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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
10
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
11
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
12
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
13
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
6. References 519
Ãlvarez-Solas, J.; Montoya, M.; Ritz, C.; Ramstein, G.; Charbit, S.; Dumas, C.; 520
Nisancioglu, K.; Dokken, T. & Ganopolski, A., Heinrich event 1: an example of 521
dynamical ice-sheet reaction to oceanic changes, Clim. Past, Copernicus Publications, 522
7, 1297-1306, 2011. 523
Bigg, G.R. and E.J. Rohling, E.J, An oxygen isotope data set for marine waters. J. 524
Geophys. res., 105, 8527-8536, 2000. 525
Birchfield, G.E., and W.S. Broecker, A salt oscillator in the glacial ocean? Part II: A 526
'scale analysis' model, Paleoceanography, 5, 835-843, 1990. 527
Bond, G., Broecker, W., Johnsen, S., McManus, J., Labeyrie, L., Jouzel, J. & Bonani, 528
G. Correlations between climate records from North Atlantic sediments and Greenland 529
ice, Nature, 365, 143-147, 1993. 530
Broecker, W.S., Bond, G. and Klas, M., A salt oscillation in the glacial Atlantic? 1. the 531
concept, Paleoceanography, 5, 469, 1990. 532
Broecker, W.S., Peteet, D.M. and Rind, D., Does the ocean-atmosphere system have 533
more than one stable mode of operation, Nature, 315, 21-26, 1985. 534
Broecker, W., Abrupt climate change: causal constraints provided by the paleoclimate 535
record, Earth-Science Reviews, 51, 137-154, 2000. 536
Carstens, J.; Hebbeln, D. & Wefer, G., Distribution of planktic foraminifera at the ice 537
margin in the Arctic (Fram Strait), Marine Micropaleontology, 29, 257-269, 1997. 538
Chiang, J. C. H.; Biasutti, M. & Battisti, D. S., Sensitivity of the Atlantic Intertropical 539
Convergence Zone to Last Glacial Maximum boundary conditions, Paleoceanography, 540
18, 2003. 541
Crowley, T.J., North Atlantic deep water cools the Southern Hemisphere, 542
Paleoceanography, 7, 489-497, 1992. 543
25
Curry, W.B., and Oppo, D.W., Synchronous, high-frequency oscillations in tropical sea 544
surface temperatures and North Atlantic Deep Water production during the last glacial 545
cycle, Paleoceanography, 12, 1–14, 1997. 546
Dansgaard, W., Johnsen, S. J.; Clausen, H. B.; Dahl-Jensen, D.; Gundestrup, N. S.; 547
Hammer, C. U.; Hvidberg, C. S.; Steffensen, J. P.; Sveinbjornsdottir, A. E.; Jouzel, J. & 548
Bond, G. , Evidence for general instability of past climate from a 250-kyr ice core 549
record. Nature, 364, 218-220, 1993. 550
Davies, S.M., Wastegård, S.; Rasmussen, T. L.; Svensson, A.; Johnsen, S. J.; 551
Steffensen, J. P. & Andersen, K. K., 2008, Identification of the Fugloyarbanki tephra in 552
the NGRIP ice core: a key tie point for marine and ice-core sequences during the last 553
glacial period, Journ. of Quaternary Science, 23(5), p. 409-414. 554
Dokken, T.M. and Jansen, E., Rapid changes in the mechanism of ocean convection 555
during the last glacial period, Nature, 401, 458-461, 1999. 556
EPICA community members, One-to-one coupling of glacial climate variability in 557
Greenland and Antarctica, Nature, 444, 195-198, 2006. 558
Fairbanks, R. G., Mortlock, R. A., Chiu, T. C., Cao, L., Kaplan, A., Guilderson, T. P., 559
Fairbanks, T. W., Bloom, A. L., Grootes, P. M. and Nadeau, M. J., Radiocarbon 560
calibration curve spanning 0 to 50,000 years BP based on paired Th-230/U-234/U-238 561
and C-14 dates on pristine corals, Quaternary Science Reviews, 24, 1781-1796, 2005. 562
Fairbanks, R.G., A 17,000-year glacio-eustatic sea-level record - influence of glacial 563
melting rates on the younger dryas event and deep-ocean circulation, Nature, 342, 637-564
642, 1989. 565
Ganopolski, A. and Rahmstorf, S., Rapid changes of glacial climate simulated in a 566
coupled climate model, Nature, 409, 153-158, 2001. 567
Genty, D., D. Blamart, R. Ouahdi, M. Gilmour, A. Baker, J. Jouzel and Sandra Van-568
Exter, Precise dating of Dansgaard–Oeschger climate oscillations in western Europe 569
from stalagmite data, Nature, 421, 833-837, 2003. 570
26
Gildor, H. & Tziperman, E., Sea-ice switches and abrupt climate change 571
Philosophical Transactions of the Royal Society A-mathematical Physical and 572
Engineering Sciences, Royal Soc, 2003, 361, 1935-1942. 573
Hebbeln, D., Dokken, T., Andersen, E. S., Hald, M., and Elverhoi, A., Moisture Supply 574
For Northern Ice-Sheet Growth During the Last- Glacial-Maximum, Nature, 370, 357-575
360, 1994. 576
Hendy, I. L.; Kennett, J. P.; Roark, E. B. & Ingram, B. L., Apparent synchroneity of 577
submillennial scale climate events between Greenland and Santa Barbara Basin, 578
California from 30-10 ka, Quaternary Science Reviews, 21, 1167-1184, 2002. 579
Holland, M. M. & Bitz, C. M., Polar amplification of climate change in coupled models, 580
Climate Dynamics, 21, 221-232, 2003. 581
Huber, C.; Leuenberger, M.; Spahni, R.; Flückiger, J.; Schwander, J.; Stocker, T. F.; 582
Johnsen, S.; Landais, A. & Jouzel, J., Isotope calibrated Greenland temperature record 583
over Marine Isotope Stage 3 and its relation to CH4, Earth and Planetary Science 584
Letters, 243, 504-519, 2006. 585
Hughen, K.A., Baillie, M.G.L., Bard, E., Beck, J.W., Bertand, C.J.H., Blackwell, P.G., 586
Buck, C.E., Burr, G.S., Cutler, K.B., Damon, P.E., Edwards, R.L., Fairbanks, R.G., 587
Friedrich, M., Guilderson, T.P., Kromer, B., McCormac, G., Manning, S., Ramsey, 588
C.B., Reimer, P.J., Reimer, R.W., Remmele, S., Southon, J.R., Stuiver, M., Talamo, S., 589
Taylor, F.W., van der Plicht, J., Weyhenmeyer, C.E., Marine04 Marine radiocarbon age 590
calibration, 0–26 ka BP, Radiocarbon, 46, 1059–1086, 2004. 591
Jansen, E.; Slettemark, B.; Bleil, U.; Henrich, R.; Kringstad, L. & Rolfsen, S., Oxygen 592
and carbon isotope stratigraphy and magnetostratigraphy of the last 2.8 Ma: 593
paleoclimatic comparisons between the Norwegian Sea and the North Atlantic, 594
Proceedings of the Ocean Drilling Program Scientific Results, 104, 255-269, 1989. 595
Jouzel, J., Masson-Delmotte, V., Stievenard, M., Landais, A., Vimeux, F., Johnsen, S.J., 596
Sveinbjornsdottir, A.E. and White, J.W., Rapid deuterium-excess changes in Greenland 597
27
ice cores: a link between the ocean and the atmosphere, Comptes Rendus Geoscience, 598
337, 957-969, 2005. 599
Kim, S.-T. and O'Neil, J.R., Equilibrium and nonequilibrium oxygen isotope effects in 600
synthetic carbonates, Geochimica et Cosmochimica Acta, 61, 3461-3475, 1997. 601
Kissel, C., Laj, C., Labeyrie, L., Dokken, T., Voelker, A., Blamart, D., 1999. Rapid 602
climatic variations during marine isotope stage 3: magnetic analysis of sediments from 603
Nordic Seas and North Atlantic, Earth and Planetary Science Letters, 171, 489–502. 604
Knutti, R., Fluckiger, J., Stocker, T. and Timmermann, A., Strong hemispheric coupling 605
of glacial climate through freshwater discharge and ocean circulation, Nature, 430, 851-606
856, 2004. 607
Landais, A.; Caillon, N.; Severinghaus, J.; Barnola, J. M.; Goujon, C.; Jouzel, J. & 608
Masson-Delmotte, V., Isotopic measurements of air trapped in ice to quantify 609
temperature changes., Comptes Rendus Geoscience, 336, 963-970, 2004. 610
Lang, C.; Leuenberger, M.; Schwander, J. & Johnsen, S., 16°C Rapid Temperature 611
Variation in Central Greenland 70,000 Years Ago, Science, 286, 934-937, 1999. 612
Lewis, S. C.; LeGrande, A. N.; Kelley, M. & Schmidt, G. A., Water vapour source 613
impacts on oxygen isotope variability in tropical precipitation during Heinrich events, 614
Clim. Past, 6, 325-343, 2010. 615
Li, C.; Battisti, D. S.; Schrag, D. P. & Tziperman, E., Abrupt climate shifts in 616
Greenland due to displacements of the sea ice edge, Geophysical Research Letters, 32, 617
L19702, 2005. 618
Li, C.; Battisti, D. S. & Bitz, C. M., Can North Atlantic Sea Ice Anomalies Account for 619
Dansgaard-Oeschger Climate Signals?, Journal of Climate, 23, 5457-5475, 2010. 620
Manabe, S., and R. J. Stouffer, Two stable equilibria of a coupled ocean-atmosphere 621
model, Journal of Climate, 1, 841-866, 1988. 622
28
Marcott, S. A.; Clark, P. U.; Padman, L.; Klinkhammer, G. P.; Springer, S. R.; Liu, Z.; 623
Otto-Bliesner, B. L.; Carlson, A. E.; Ungerer, A.; Padman, J.; He, F.; Cheng, J. & 624
Schmittner, A., Ice-shelf collapse from subsurface warming as a trigger for Heinrich 625
events, Proceedings of the National Academy of Sciences, 2011. 626
Marotzke, J. and Willebrand, J., Multiple Equilibria of the Global Thermohaline 627
Circulation, Journal of Physical Oceanography, 21, 1372-1385, 1991. 628
Masson-Delmotte, V., Jouzel, J., Landais, A., Stievenard, M., Johnsen, S.J., White, 629
J.W.C., Werner, M., Sveinbjornsdottir, A. and Fuhrer, K., GRIP deuterium excess 630
reveals rapid and orbital-scale changes in Greenland moisture origin, Science, 309, 118-631
121, 2005. 632
Mayewski, P.A., Meeker, L.D., Twickler, M.S., Whitlow, S.I., Yang, Q., Lyons,W.B. 633
and Prentice, M., Major features and forcing of high latitude northern hemisphere 634
atmospheric circulation over the last 110,000 years, Journal of Geophysical Research, 635
102, 26,345-26,366, 1997. 636
Mignot, J.; Ganopolski, A. & Levermann, A., Atlantic subsurface temperatures: 637
Response to a shutdown of the overturning circulation and consequences for its 638
recovery, Journal of Climate, 20, 4884-4898, 2007. 639
North Greenland Ice Core Project members, High-resolution record of Northern 640
Hemisphere climate extending into the last interglacial period, Nature, 431, 147–151, 641
2004. Rev., 23, 1513–1522, 2004. 642
North Greenland Ice Core Project members, High-resolution record of Northern 643
Hemisphere climate extending into the last interglacial period, Nature, 431, 147–151, 644
2004. Rev., 23, 1513–1522, 2004. 645
North Greenland Ice Core Project members: High-resolution record of Northern 646
Hemisphere climate extending into the last inter-glacial period, Nature, 431, 147–151, 647
2004. 648
29
Otto-Bliesner, B. L. & Brady, E. C., The sensitivity of the climate response to the 649
magnitude and location of freshwater forcing: last glacial maximum experiments, 650
Quaternary Science Reviews, 29, 56-73, 2010. 651
Paillard, D.L., Labeyrie, L., and Yiou, P., Macintosh program performs time-series 652
analysis, EOS Transactions AGU, 77, 379, 1996. 653
Pausata, F. S. R.; Battisti, D. S.; Nisancioglu, K. H. and Bitz, C. M., Chinese stalagmite 654
delta(18)O controlled by changes in the Indian monsoon during a simulated Heinrich 655
event, Nature Geoscience, 4, 474-480, 2011. 656
Petersen, S. V.; D. P. Schrag, and P. Clark, A new mechanism for Dansgaard-Oeschger 657
cycles, Paleoceanography, doi:10.1029/2012PA002364, in press (2013). 658
Peterson, L.C., Haug, G.H., Hughen, K.A. and Rohl, U., Rapid changes in the 659
hydrologic cycle of the tropical Atlantic during the last glacial, Science, 290, 1947-660
1951, 2000. 661
Pflaumann, U.; Sarnthein, M.; Chapman, M.; d'Abreu, L.; Funnell, B.; Huels, M.; 662
Kiefer, T.; Maslin, M.; Schulz, H.; Swallow, J.; van Kreveld, S.; Vautravers, M.; 663
Vogelsang, E. and Weinelt, M., Glacial North Atlantic: Sea-surface conditions 664
reconstructed by GLAMAP 2000, Paleoceanography, 18, 1-21, 2003. 665
Rahmstorf, S., Ocean circulation and climate during the past 120,000 years, Nature, 666
419, 207-214, 2002. 667
Rasmussen, T. L. & Thomsen, E., The role of the North Atlantic Drift in the millennial 668
timescale glacial climate fluctuations, Palaeogeography, Palaeoclimatology, 669
Palaeoecology, 210, 101-116, 2004. 670
Renssen, H. & Isarin, R., The two major warming phases of the last deglaciation at 671
~14.7 and ~11.5 ka cal BP in Europe: climate reconstructions and AGCM experiments, 672
Global and Planetary Change, 30, 117-153, 2001. 673
30
Rudels, B.; Jones, E. P.; Schauer, U. & Eriksson, P., Atlantic sources of the Arctic 674
Ocean surface and halocline waters, Polar Research, 23, 181-208, 2004. 675
Sachs, J.P. and Lehman, S.J. Subtropical North Atlantic temperatures 60,000 to 30,000 676
years ago, Science, 286, 756-759, 1999. 677
Sachs, J.P. and Lehman, S.J. Subtropical North Atlantic temperatures 60,000 to 30,000 678
years ago, Science, 286, 756-759, 1999. 679
Schulz, H., von Rad, U. and Erlenkeuser, H. Correlation between Arabian Sea and 680
Greenland climate oscillations of the past 110,000 years, Nature, 393, 54-57, 1998. 681
Severinghaus, J.P. and Brook, E.J., Abrupt climate change at the end of the last glacial 682
period inferred from trapped air in polar ice, Science, 286, 930-934, 1999. 683
Siddall, M., Rohling, E.J., Thompson, W.G. and Waelbroeck, C., Marine isotope stage 3 684
sea level fluctuations: data synthesis and new outlook, Reviews of Geophysics 46, 685
RG4003, 2008. 686
Smedsrud, L. H.; Sorteberg, A. & Kloster, K., Recent and future changes of the Arctic 687
sea-ice cover, Geophysical Research Letters, 35, L20503, 2008. 688
Steffensen, J. P., Andersen, K. K., Bigler, M., Clausen, H. B., Dahl-Jensen, D., Fischer, 689
H., Goto-Azuma, K., Hansson, M., Johnsen, S. J., Jouzel, J.; Masson-Delmotte, V., 690
Popp, T., Rasmussen, S. O., Rothlisberger, R., Ruth, U., Stauffer, B., Siggaard-691
Andersen, M. L., Sveinbjornsdottir, A. E., Svensson, A. and White, J. W. C., High-692
resolution Greenland Ice Core data show abrupt climate change happens in few years, 693
Science, 321, 680-684, 2008. 694
Stocker, T. F. and Johnsen, S. J., A minimum thermodynamic model for the bipolar 695
seesaw, Paleoceanography, 18, 2003. 696
Stommel, H., Thermohaline Convection with Two Stable Regimes of Flow, Tellus, 13, 697
224-230, 1961. 698
31
Stuiver, M. and Reimer, P. J., Extended C database and revised CALIB 3.0 14C age 699
calibration program, Radiocarbon, 35, 215–230, 1993. 700
Svensson, A., Andersen, K.K., Bigler, M., Clausen, H.B., Dahl-Jensen, D., Davies, 701
S.M., Johnsen, S.J., Muscheler, R., Rasmussen, S.O., Rothlisberger, R., Steffensen, J.P. 702
and Vinther, B.M., The Greenland Ice Core Chronology, 2005, 15-42 ka. Part 2: 703
comparison to other records, Quaternary Science Reviews, 25, 3258-326, 2006. 704
Svensson, A., Andersen, K. K., Bigler, M., Clausen, H. B., Dahl-Jensen, D., Davies, S. 705
M., Johnsen, S. J., Muscheler, R., Parrenin, F., Rasmussen, S. O., Roethlisberger, R., 706
Seierstad, I., Steffensen, J. P. and Vinther, B. M., A 60 000 year Greenland stratigraphic 707
ice core chronology, Climate Of The Past, 4, 47-57, 2008. 708
Telford, R.J., and Birks, H.J.B., The secret assumption of transfer functions: problems 709
with spatial autocorrelation in evaluating model performance, Quaternary Science 710
Reviews 24, 2173–2179, 2005. 711
Thomas, E.R., Wolff, E.W., Mulvaney, R., Johnsen, S.J., Steffensen, J.P., Arrowsmith, 712
C., Anatomy of a Dansgaard-Oeschger warming transition: High resolution analysis of 713
the North Greenland Ice Core Project ice core, Journal. of Geophys. Res., 114, doi: 714
10.1029/2008JD011215, 2009. 715
Thornalley, D. J. R.; Barker, S.; Broecker, W. S.; Elderfield, H. & McCave, I. N., The 716
Deglacial Evolution of North Atlantic Deep Convection, Science, 331, 202-205, 2011. 717
Vellinga, M. & Wood, R. A., Global climatic impacts of a collapse of the Atlantic 718
thermohaline circulation, Climatic Change, 54, 251-267, 2002. 719
Voelker, A.H.L., Global distribution of centennial-scale records for Marine Isotope 720
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