6.2. Ice Sheet–Upper Ocean Connection
 Today N. pachyderma sin. thrives in the Nordic Sea and Arctic Ocean at the base of the summer seasonal surface mixed layer at ∼50 m [Bauch et al., 1997; Jonkers et al., 2010]. In the subpolar North Atlantic, G. bulloides calcifies within 0–60 m water depth, blooming in the summer months [Cortijo et al., 1997; Ganssen and Kroon, 2000]. Field and modeling evidence from present-day and LGM assemblages in polar and subpolar regions [Fraile et al., 2009a, 2009b; Jonkers et al., 2010; Kohfeld et al., 1996] suggest that glacial blooms of both G. bulloides and of N. pachyderma sin. occurred during the summer months. However, given the habitat preferences of the two species we can interpret the G. bulloides δ18O record as being a summer SST signal and N pachyderma sin. δ18O as representing mean annual subsurface conditions [Peck et al., 2008]. The ∼0.5‰ offset between both δ18O records at this site prior to the LGM (Figure 9b) points to a strong gradient between surface and subsurface waters during the summer months, with denser (cooler and/or saltier) waters below lighter (warmer/less saline) surface waters. At the LGM the two δ18O records converge suggesting that both species are living in the same water mass, suggestive of year-round mixing of the upper waters/no seasonal thermocline.
 A comparison of both isotopic records with the equivalent from core MD01-2461 in the Porcupine Seabight (SW Ireland) [Peck et al., 2006] (Figure 8) reveals clear similarities in upper water column structure at the two sites. The surface isotopic signals seem to represent roughly equivalent water masses at both locations. Since G. bulloides from the LGM seem to calcify in summer in both areas [Fraile et al., 2009b], the small occasional difference between both records can be interpreted as a minor latitudinal temperature gradient and/or meltwater lenses.
 Comparable N. pachyderma sin. δ18O records from the two sites suggest similar subsurface water mass properties at both locations up until ∼27 ka B.P. After this time, the δ18O N. pachyderma sin. record at the more southerly site becomes up to 0.5‰ lighter than in Rosemary Bank, evidencing clear differences between both records.
 A likely explanation for the different behavior of both records from ∼27 ka B.P. relies in the migration of the Polar Front. The comparison of N. pachyderma sin. relative abundance records (Figure 8) during MIS 3 shows near synchronous variations between both sites until ∼27 ka B.P., allowing us to infer an unstable Polar Front migrating rapidly along the western BIIS margin, across both sites, simultaneously to GS-GI cyclicity. Additional evidence of the migration of the Polar Front as far south as 40°N during HE and some of the GS is observed in records from the western Iberian Margin [Eynaud et al., 2009; Salgueiro et al., 2010; Voelker et al., 2009]. In contrast, at the onset of MIS 2 the persistent dominance of N. pachyderma sin. in the faunal assemblage and heavier δ18O values at Rosemary Bank are consistent with a scenario of BIIS build-up and a likely almost stable position of the Polar Front south of core MD04-2829CQ. Periodic migration of the Polar Front south of MD01-2461 is reflected in our records by the convergence of both isotopic and faunal signals, at 25.7 ka B.P., 25 ka B.P. and HE 2. After HE 2 the Polar Front remains between the two sites but appears to periodically retreat north of core MD04-2829CQ when N. pachyderma sin. relative abundance from this site decreases to converge with MD01-2461, although this is not expressed in the isotopic records. BIIS advance at ∼21.5 ka B.P. (also observed by Knutz et al.  and Peck et al. ) is associated with a final advance of the Polar Front toward MD01-2461, while three further Polar Front retreats northward of Rosemary Bank are noted at 20.1, 19.9 and 19 ka B.P.
 Cold events such as the ones at ∼25.7 and ∼21.5 ka B.P. produced meltwater plumes both north and south of the BIIS, as reflected in the light δ18O peaks in G. bulloides recorded at both sites. At HE 2, the amount of melt waters and icebergs seems to be much larger at the southern front of the BIIS, as shown by the large isotopic spike and IRD flux recorded in core MD01-2461 (Figure 8) [Peck et al., 2007a] and in core MD95-2002 further southeast along the continental margin [Grousset et al., 2000; Auffret et al., 2002].
 Contrary to the N. pachyderma sin. record, the small response of G. bulloides to the events recorded during HE 2 and MIS 2 in core MD04-2829CQ likely show a partially biased record affected by the harsh conditions of average years, with low to very low abundances (Figure 4) and/or summer blooming restricted to the warmer years.
 Isotopic data from core MD04-2829CQ suggest that MIS 3 GS conditions were characterized by cooling synchronously to iceberg calving from the BIIS (Figure 9b). The frequency and scale of iceberg discharge events and hence the input of low-salinity waters to the surface layers water are greatly enhanced after ∼31 ka B.P. as may be inferred from the increased magnitude of the light G. bulloidesδ18O excursions and of simultaneous maxima in IRD flux during the centennial-scale BIIS-sourced events prior to HE 2 and during the LGM. The absence or significantly smaller scale of the related decreases in the N. pachyderma sin. δ18O record during most of these cold intervals is remarkable since the associated IRD flux peaks are of similar or even greater magnitude to those recorded during HE 3 and 4. This may reflect a stronger water column stratification resulting in the migration of N. pachyderma sin. to deeper depths below the halocline [Kohfeld et al., 1996]. Alternatively, Elliot et al.  attributed an analogous N. pachyderma sin. δ18O signal to a combination of decreased meltwater flux and its origin in coastal ice sheets, therefore showing a less fractionated isotopic signature. However, since the BIIS is in all cases the likely source of meltwater and the ice sheet size was far larger after ∼28 ka B.P. than during MIS 3, the degree of distance-based isotopic fractionation between ice accumulation areas and the ice sheet margins should have increased rather than decreased as the BIIS increased its ice volume, and for the interval before ∼28 ka B.P. the relatively low fractionation may have reduced the meltwater signal recorded in the N. pachyderma sin. δ18O. The magnitude of the G. bulloides δ18O light anomalies and IRD fluxes between 28 and 18 ka B.P. in MD04-2829CQ suggest that a meltwater lens may have been a persistent feature at that time/site, although, because of the low abundances of G. bulloides, this interpretation should be treated with some caution. Overall, it seems probable that the freshening associated with the BIIS IRD surging only influenced the surface waters and did not fully affect deeper layers. The scale of some of these calving events (such as those at ∼25.7 ka B.P. and ∼21.5 ka B.P.) must have involved much of the western margin of the BIIS since simultaneous meltwater plumes were also recorded in core MD01-2461 (Figure 8). Larger fluxes of IRD at MD01-2461 during these two events [Peck et al., 2007a] compared with our core further north may be indicative of higher instability in the southwestern BIIS margin at those times.
 As highlighted in Figure 9 it is evident that in our core HE 4 and 2 are characterized by prominent light peaks in the isotopic record of N. pachyderma sin. coincident with the arrival of LIS-sourced material. These excursions are not accompanied by large depletions of the isotopic values of G. bulloides as during centennial-scale cooling events; rather, negative values in G. bulloides δ18O are less clearly defined, and tend to lag those of N. pachyderma sin. It has been suggested that convergence of surface and subsurface δ18O values during HE represents strong mixing of the upper water column down to ∼100–200 m [Simstich et al., 2003], and a recent study suggest that this deepening of the surface mixed layer might have been caused by a brief increase in storminess during GSs containing HEs [Rashid and Boyle, 2007]. However, convergence of our isotopic records is not evident in HE 4 and HE 2 and only HE 3 shows possible evidence of this process although leading the maximum IRD flux by a few hundreds of years.
 HE 3 stands out in our record as a distinct event due to the absence of LIS-sourced material and to the particular isotopic signals recorded in this interval. Before this event and for a ∼500 year interval, δ18O values from three different species representing diverse depths in the upper water masses decrease simultaneously (Figure 9b). A prominent negative anomaly is also observed in the N. pachyderma sin. δ18O record at MD01-2461 around the same time (Figure 8), with an associated increase in benthic δ18O and decrease in benthic δ13C values and 14C marine reservoir ages [Peck et al., 2006, 2007b]. We note that these anomalies coincide with extremely low IRD fluxes at MD04-2829CQ (Figure 9f), consisting predominantly of mica flakes, and of a faunal assemblage dominated almost entirely by T. quinqueloba in the planktonic foraminiferal fraction (Figure 4). Together, all these indicators seem to show either the presence of an exceptionally warm or less saline water mass apparently unconnected to BIIS calving events, since its IRD composition is unique. This event can be correlated to the so-called “DO 4.1” in Greenland and Antarctica ice core records [EPICA Community Members, 2006] and represents a warming phase in those records and in the Western Mediterranean [Sierro et al., 2009]. Warming is not obvious from our records as T. quinqueloba is characterized as an extremely resistant species that can survive to very low salinity conditions [Simstich et al., 2003], and has been previously interpreted as a marker for the position of the Arctic front [Johannessen et al., 1994; Eynaud et al., 2009]. Moreover, simultaneously to our event Lekens et al.  registered a “low salinity” event focused in the southern Nordic Seas apparently caused by the breakage of ice dams and flooding of freshwater glacial lakes from the North Sea. These authors consider this low δ18O as representing HE 3 in this area, but our records indicate that, if being the same event, it precedes HE 3 itself (or at least BIIS IRD response to HE 3) by 2–3 centuries, in line with the observations by van Kreveld et al.  in the Irminger Sea.