al. (2013), suggest that the duration of the interstadial and 483

stadial phases of the D-O cycles are related to ice shelves, while sea ice changes provide 484

the abrupt transitions. Critical support for this ice shelf mechanism is an increase in ice-485

rafting during each D-O stadial inferred from IRD and fresh surface anomalies (based 486

22

on planktonic δ18O) in the Nordic Seas and Irminger Basin (Dokken and Jansen 1990, 487

van Kreveld et al. 2000, Elliot et al. 1998, 2001). However, the postulated occurrence of 488

ice rafting during cold Greenland stadials is inconsistent with the chronology of our 489

marine sediment core. The occurrence of ice rafting at our site in the Nordic Seas is 490

enhanced during the latter part of the interstadials (Figure 3e), although our 491

hypothesized reorganization of the Nordic Seas circulation during D-O cycles is not 492

dependent on the input of meltwater from melting ice sheets. 493

Theses issues show that open questions remain, requiring better spatial coverage of 494

marine sediment cores with adequate chronology to pin down the correct timing of 495

reconstructed oceanic signals compared to the ice core records. 496

Our conceptual model of D-O cycles requires the presence of a shallow ocean over a 497

continental shelf adjacent to the Fennoscandian ice sheet. Should the Fennoscandian ice 498

sheet extend over the continental shelf, the stadial phase of the proposed D-O cycle 499

could not exist because brine production through sea ice formation would not be 500

possible and thus a stable halocline could not be maintained in the Nordic Seas. Hence, 501

consistent with observations, our conceptual model precludes the possibility of D-O 502

cycles at the Last Glacial Maximum, when the Fennoscandian ice margin was located 503

near the shelf break. Furthermore, with less efficient brine production at LGM, we 504

expect sea ice formation will be reduced compared to a typical stadial, resulting in ice-505

free conditions over large parts of the Nordic Seas (Hebbeln et al., 1994). 506

Finally, we can draw an analogy between the stadial phase in the Nordic Seas during the 507

last glacial period and the Arctic Ocean today. Both feature a fresh surface layer, 508

halocline, and extensive sea ice. The existence of these features is sensitive to the heat 509

23

transported by inflowing Atlantic water and to brine release during sea ice formation on 510

the continental shelves. Projections of global warming (Holland & Bitz, 2003) are 511

consistent with observed trends (Smedsrud et al., 2008) showing increased heat 512

transport to the Arctic Ocean by the Atlantic inflow. Together with reduced sea ice 513

formation this will weaken the halocline, and could rapidly tip the Arctic Ocean into a 514

perennially ice-free phase. In such a phase, warm Atlantic water will be brought to the 515

surface and have a severe impact on Arctic climate and the stability of the Greenland ice 516

sheet. 517

518

24

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Stage (MIS) 3: a database, Quaternary Science Reviews, 21, 1185-1212, 2002. 721

Waelbroeck, C., Labeyrie, L., Michel, E., Duplessy, J. C., McManus, J. F., Lambeck, 722

K., Balbon, E. and Labracherie, M., Sea level and deep water temperature changes 723

derived from benthonic foraminifera isotopic records, Quaternary Science Rev., 21, 724

295-305, 2002. 725

32

Wagner, J.D.M. et al. Moisture variability in the southwestern United States linked to 726

abrupt glacial climate change. Nature Geosciences, 3, 110-113, 2010. 727

Wastegard, S.; Rasmussen, T. L.; Kuijpers, A.; Nielsen, T. & van Weering, T. C. E., 728

Composition and origin of ash zones from Marine Isotope Stages 3 and 2 in the North 729

Atlantic, Quaternary Science Reviews, 25, 2409-2419, 2006. 730

731

33

732

Table 1: Radiocarbon dated intervals used as a first approximation to create the age 733

model. All AMS 14C ages are measured on N. pachyderma sin. Also shown in the table 734

are two identified ash layers. The ages given for Fugloyarbanki Tephra (FMAZII) and 735

Faroe Marine Ash Zone III (FMAZIII) are from Svensson et al. (2008). The ash ages 736

are referred to in b2k (before year A.D. 2000). Sample with laboratory reference “TUa-737

“ are measured at the Uppsala Accelerator in Sweden, but the samples have been 738

prepared at the radiocarbon laboratory in Trondheim, Norway. Samples labelled “POZ-739

“ are prepared and measured in the radiocarbon laboratory in Poznan, Poland. 740

741

Laboratory Reference and ash

identity

Mean depth

(cm)

14C Age, uncorrected

(yr BP)

Error,1σ

(yr)

Calibrated ages - intercept

(kyr)

Max. 1σ

cal. age (yr BP)

Min. 1σ

cal. age (yr BP)

TUa-3310 1100,5 21975 160 24,64 25018 24270

Fugloyarbanki ash (FMAZ II)

1407,5 26, 69 (b2k)

POZ-29522 2106,5 29100 240 34,06 33770 34350

POZ-29523 2324,5 29900 600 34,87 34260 35480

POZ-17620 2775,5 29920 240 35,5 36000 35000

POZ-17621 2996,5 32500 300 37,5 37800 37100

Faroe Marine Ash Zone III (FMAZ III)

3040,5 38,07 (b2k)

POZ-29524 3075,5 34600 700 39,53 38830 40230

34

Figure 1: Surface air temperature response during the extended winter season 742

(December-April) to changes in Nordic Seas ice cover (Figure redrawn from 743

simulations in Li et al., 2010). Black contours indicate maximum (March) ice extent in 744

the two scenarios. The white dot is the location of the core. 745

Figure 2: (a) δ18O of NGRIP plotted versus depth. (b) Measured low-field magnetic 746

susceptibility, Anhysteretic Remanent Magnetization (ARM) (log scale/unit 10-6A/m) in 747

MD992284 plotted versus depth. Red stars indidate the stratigraphical level (depths) 748

from where AMS 14C ages are measured (see also Table 1). (c) NGRIP δ18O and ARM-749

record from MD992284 after converting the MD992284 age to the NGRIP age (b2k). 750

Green vertical lines indicate tuning points for the (numbered) onset of interstadials in 751

the ice core- and marine records. Red vertical lines connect identical ash layers 752

(FMAZIII and FMAZII) in NGRIP and MD992284 as described in the text and in Table 753

1. 754

Figure 3: Down core data sets from NGRIP (a) and MD992284 (b-g) covering the 755

period 41ka to 31ka plotted on the GICC05 (b2k) time scale. a) NGRIP δ18O (proxy for 756

Greenland temperature), b) reconstructed temperature based on planktonic foram 757

assemblages, c) near surface δ18OSMOW, d) benthic δ18O, e) flux corrected Ice Rafted 758

Debris (IRD), f) near surface δ13C, and g) benthic δ13C. Interstadial periods are 759

indicated by grey shading. The near-surface δ18OSMOW is determined from the Kim and 760

O’Neil (1997) calibration curve using the δ18O of N. pachyderma sin. and the 761

temperature obtained using transfer functions (see section 2). The near surface δ13C is 762

measured on N. pachyderma sin. and the benthic δ18O and δ13C are measured on 763

Cassidulina teretis 764

35

Figure 4: A schematic column model of the stadial (left) and interstadial (right) phases 765

consistent with the Nordic Sea sediment core MD992284. 766

Figure 5: A schematic timeline showing the evolution of a D-O cycle that is consistent 767

with the Nordic Sea sediment core MD992284 and the NGIP ice core. 768

Figure 6: Schematic showing wintertime conditions in the North Atlantic and Nordic 769

Seas during typical cold stadial periods (left column) and warm interstadial periods 770

(right column) of a D-O cycle. Top panels show maps of the North Atlantic region with 771

sea ice extent and land ice. The location of the sediment core is marked with a blue dot 772

and the sections described in the lower panels are indicated. Middle panels A) and C) 773

show north-south sections of the North Atlantic during stadial and interstadial 774

conditions, respectively. Bottom panels B) and D) show east-west sections of the 775

Norwegian Sea during stadial and interstadial conditions, respectively. 776

0

5

10

15

20

25

30

26 28 30 32 34 36 38 40

3 4 5 6 7 8 9 10

3 4 5 6 7 8 9 10

teph

ra (F

MA

Z II)

1400 1600 1800 2000 2200 2400 2600 2800 3000 3200Core depth (cm)

N-GRIP (GICC05 age)

teph

ra (F

MA

Z III

)

Mon

o La

ke

Lasc

ham

p ev

ent

ARM

(*10

00)

100

200

300

400

b18O

NG

RIP

3 4 5 6 7 8 9 10

teph

ra (F

MA

Z II)

teph

ra (F

MA

Z III

)

Mono Lake Lachamp event

26 28 30 32 34 36 38 40

N-GRIP (GICC05 age - ka)

42

ARM

(*10

00)

100

200

300

400

b18O

NG

RIP

-36

-38

-40

-42

-44

-46

-36

-38

-40

-42

-44

-46

-34

a)

c)

b)

d)

e)

f)

g)

-36

-38

-40

-42

-44

-46

NG

RIP b18

O

Reco

nstr

ucte

d SS

T (O

C)

1098765

424140393837363534333231

424140393837363534333231

Interstadial numbersH4H3

b18O

(sm

ow) N

.pac

hyde

rma

sin.

-2,5

-2,0

-1,5

-1,0

-0,5

-0

0,5

0

2

4

6

8

10

b18O

C. t

eret

is (b

enth

ic)6,0

5,5

5,0

4,5

4,0

3,5

b13C

N.p

achy

derm

a si

n.

-0,3

-0,1

0,5

0,3

0,1

-1,8

-1,4

-0,4

-0,8

-1,2

b13C

C. t

eret

is (b

enth

ic)

Age (Ka) (GICC5)

15

10

5

IRD

(No.

gra

in>

0,5m

m/g

300 m

600 m

100 m

Halocline

Fresh surface

Brine

fossil/aged/leftoverNordic Sea water

(intermediate and deep/ventilated)Nordic Sea water

GlacialAtlantic inflow Atlantic inflow

open ocean increasingwinter ice perennial sea iceperennial sea ice open ocean

time

Greenlandtemperature

Subsurfacesalinity

Brine production /Sea ice cover

Interstadial conditionsStadial conditions

Equator Arctic

Nordic Seas

Gre

enla

nd -

Scot

land

rid

ge

Nordic Seas

Gre

enla

nd -

Scot

land

rid

ge

Sea ice Cold, fresh water

Equator Arctic

Norwegian sea Norway Norwegian sea Norway

Cold air freezes seawater

Sea ice

Cold, fresh water Brine release

500 meters 500 meters

Cold, saltier water

Warm, saltyAtlantic water

Denser, salty water sinks200 meters

Newly formed ice

Halocline

Heat loss

Heat loss

Northwardheat transport

Northward heat transport

A

A

B

B

C

C

D

D

Figure 3