1 Dansgaard-Oeschger cycles: interactions between ocean and 2 sea ice intrinsic to the Nordic Seas
3 Trond M. Dokken1,3, Kerim H. Nisancioglu2,1,3, Camille Li4,3, David S. Battisti5,3, 4 Catherine Kissel6
5 1. UNI Research AS, Allegaten 55, 5007 Bergen, Norway; 2. Department of Earth Science, University of 6 Bergen, 5007 Bergen, Norway; 3. Bjerknes Centre for Climate Research, 5007 Bergen, Norway; 4. 7 Geophysical Institute, University of Bergen, 5007 Bergen, Norway; 5. Department of Atmospheric 8 Sciences, University of Washington, Seattle WA 98195, USA; 6. Laboratoire des Sciences du Climat et de 9 l’Envionnement, CEA/CNRS/UVSQ, 91198 Gif-sur-Yvette Cedex, France.
10 ABSTRACT
11 Dansgaard-Oeschger (D-O) cycles are the most dramatic, frequent, and wide reaching
12 abrupt climate changes in the geologic record (Voelker, 202). On Greenland, D-O
13 cycles are characterized by an abrupt warming of up to 15°C from a cold stadial to a
14 warm interstadial phase (Dansgaard et al., 1993; Severinghaus & Brook, 1999; NGRIP,
15 2003), followed by gradual cooling before a rapid return to stadial conditions. The
16 mechanisms responsible for these millennial cycles are not fully understood, but are
17 widely thought to involve abrupt changes in Atlantic Meridional Overturning
18 Circulation (AMOC) due to freshwater perturbations (Broecker et al., 1985; Broecker et
19 al., 1990; Birchfield & Broecker, 1990). Here we present a new, high resolution multi-
20 proxy marine sediment core monitoring changes in the warm Atlantic inflow to the
21 Nordic Seas as well as in local sea ice cover and influx of Ice Rafted Debris (IRD). In
22 contrast to previous studies, the freshwater input is found to be coincident with warm
23 interstadials on Greenland and has a Fennoscandian rather than Laurentide source.
24 Furthermore, the data suggest a different thermohaline structure for the Nordic Seas
25 during cold stadials, in which there is a continuous circulation of relatively warm
26 Atlantic water beneath a fresh surface layer topped by sea ice. This implies a delicate
1 27 balance between the warm sub-surface Atlantic water and fresh surface layer, with the
28 possibility of abrupt changes in sea ice cover, and suggests a novel mechanism for the
29 abrupt D-O events observed in Greenland ice cores.
30 1. INTRODUCTION
31 There is a wealth of proxy data showing that the climate system underwent large, abrupt
32 changes throughout the last ice age. Among these are about one score large, abrupt
33 climate changes known as Dansgaard-Oeschger events that occurred every 1-2 kyr.
34 These events were first identified in ice cores taken from the summit of Greenland
35 (Dansgaard et al. 1993; Bond et al., 1993), and are characterized by an abrupt warming
36 of up to 15C in annual average temperature (Severinghaus & Brook, 1999; Lang et al.
37 1999; Landais et al. 2004; Huber et al. 2006). Subsequent observational studies
38 demonstrated the D-O events represent a near-hemispheric scale climate shift (see e.g.
39 Rahmstorf, 2002 for an overview). Warming in Greenland is coincident with warmer,
40 wetter conditions in Europe (Genty et al., 2003), an enhanced summer monsoon in the
41 northwest Indian Ocean (Schulz et al., 1998; Pausata et al., 2011), a northward shift of
42 precipitation belts in the Cariaco Basin (Peterson et al., 2000), aridity in the
43 southwestern United States (Wagner et al., 2010), and changes in ocean ventilation off
44 the shore of Santa Barbara, California (Hendy et al., 2002). Each abrupt warming
45 shares a qualitatively similar temperature evolution: about 1000 years of relatively
46 stable cold conditions, terminated by an abrupt (less than 10 years) jump to much
47 warmer conditions that persist for 200-400 years, followed by a more gradual transition
48 (~50 to 200 years) back to the cold conditions that precede the warming event. This
49 sequence of climate changes is often referred to as a D-O cycle.
2 50 The leading hypothesis for D-O cycles attributes them to an oscillation of the Atlantic
51 Meridional Overturning Circulation (AMOC) (Broecker et al., 1985), with "on" and
52 "off" (or "weak") modes of the AMOC (Stommel, 1961; Manabe & Stouffer, 1985)
53 producing warm interstadials and cold stadials, respectively, through changes in the
54 meridional ocean heat transport. Indeed, interstadial and stadial conditions on
55 Greenland summit are associated with markedly different water mass properties
56 (temperature and salinity) in the subtropical Atlantic (Sachs & Lehman, 1999), North
57 Atlantic (Curry & Oppo, 1997) and Nordic Seas (Dokken & Jansen, 1990). Transitions
58 between the two phases are thought to arise from changes in the freshwater budget of
59 the North Atlantic (Broecker et al., 1990). In particular, Birchfield and Broecker (1990)
60 proposed that the continental ice sheets covering North America and Eurasia were the
61 most likely sources for the freshwater.
62 A number of studies with climate models of intermediate complexity (EMICs) have
63 produced two quasi-stable phases in the AMOC by imposing ad hoc cyclic freshwater
64 forcing (Marotzke & Willebrand, 1991; Ganopolski & Rahmstorf, 2001; Knutti et al.,
65 2004). Although these models simulate a temperature evolution on Greenland Summit
66 that is qualitatively similar to observations, the changes in the simulated local and far-
67 field atmospheric response are small compared to observations. Similar experiments
68 with state-of-the-art climate models produce local and far-field responses that are in
69 better agreement with the proxy data (e.g. Vellinga & Wood, 2002). However, these
70 models require unrealistically large freshwater perturbations, and on Greenland, the
71 simulated stadial to interstadial transitions are not as abrupt as observed. In both cases,
72 the prescribed external forcing is ad hoc, and yet is entirely responsible for the existence
73 and the longevity of the two phases (i.e., the cold phase exists as long as the freshwater
3 74 forcing is applied).
75 Despite the unresolved questions surrounding D-O cycles, there has been progress. The
76 direct, causal agent for the abrupt climate changes associated with D-O warming events
77 is thought to be abrupt reductions in North Atlantic sea ice extent (Broecker 2000;
78 Gildor and Tziperman 2003; Masson-Delmotte et al. 2005; Jouzel et al. 2005). Studies
79 using atmospheric general circulation models (AGCMs) show that the local
80 precipitation and temperature shifts (inferred from oxygen isotopes and accumulation
81 changes) on Greenland summit associated with a typical D-O event are consistent with
82 the response of the atmosphere to a reduction in winter sea ice extent, in particular in
83 the Nordic Seas region (e.g. Figure 1; see also Renssen and Isarin 2001; Li et al 2005;
84 Li et al. 2010). Away from the North Atlantic, the northward shift of tropical
85 precipitation belts associated with D-O warming events mirrors the southward shift
86 simulated by coupled climate models in response to freshwater hosing and expansion of
87 sea ice (Vellinga and Wood 2002; Otto-Bliesner & Brady, 2010; Lewis et al., 2010),
88 and by AGCMs coupled to slab ocean models in response to a prescribed expansion of
89 North Atlantic sea ice (Chiang et al., 2003). Thus supporting the existence of a robust
90 link between North Atlantic sea ice cover and far-field climate.
91 More recently, a number of studies have shown that cold stadial conditions are
92 associated with subsurface and intermediate depth warming in the North Atlantic
93 (Rasmussen & Thomsen 2004; Marcott et al. 2011). This subsurface warming may have
94 bearing on many aspects of D-O cycles, from the abruptness of the warming (stadial-to-
95 interstadial) transitions (Mignot et al., 2007) to the possibility of ice shelf collapses
96 during D-O cycles (Alvarez Solas et al., 2011; Marcott et al., 2011; Petersen et al.,
4 97 2013).
98 In this study, we focus on the subsurface warming rather than ice-rafting to build
99 physical arguments for the mechanisms behind D-O cycles. Changes in sea ice and
100 subsurface ocean temperatures during D-O transitions are linked to each other by the
101 ocean circulation, and to changes in Greenland by the atmospheric circulation. To
102 elucidate these links, requires marine records that (i) are situated in key regions of the
103 North Atlantic, (ii) are representative of conditions through the depth of the water
104 column, and (iii) can be compared to records of atmospheric conditions (as measured
105 from the Greenland summit) with a high degree of certainty in terms of timing. In
106 addition, because of the abruptness of the transitions (on the order of decades or less),
107 the marine records must be of exceptionally high resolution.
108 Here we present new data from a Nordic Seas core that meet these criteria. The core site
109 is strategically located in the Atlantic inflow to the Nordic Seas, where extremely high
110 sedimentation rates allow for decadal scale resolution of both surface and deep-water
111 proxies. In addition to an age model based on calibrated C-14 ages and magnetic
112 properties, ash layers provide an independent chronological framework to synchronize
113 signals in the marine core with signals in an ice core from the summit of Greenland.
114 The new marine proxy data reveal systematic changes in the hydrography of the Nordic
115 Seas as Greenland swings from stadial (cold) to interstadial (warm) conditions and back
116 through several D-O cycles, with details of the transitions better resolved than in any
117 previously studied core. The interstadial phase in the Nordic Seas is shown to resemble
118 current conditions, while the stadial phase has greatly enhanced sea ice coverage and a
119 hydrographic structure similar to that in the Arctic Ocean today. The variations recorded
5 120 by the marine proxies near the transitions are used to determine the possible
121 mechanisms that cause the North Atlantic climate system to jump abruptly from stadials
122 to interstadials and to transition slowly back from interstadials to stadials. In a departure
123 from most existing hypotheses for the D-O cycles, the conceptual model presented here
124 does not require freshwater fluxes from ice sheets to enter the ocean at specific times
125 during the D-O cycle. Rather, it focuses on the role of ocean-sea ice interactions within
126 the Nordic Seas in triggering stadial-interstadial transitions.
127 The paper proceeds as follows. Section 2 contains a description of the marine sediment
128 core, analysis methods, the construction of the age model for the core, and the
129 construction of a common chronology with an ice core from the summit of Greenland.
130 Section 3 describes the cold stadial and warm interstadial climate of the Nordic Seas as
131 seen in the marine sediment data. Section 4 analyzes the relevance of the new proxy
132 data for the D-O cycles as observed on Greenland, and in particular the mechanisms
133 behind the abrupt transitions. Section 5 presents a summary of the results and a
134 discussion of the implications.
135 2. MATERIAL AND METHODS
136 Core MD992284 was collected in the Nordic Seas during the MD114/IMAGES V
137 cruise aboard R.V. Marion Dufresne (IPEV) at a water depth of 1500 m on the
138 northeastern flank of the Faeroe-Shetland channel. The core is strategically located in
139 the Atlantic inflow to the Nordic Seas and has an exceptionally high sedimentation rate,
140 allowing for decadal scale resolution of both surface and deep water proxies. The
141 unique, high-resolution marine core with a well constrained age model (Figure 2) shows
142 evidence for systematic differences in the Nordic Seas during Greenland interstadials
6 143 versus Greenland stadials. These new data are used to characterize the important
144 hydrographic features of the Nordic Seas in each of these two phases of the D-O cycle
145 as well as the transition between the phases.
146 Figure 2a shows a segment of the NorthGRIP ice core (NGRIP) from Greenland
147 (NGRIP, 2004) during late Marine Isotope Stage 3 (MIS3) containing several typical D-
148 O cycles. To compare the marine record to the ice core record, we have constructed an
149 age model by matching rapid transitions in MD992284 to NGRIP using the Greenland
150 Ice Core Chronology 2005 (GICC05; Svensson et al., 2006) and ash layers found in
151 both cores. The common chronology allows the Nordic Seas data to be linked directly to
152 Greenland ice core signals.
153 Age model for marine sediment core
154 Age models based on the ice core and calibrated 14C dates can be used to estimate “true
155 ages” for the marine core, but will be subject to errors due to missing or misinterpreted
156 ice layers, uncertainty in reservoir ages and uncertainty in sedimentation rates. Our
157 approach here is thus not to rely on “true ages”, but rather on a stratigraphic tuning of
158 the marine record to the ice core record. We use the high frequency variations in
159 measured Anhysteretic Remanent Magnetization (ARM) during Marine Isotope Stage 3
160 (MIS3) in the marine core (Figure 2b). The fast oscillations in magnetic properties
161 during MIS3 in the North Atlantic/Nordic Seas have been shown to be in phase with
162 changes in d18O of Greenland ice cores (Kissel et al., 1999). We use the ARM record
163 for tuning only, with no further subsequent paleo-proxy interpretation of this parameter.
164 The tuning has been performed using the tuning points as indicated in Figure 2 and the
165 software package AnalySeries 2.0 (Paillard et al., 1996).
7 166 A final step in the synchronization between the ice core and the marine core is done
167 with tephra layers. Two ash zones identified in the marine record have also been
168 identified in the NGRIP ice core (Svensson et al., 2008), the Fugloyarbanki Tephra
169 (FMAZII) (Davies et al., 2008) and Fareo Marine Ash Zone III (FMAZIII)
170 (Wastegaard et al., 2006). These ashes allow for a direct, independent comparison
171 between the marine and ice core records. Our chronology places the ash layer FMAZIII
172 less than 100 years after the onset of IS8 in NGRIP, and thus provides a verification of
173 the initial tuning of the ARM record in MD992284 to the d18O record in NGRIP. These
174 two ash layers are also used as tuning points in the final age model. The final age model
175 gives the best possible synchronization of the marine record to the NGRIP ice core
176 record (NGRIP, 2004) making it possible to determine the relative timing of observed
177 changes in the marine proxies compared to temperature changes on Greenland.
178 The final tuned age model presented for MIS3 (Figure 3) only uses the ages given by
179 the tuning to NGRIP and the ash layers. As an independent check, we note that the new
180 age model for the marine sediment core is consistent with calibrated 14C dates (Table 1).
181 These are calculated using the calibration software CALIB 6.0 (Stuiver & Reimer,
182 1993) and an ocean reservoir age of 400 years, applying the “Marine09” calibration
183 curve (Hughen et al., 2004) for ages younger than 25 kyr BP (14C age, kilo years before
184 present) and the “Fairbanks0107” calibration curve for AMS 14C samples of ages older
185 than 25 kyr BP (Fairbanks et al., 2005). Offsets of at least a few thousand years from
186 the “true age” are expected due to the relatively large error in 14C-measurements in the
187 interval 30 - 40 kyr BP, as well as errors in transformation from 14C-ages to calibrated
188 ages. Additional error is introduced by the assumption that sedimentation rates are
189 constant between the dated levels.
8 190 Near surface ocean temperature estimates
191 For the planktonic foraminifera, raw census data of the planktonic assemblages were
192 analyzed and the transfer function technique (Maximum Likelihood - ML) was used to
193 calculate near surface temperature. The ML method operates with a statistical error
194 close to ±1ºC, and is found to give the least autocorrelation compared to other statistical
195 methods (Telford & Birks, 2005).
196 Transforming the δ18O into a measure of sea water isotopic composition
197 The ratio of oxygen-18 to oxygen-16 as they are incorporated into foraminiferal calcite
18 18 198 (δ Ocalcite) is dependent on both seawater δ O (expressed as per mil deviations from
199 Vienna Standard Mean Ocean Water - SMOW) and calcification temperature.
200 Foraminiferal based SST estimates provide an independent estimate of the calcification
201 temperature, and therefore the temperature effects on calcite δ18O. The temperature
202 to δ18O calibration used is based on Kim and O´Neil (1997):
18 18 18 18 2 203 T = 16.1 - 4.64 (δ Ocalcite - δ OSMOW) + 0.09 (δ Ocalcite - δ OSMOW)
204 Knowing the calcification temperature (T) from the foram transfer function and the
18 205 measured δ Ocalcite we can extract the isotopic composition of the ambient sea water
18 18 206 (δ OSMOW). The relationship between salinity and δ OSMOW in the ocean depends on
207 the spatial variation of δ18O associated with the fresh water flux to the ocean, the
208 amount of sea water trapped in ice caps, as well as ocean circulation (Bigg & Rohling,
209 2000).
18 18 210 Before translating glacial δ O data into a measure of δ OSMOW, the effect of global ice
211 volume must be accounted for. Fractionation processes of oxygen isotopes when water
9 212 freezes into ice produce an ocean-wide δ18O shift of about 0.1 ‰ per 10 m of sea level
213 stored in continental ice sheets (Fairbanks, 1989). A review of recent advances in the
214 estimates of sea level history during MIS3 is given by Siddall et al. (2008). The
215 estimated sea level changes during MIS3 vary in amplitude and timing between the
216 different reconstructions. In our study, we use the sea level reconstruction of
217 Waelbroeck et al. (2002), which fits well with available coral based estimates for MIS3,
218 to remove the effects of changing ice volume on the δ18O of ocean water.
219 3. Interstadial and Stadial phases of a D-O cycle
220 Positioned in the path of inflowing warm Atlantic water to the Nordic Seas and close to
221 the routing of freshwater from the Fennoscandian ice sheet, our marine sediment core is
222 ideally suited to better understanding the changes to the hydrography and sea ice cover
223 of the Nordic Seas during MIS3. Here we will describe the warm and cold phases of D-
224 O cycles in the Nordic Seas as suggested by the new sediment core data, before
225 examining the transitions in section 4.
226 The Nordic Seas during Greenland stadials
227 Foraminiferal assemblages used to reconstruct near surface ocean temperature show that
228 water masses at the core site in the Nordic Seas are ~2°C warmer during cold Greenland
229 stadials than during warm interstadials (Figure 3b). Further, there is a warm overshoot
230 at the onset of each interstadial. Despite difficulties in acquiring reliable low-
231 temperature reconstructions in polar and Arctic water masses (e.g., Pflaumann et al.
232 2003), the estimated higher temperatures during stadials compared to interstadials is a
233 robust result. This is based on higher percentages of “Atlantic species” associated with
234 relatively warm conditions (such as Globigerina bulloides, Neogloboquadrina
10 235 pachyderma (dextral) and Turborotalita quinqueloba) during stadials. Interstadials are
236 instead dominated by Neogloboquadrina pachyderma (sinistral).
237 Planktonic foraminifera are known to change their depth habitat in Arctic environments
238 to avoid the low salinity surface, which is separated from the subsurface warm Atlantic
239 layer by a halocline (strong vertical salinity gradient). This results in different depth
240 distributions of foraminifera under different hydrographic conditions (Carstens et al.
241 1997). Therefore a reasonable interpretation of these data is that, during cold stadials in
242 the Nordic Seas, the planktonic assemblages are located below the fresh surface layer
243 and thus record the relatively warm temperatures just below the halocline (Figure 4a).
244 At these depths, the foraminifera are in contact with relatively warm Atlantic water that
245 is isolated from the atmosphere by the halocline and sea ice cover, both of which are
246 features of the stadial phase of the D-O cycle. Note, that at the end of each stadial
247 phase, there is a warm overshoot that provides a clue to understanding the stadial-to-
248 interstadial transition (discussed in section 4).
18 249 Combining stable isotope (δ Ocalcite) measurements on the planktonic foraminifera N.
250 pachyderma sin. with the ocean temperature reconstruction produces an estimate of the
18 18 251 isotopic composition of sea water (δ OSMOW). δ OSMOW is partly related to salinity, but
252 it also reflects spatial shifts in water masses. The isotopic composition of polar water
253 may be different compared to water of subtropical origin, but because the planktonic
254 foraminifera prefer to stay within the warm Atlantic layer, we do not expect that the
18 255 mean δ OSMOW (salinity) at the core site will be substantially different during stadials
256 compared to interstadials. Indeed, we see in Figure 3c that there is no systematic
18 257 difference in the mean δ OSMOW during stadial and interstadial periods. The common
11 258 spikes at the transitions between stadial and interstadial phases and the consistent trends
259 within the stadial and interstadial periods provide insight into the transition mechanisms
260 and are discussed in section 4.
261 The benthic δ18O record measured on Cassidulina teretis (Figure 3d) is remarkably
262 similar to the NGRIP record, exhibiting the characteristic shape and abrupt transitions
263 of D-O cycles. One explanation for the link between the benthic δ18O and NGRIP is
264 provided by Dokken and Jansen (1999), who presented evidence that there are two
265 different forms of deep water production in the Nordic Seas during the last glacial
266 period, depending on the sea ice coverage. As in the modern climate, there is deep open
267 ocean convection in the Nordic Seas during the last glacial period when the surface is
268 free of sea ice. When Greenland is cold (the stadial phase of the D-O cycle), however,
269 they argued that the benthic δ18O becomes lighter in the deep Nordic Seas because there
270 is sea ice formation along the Norwegian continental shelf, which creates dense and
271 istopically light brine water that is subsequently transported down the continental slope
272 and injected into the deeper water. In this scenario, the close link between the benthic
273 δ18O record and NGRIP is due to the deep water being affected by continuous sea ice
274 formation throughout the stadial phase and by open ocean convection during the
275 interstadial phase of the D-O cycle. This interpretation of the benthic record is
276 commensurate with all of the other proxy data in Figure 3, which indicates that during
277 the stadial phase of the D-O cycle, the Nordic Seas was covered with sea ice, the
278 surface waters were fresh (Figure 3c) and at the freezing point. Further evidence for the
279 existence of extensive sea ice stems from Figure 3a and modeling work that links the
280 Greenland δ18O record to sea ice extent in the Nordic seas (see Section 1).
12 281 An alternative interpretation of benthic δ 18O at nearby core sites attributes the changes
282 to bottom water temperature rather than brine (Rasmussen & Thomsen, 2004). The
283 temperature change that would be required to explain the 1-1.5 per mil shifts seen in
284 benthic δ18O at our site implies a 4-6°C warming at a depth of 1500 m during stadials,
285 and almost twice as much during Heinrich events. We believe that enhanced brine
286 rejection at the surface and moderately increased temperatures at depth act together to
287 create the light benthic δ18O values during stadials, and that both processes are
288 necessary for maintaining stadial conditions.
289 The record of Ice Rafted Debris (IRD) indicates that there is less frequent ice rafting
290 during cold stadial periods on Greenland (Figure 3e). In this core, the IRD is of
291 Scandinavian origin, and the record implies that there is less calving from the
292 Fennoscandian ice sheet during the cold stadial, and therefore less meltwater derived
293 from ice bergs. Note that there are two Heinrich events during the time interval shown
294 in Figure 3 (denoted H3 and H4 in Figure 3), both of which occurred during the latter
295 part of a cold stadial phase of a D-O cycle. In this study, we are not concerned with the
296 infrequent, large Heinrich ice rafting events and do not consider Heinrich events to be
297 fundamental to the physics of a D-O cycle. Instead, the reader is referred to the study of
298 Alvarez-Solas et al. (2011) for a discussion of the possible relationship between these
299 two phenomena.
300 In general, the planktonic δ13C exhibits more negative values during stadials (Figure
301 3f), suggesting that the water at the preferred depth for the planktonic foraminifera is
302 isolated from the atmosphere. Thus, benthic and planktonic data collectively suggest
303 that during the stadial phase of the D-O cycle, the eastern Nordic Seas is characterized
13 304 by extensive sea ice cover, a surface fresh layer separated by a halocline from warmer,
305 saltier sub-surface waters, and by deep water production by brine rejection along the
306 Norwegian coast.
307 The Nordic Seas during Greenland interstadials
308 In the interstadial phase of the D-O cycle, when Greenland is warm (Figure 3a), the
309 Nordic Seas are thought to be mostly free of sea ice. Without the presence of sea ice and
310 an associated fresh surface layer, the water column is weakly stratified and the
311 incoming Atlantic water efficiently releases heat to the atmosphere, as is the case in
312 today’s Nordic Seas (Figure 4b). As a consequence, subsurface temperatures in the
313 Nordic Seas are found to be relatively cold (Figure 3b), and any fresh water that may be
314 deposited in the surface layer is easily removed by strong wind mixing, particularly
315 during the winter.
316 As shown by Li et al. (2005), reduced North Atlantic sea ice cover can account for the
317 warm temperatures observed in the Greenland ice cores during interstadials of the last
318 glacial cycle (Figure 1). Due to the exceptional chronology of the high resolution
319 marine sediment record, we can ubiquitously link the timing of warm interstadial
320 conditions on Greenland to periods in the marine record with high benthic δ18O (Figure
321 3d) and well ventilated conditions in the Nordic Seas as evidenced by planktonic δ13C
322 (Figure 3f). The high benthic δ18O values indicate less efficient brine formation
323 combined with lower temperatures, supporting the presence of reduced sea ice cover
324 during interstadials in the Nordic Seas (Dokken & Jansen, 1999). Note that the benthic
325 δ13C record (Figure 3g) is not as straightforward to interpret as the planktonic δ13C
14 326 record: the strong overshoot seen during the transitions between stadial and interstadial
327 conditions is discussed in detail in section 4.
328 4. IMPLICATIONS FOR D-O CYCLES
329 Transition between stadial and interstadial phases
330 The records presented thus far support the idea of a reorganization of the vertical
331 thermohaline structure of the Nordic Seas between the stadial and interstadials phases of
332 D-O cycles. In the interstadial phase, conditions in the Nordic Seas resemble those
333 existing today: the heat transported into the region by the ocean and stored seasonally in
334 the mixed layer is released to the atmosphere, resulting in a moderate, seasonal sea ice
335 cover. In the stadial phase, however, conditions in the Nordic Seas resemble those
336 existing in the Arctic Ocean today: there is a fresh surface layer buffered from the warm
337 Atlantic water below by a halocline. In essence, there is an expansion of sea ice
338 southwards from the Arctic into the Nordic Seas. Indeed, the stratification and vertical
339 structure of the Nordic Seas inferred from our proxy data (Figure 3 & 4) are similar to
340 those of today’s Arctic, where water is several (~3°C) degrees warmer at a depth of
341 400—500 m than at the surface in winter (e.g. Rudels et al., 2004).
342 During the stadial phase, the planktonic foraminfera are mainly recording the
343 temperature of water within, or just below, the halocline (Figure 4a). The sea ice cover
344 prevents wind mixing, and sea ice formation along the Norwegian coast creates dense
345 brines that are rejected from the surface to deeper in the water column, thus helping to
346 maintain the halocline. As the stadial phase progresses, the planktonic foraminfera show
347 an increase in temperature (Figure 3b) consistent with their depth habitat being
348 continuously fed by the northward advection of relatively warm and salty Atlantic
15 349 water. With no possibility of venting heat to the atmosphere above, because of extensive
350 sea ice cover, the warming due to the inflow of warm Atlantic water gradually reduces
351 the density of the subsurface waters, and weakens the stratification that allows the
352 halocline and sea ice cover to exist. The transition to warm interstadial conditions on
353 Greenland occurs when the stratification and halocline weaken to the point of collapse,
354 at which point heat from the subsurface layer is rapidly mixed up to the surface, melting
355 back the sea ice.
356 The removal of sea ice in the Nordic Seas allows efficient exchange of heat between the
357 ocean and the atmosphere, and is seen as an abrupt warming (D-O event) of up to 15°C
358 in Greenland at the start of each interstadial (NGRIP, 2003; Figure 1). There is a distinct
359 overshoot (< 100 years long) in foraminifera-derived subsurface temperatures right at
360 the transition (Figure 3b). The planktonic foraminifera live at the interface between the
361 surface and subsurface layers, and immediately feel the temperature effect as warm
362 subsurface water is mixed up to the surface. Once the fresh surface layer has been
363 mixed away, the ocean is able to vent heat to the atmosphere. The foraminifera now find
364 themselves living in a new habitat, an unstratified surface layer, which is colder than
365 their subsurface stadial habitat.
366 Once in the interstadial (Figure 4b), conditions are comparable to those of today: the
367 Nordic Seas are relatively ice free; inflowing warm Atlantic water is in contact with the
368 surface; and there is an efficient release of heat to the atmosphere. The planktonic-based
369 temperatures show little change (Figure 3b), but there is a consistent reduction in
370 salinity, in particular towards the end of each interstadial (Figure 3c). This freshening of
371 the upper part of the Nordic Seas water column is most likely due to melting and
16 372 calving of the nearby Fennoscandian ice sheet in response to the warm climate
373 conditions, as is supported by the increased frequency of IRD deposits at the site
374 (Figure 3e).
375 Supporting evidence from δ13C
376 The carbon isotope data from planktonic (N. Pachyderma sin.) and benthic (C. teretis)
377 foraminifera at the site (Figure 3f and g) give additional clues about the transition
378 between stadial and interstadial phases in the Nordic Seas. The isotopic signature of
379 carbon (δ13C) preserved in the calcium carbonate shells of foraminfera is related to
380 ventilation and water mass age: at the surface CO2 exchange with the atmosphere and
381 marine photosynthesis preferentially extracts 12C from seawater, causing enrichment of
382 surface water in dissolved inorganic 13C. When the water mass is isolated from the
383 surface mixed layer, its δ13C value decreases with age due to mixing with different
384 water masses and gradual decomposition of low δ13C organic matter. Although the
385 benthic foraminifera C. teretis do not live directly on the surface of the marine
386 sediments, evidence from studies in the Nordic Seas show that the δ13C of C. teretis
387 tracks the δ13C recorded by epibenthic species (benthic fauna living on top of the
388 sediment surface at the seafloor). In the absence of epibenthic foraminifera in high-
389 deposition regions such as the core site, C. teretis can be used with caution to infer past
390 changes in deep water δ13C (Jansen et al., 1989).
391 In the stadial phase, with extensive sea ice, relatively young water of Atlantic origin
392 with high δ13C (Figure 3f) enters the Nordic Seas below the fresh surface layer and
393 halocline. At the onset of the interstadial there is a brief period with extremely light
394 benthic δ13C (Figure 3f). This is consistent with a mixing in of old, poorly ventilated
17 395 deep water masses of Arctic origin from below the Atlantic layer in the Nordic Seas
396 (e.g. Thornally et al., 2011). Enhanced mixing at the transition will also bring well-
397 ventilated waters with high δ13C down from the surface; however, this signal is
398 overwhelmed by the mixing in of the old water from below with extremely light δ13C.
399 Early in the interstadial phase, following the abrupt transition, the core site is
400 dominated by high δ13C as seen in both the benthic (Figure 3f) and planktonic (Figure
401 3e) records. At this point, the Nordic Seas are free of sea ice, well ventilated, and the
402 stratification is weak due to efficient winter mixing and absence of a halocline.
403 Continuous input of terrestrial freshwater from melting ice sheets, such as the
404 Fennoscandian, also contribute to the high planktonic δ13C. However, towards the end
405 of the interstadial the influence of sea ice is increasing in the Nordic Seas (Figure 3d)
406 and the stratification is increasing, as evidenced by decreasing surface salinity (Figure
407 3b). As a consequence, benthic δ13C is reduced (Figure 3f) indicating increased
408 stratification and less ventilation of waters at the depth of the core site.
409 Returning to the stadial phase (Figure 5I), with an extensive sea ice cover and return of
410 the halocline, the surface and intermediate waters are now isolated from the atmosphere,
411 as evidenced by relatively low planktonic δ13C values (Figure 3e). Although the
412 qualitative features we have described are robust for each D-O cycle, there are
413 quantitative differences between each cycle. For example: convection may reach deeper
414 during some interstadials than others, making the light spike more pronounced; or the
415 ocean may re-stratify more quickly during some interstadials, resulting in a more
416 immediate recovery to heavier δ13C values. Such differences will also be reflected in
417 the other proxy records.
18 418 5. DISCUSSION
419 The conceptual model for D-O cycles presented in the previous section, and
420 summarized in figures 5 and 6, is based on the proxy records measured in MD992284,
421 and specifically, on the relationship between signals in these marine records and signals
422 in the Greenland oxygen isotope record. The marine records themselves provide direct
423 information on the water column in the Atlantic inflow region of the Nordic Seas.
424 However, the expected changes in sea ice cover illustrated in the conceptual model have
425 implications for the response and role of other parts of the climate system in D-O
426 cycles. We have focused on the Greenland oxygen isotope record as a reference time
427 series for D-O cycles, but additional information is provided by the Greenland
428 deuterium excess record (Jouzel et al., 2005; Masson-Delmotte et al., 2005; Thomas et
429 al., 2009). Deuterium excess shows a rapid decrease in the temperature of source waters
430 for Greenland precipitation at stadial to interstadial transitions. Assuming that the
431 abrupt transition from stadial to interstadial conditions is mediated by the removal of
432 sea ice in the Nordic Seas, this would provide a new, local source of moisture for
433 precipitation on the Greenland ice sheet during warm interstadials. The ocean
434 temperature of this source would be significantly lower than sources further to the south
435 (Sachs & Lehman, 1999), thus explaining the observed 3°C rapid decrease in the source
436 temperature for precipitation on Greenland during phases of rapid increase in surface
437 air temperature of D-O cycles (Steffensen et al., 2008).
438 On a larger scale, changes in sea ice cover are expected to be associated with a
439 reorganization of atmospheric circulation, as is suggested by changes in dust delivered
440 to Greenland (Mayewski et al., 1997). We hypothesize that during interstadials, the
19 441 storm track extends northeastwards into the Nordic Seas, transporting heat into the
442 region and inhibiting formation of sea ice by mechanical mixing. During stadials, the
443 storm track is more zonal and positioned south of the Fennoscandian ice sheet, allowing
444 for the quiescent conditions required to maintain strong katabatic winds and sea ice
445 production in leads and polynyas on the continental shelf. Shifts in the orientation of the
446 Atlantic storm track could help explain the dust record in Greenland, as well as far field
447 monsoon and precipitation changes associated with D-O cycles (Pausata et al., 2011).
448 Many previous models of D-O cycles require, or assume, a flux of freshwater to the
449 North Atlantic mainly from the Laurentide ice sheet in order to maintain extensive sea
450 ice cover over the Nordic Seas during stadials. In contrast, the provenance of IRD in
451 MD992284 indicates that the freshwater forcing originates from the Fennoscandian ice
452 sheet, and is associated with interstadials as much as, if not more than, stadials. This
453 provides a local freshwater source in the path of the Atlantic inflow in a highly sensitive
454 area of the Nordic Seas, which can aid more directly in creating a halocline and
455 facilitating expansion of sea ice compared to freshwater sources on the western rim of
456 the North Atlantic.
457 We further expect that the presence of a halocline and extensive sea ice during stadials
458 greatly reduces the ventilation of the deep ocean as it prevents contact between
459 inflowing Atlantic water and the atmosphere in the Nordic Seas. This will impact the
460 properties of the return flow of Atlantic water leaving the Nordic Seas, and should be
461 recorded by intermediate and deep water proxies throughout MIS3. The marine proxy
462 data presented in this study support the existence of an active circulation of relatively
463 warm Atlantic water into the Nordic Seas during stadials, but do not give a direct
20 464 measure of the rate of Atlantic meridional overturning circulation. Evaluating changes
465 in the AMOC as a response to enhanced stratification in the Nordic Seas during stadials
466 is beyond the scope of this study. However, such changes are expected, as the returning
467 outflow of Atlantic water from the Nordic Seas would not be sufficiently dense to sink
468 into the deep abyss during stadials. As discussed in several studies (Crowley, 1992;
469 Stocker & Johnsen, 2003), a reduction in the AMOC could lead to warming of the
470 Southern Ocean and possibly explain the “seesaw” pattern with warm temperatures on
471 Antarctica when Greenland is cold (EPICA , 2006).
472 MIS3 records from other Nordic Seas marine cores share many consistent features with
473 the records presented in this study. However, different interpretations of these features
474 lead to hypothesized mechanisms for D-O cycles that are different from ours in
475 important ways.
476 For example, Rasmussen & Thomsen (2004) suggest that there is warming at depth
477 during stadials, based on the depleted benthic δ18O from a sediment core obtained from
478 the central Nordic Seas basin at 1000 m depth. We also observe depleted benthic δ18O
479 during stadials at our core site, but we interpret this primarily to be due to brine
480 production. Given the position of our core on the Norwegian slope, interpreting the
481 depleted benthic δ18O as a temperature signal would imply temperatures 4-6°C warmer
482 during stadial periods at a depth of 1500 m in the Nordic Seas, which we find unlikely.
483 Another example, Petersen et al. (2013), suggest that the duration of the interstadial and
484 stadial phases of the D-O cycles are related to ice shelves, while sea ice changes provide
485 the abrupt transitions. Critical support for this ice shelf mechanism is an increase in ice-
486 rafting during each D-O stadial inferred from IRD and fresh surface anomalies (based
21 487 on planktonic δ18O) in the Nordic Seas and Irminger Basin (Dokken and Jansen 1990,
488 van Kreveld et al. 2000, Elliot et al. 1998, 2001). However, the postulated occurrence of
489 ice rafting during cold Greenland stadials is inconsistent with the chronology of our
490 marine sediment core. The occurrence of ice rafting at our site in the Nordic Seas is
491 enhanced during the latter part of the interstadials (Figure 3e), although our
492 hypothesized reorganization of the Nordic Seas circulation during D-O cycles is not
493 dependent on the input of meltwater from melting ice sheets.
494 Theses issues show that open questions remain, requiring better spatial coverage of
495 marine sediment cores with adequate chronology to pin down the correct timing of
496 reconstructed oceanic signals compared to the ice core records.
497 Our conceptual model of D-O cycles requires the presence of a shallow ocean over a
498 continental shelf adjacent to the Fennoscandian ice sheet. Should the Fennoscandian ice
499 sheet extend over the continental shelf, the stadial phase of the proposed D-O cycle
500 could not exist because brine production through sea ice formation would not be
501 possible and thus a stable halocline could not be maintained in the Nordic Seas. Hence,
502 consistent with observations, our conceptual model precludes the possibility of D-O
503 cycles at the Last Glacial Maximum, when the Fennoscandian ice margin was located
504 near the shelf break. Furthermore, with less efficient brine production at LGM, we
505 expect sea ice formation will be reduced compared to a typical stadial, resulting in ice-
506 free conditions over large parts of the Nordic Seas (Hebbeln et al., 1994).
507 Finally, we can draw an analogy between the stadial phase in the Nordic Seas during the
508 last glacial period and the Arctic Ocean today. Both feature a fresh surface layer,
509 halocline, and extensive sea ice. The existence of these features is sensitive to the heat
22 510 transported by inflowing Atlantic water and to brine release during sea ice formation on
511 the continental shelves. Projections of global warming (Holland & Bitz, 2003) are
512 consistent with observed trends (Smedsrud et al., 2008) showing increased heat
513 transport to the Arctic Ocean by the Atlantic inflow. Together with reduced sea ice
514 formation this will weaken the halocline, and could rapidly tip the Arctic Ocean into a
515 perennially ice-free phase. In such a phase, warm Atlantic water will be brought to the
516 surface and have a severe impact on Arctic climate and the stability of the Greenland ice
517 sheet.
518
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731
32 Mean 14C Age, Error, Calibrated ages Max. 1σ Min. 1σ Laboratory depth uncorrected 1σ - intercept Reference and ash cal. age cal. age identity (cm) (yr BP) (yr) (kyr) (yr BP) (yr BP)
TUa-3310 1100,5 21975 160 24,64 25018 24270
Fugloyarbanki ash 1407,5 26, 69 (b2k) (FMAZ II) 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 3040,5 38,07 (b2k) Zone III (FMAZ III) POZ-29524 3075,5 34600 700 39,53 38830 40230
732
733 Table 1: Radiocarbon dated intervals used as a first approximation to create the age
734 model. All AMS 14C ages are measured on N. pachyderma sin. Also shown in the table
735 are two identified ash layers. The ages given for Fugloyarbanki Tephra (FMAZII) and
736 Faroe Marine Ash Zone III (FMAZIII) are from Svensson et al. (2008). The ash ages
737 are referred to in b2k (before year A.D. 2000). Sample with laboratory reference “TUa-
738 “ are measured at the Uppsala Accelerator in Sweden, but the samples have been
739 prepared at the radiocarbon laboratory in Trondheim, Norway. Samples labelled “POZ-
740 “ are prepared and measured in the radiocarbon laboratory in Poznan, Poland.
741
33 742 Figure 1: Surface air temperature response during the extended winter season
743 (December-April) to changes in Nordic Seas ice cover (Figure redrawn from
744 simulations in Li et al., 2010). Black contours indicate maximum (March) ice extent in
745 the two scenarios. The white dot is the location of the core.
746 Figure 2: (a) δ18O of NGRIP plotted versus depth. (b) Measured low-field magnetic
747 susceptibility, Anhysteretic Remanent Magnetization (ARM) (log scale/unit 10-6A/m) in
748 MD992284 plotted versus depth. Red stars indidate the stratigraphical level (depths)
749 from where AMS 14C ages are measured (see also Table 1). (c) NGRIP δ18O and ARM-
750 record from MD992284 after converting the MD992284 age to the NGRIP age (b2k).
751 Green vertical lines indicate tuning points for the (numbered) onset of interstadials in
752 the ice core- and marine records. Red vertical lines connect identical ash layers
753 (FMAZIII and FMAZII) in NGRIP and MD992284 as described in the text and in Table
754 1.
755 Figure 3: Down core data sets from NGRIP (a) and MD992284 (b-g) covering the
756 period 41ka to 31ka plotted on the GICC05 (b2k) time scale. a) NGRIP δ18O (proxy for
757 Greenland temperature), b) reconstructed temperature based on planktonic foram
18 18 758 assemblages, c) near surface δ OSMOW, d) benthic δ O, e) flux corrected Ice Rafted
759 Debris (IRD), f) near surface δ13C, and g) benthic δ13C. Interstadial periods are
18 760 indicated by grey shading. The near-surface δ OSMOW is determined from the Kim and
761 O’Neil (1997) calibration curve using the δ18O of N. pachyderma sin. and the
762 temperature obtained using transfer functions (see section 2). The near surface δ13C is
763 measured on N. pachyderma sin. and the benthic δ18O and δ13C are measured on
764 Cassidulina teretis
34 765 Figure 4: A schematic column model of the stadial (left) and interstadial (right) phases
766 consistent with the Nordic Sea sediment core MD992284.
767 Figure 5: A schematic timeline showing the evolution of a D-O cycle that is consistent
768 with the Nordic Sea sediment core MD992284 and the NGIP ice core.
769 Figure 6: Schematic showing wintertime conditions in the North Atlantic and Nordic
770 Seas during typical cold stadial periods (left column) and warm interstadial periods
771 (right column) of a D-O cycle. Top panels show maps of the North Atlantic region with
772 sea ice extent and land ice. The location of the sediment core is marked with a blue dot
773 and the sections described in the lower panels are indicated. Middle panels A) and C)
774 show north-south sections of the North Atlantic during stadial and interstadial
775 conditions, respectively. Bottom panels B) and D) show east-west sections of the
776 Norwegian Sea during stadial and interstadial conditions, respectively.
35 30
25
20
15
10
5
0 N-GRIP (GICC05 age) 26 28 30 32 34 36 38 40 -34
-36 3 4 5 6 7 8 9 10 -38 Mono Lake Laschamp event -40 O NGRIP tephra (FMAZ II) tephra 18 -42
-44
-46 (FMAZ III) tephra
3 4 5 6 7 8 9 10 400
300
200 ARM (*1000)
100
1400 1600 1800 2000 2200 2400 2600 2800 3000 3200 Core depth (cm)
-36 3 4 5 6 7 8 9 10 -38 400
-40 (FMAZ II) tephra 300 O NGRIP
18 -42 200 ARM (*1000)
-44 100 Mono Lake tephra (FMAZ III) tephra Lachamp event
-46 26 28 30 32 34 36 38 40 42 N-GRIP (GICC05 age - ka) 31 32 33 34 35 36 37 38 39 40 41 42 -36 5 6 7 8 9 10 Interstadial numbers H3 H4 -38 O 18
-40 a) -42 10 NGRIP C)
-44 8 O
-46 6 b) 4
sin. 0,5 2 Reconstructed SST ( -0 c) 0 -0,5 -1,0 N.pachyderma 6,0 -1,5 5,5 -2,0 O (smow) O (smow) d) 5,0 (benthic)
18 -2,5 4,5
15 teretis C.
4,0 O 18 e) 3,5 10
5 IRD (No. grain>0,5mm/g IRD (No. 0,5
sin. f) 0,3
0,1
-0,4 N.pachyderma -0,1 C 13 -0,8 -0,3 g) (benthic) -1,2
-1,4 teretis C. C 13 -1,8 31 32 33 34 35 36 37 38 39 40 41 42 Age (Ka) (GICC5) 100 m Fresh surface Halocline 300 m Glacial Atlantic inflow Atlantic inflow
600 m
fossil/aged/leftover (intermediate and Nordic Sea water deep/ventilated) Nordic Sea water
Brine increasing perennial sea ice open ocean winter ice perennial sea ice open ocean
Greenland temperature
Brine production / Sea ice cover
Subsurface salinity
time Figure 3
Stadial conditions Interstadial conditions
A C
B D
A C
Heat loss Sea ice Cold, fresh water
500 meters 500 meters Northward heat transport
Nordic Seas Nordic Seas Greenland - Greenland Scotland ridge - Greenland Scotland ridge
Equator Arctic Equator Arctic
B Cold air freezes seawater D Heat loss Sea ice Newly formed ice
Cold, fresh water Brine release
Cold, saltier water Halocline Northward heat transport Denser, salty 200 meters water sinks
Warm, salty Atlantic water
Norwegian sea Norway Norwegian sea Norway