3.1. General trends

Figure 1a shows the simulated ocean temperature at 250 m depth, time-averaged for the entire 1992–2009 period. The warmest, saltiest waters are found to the east of the Reykjanes Ridge. These are subtropical-origin waters transported northward by the North Atlantic Current, with temperatures ≥8˚C at this depth. These waters are transported west by the Irminger Current (IC, Fig. 1b) towards southeast Greenland and north towards Svalbard and northeast Greenland by the Norwegian Atlantic and West Spitsburgen Currents. At the shelf along southeast Greenland they combine with the cold (1˚C), low-salinity shelf waters of the EGC coming from the Arctic and Nordic Seas (Reference Sutherland and PickartSutherland and Pickart, 2008) to form a two-component current with cold fresh waters above the shelf and warm salty waters along the slope. The EGC rounds the southern tip of Greenland at Cape Farewell to become the northward-flowing WGC, with subsurface ocean temperatures 2–3˚C lower than the IC, due to air–sea heat losses and lateral mixing. IW on the WGC is traceable in Figure 1 as the warm and fast-moving water flowing northward along the slope at 250 m. A portion of the WGC bifurcates west into the Labrador Sea and joins the Labrador Current (LC) to complete the western branch of the SPG (Fig. 1b). The non-bifurcating northward-flowing branch of the WGC passes through eastern Davis Strait into Baffin Bay (Reference BuchBuch, 2002). In northern Baffin Bay, the now significantly cooler IW on the WGC (1˚C) meets colder (≤0˚C), fresher southward-flowing Arctic currents from Nares Strait, and from Jones and Lancaster Sound. Together, these waters flow southward along Baffin Island in the Baffin Island Current (BIC) through western Davis Strait to the LC. In northeast Greenland, the subtropical-origin waters joining the EGC are much cooler (3.5–5˚C) than those waters joining farther south, a result of cooling during its northward transit.

The 4km ECCO2 solution reveals major changes in subsurface ocean temperature around Greenland between 1992 and 2009 (Fig. 1c and d). As the NAO ceased its sustained positive phase, IW in the SPG warmed and propagated along the EGC/WGC, yielding positive subsurface temperature anomalies in southeast, southwest and northwest Greenland. The warming around north and northeast Greenland took considerably longer, a result of the longer advective timescales associated with the propagation of subtropical-origin waters into the Greenland Sea and Arctic Ocean (Fig. 1c). During the second half of the simulation period the model indicates widespread warming in Baffin Bay and the Greenland Sea, and a cooling in the eastern Irminger Sea and southeast Greenland (Fig. 1d).

3.2. Comparison with in situ data

We compare the 4 km ECCO2 solution with in situ observations in the central Labrador Sea (within 150km of the AR7W line, in depths exceeding 3250 m), the western Irminger Sea (within 150km of the AR7E line, in depths exceeding 2500 m and west of 36˚ W) and in eastern Baffin Bay (71–74˚ N; east of 60˚W and in waters shallower than 550 m) (Fig. 2). Temperature data are drawn from conductivity-, temperature- and depth-measuring devices (CTDs), autonomous profiling floats and expendable bathythermographs compiled from Hydrobase 2 (R. Curry, unpublished information), the Global Temperature and Salinity Profile Program of the US National Oceanographic Data Center Operational Oceanography Group (2006) and the ICES (International Council for the Exploration of the Sea) Oceanographic Database and Services (2012).

Fig. 2.

Comparison of subsurface temperature along the AR7W line (see dotted black box in Fig. 1a) in the central Labrador Sea from (a) in situ data and (b) ECCO2 4 km Arctic Ocean solution for the time period 1992–2009 minus an absolute offset of 0.4˚C. Similar plots in (c) and (d) along the AR7E line in the Irminger Sea (Fig. 1a), and (e) and (f) for the Eastern Baffin Bay evaluation box (Fig. 1a). Triangles in (a), (c) and (e) denote the availability of observations. Temperature is linearly interpolated in between except in (e) between 1991 and 1996 where the data gap is too large.

The 4 km ECCO2 simulation is initialized from the WOA05, a temperature and salinity atlas, which, by coincidence, includes several important observed features of the SPG in 1992 including a relatively cool and weakly stratified interior. Compared to observations in the Labrador and Irminger Seas between 1992 and 1996, the WOA05 has approximately +0.4˚C bias. This constant bias has been removed from the model results in Figures 2 and 3 in order to focus the discussion on temporal trends.

Fig. 3.

Regional changes in subsurface ocean temperature from the 4 km ECCO2 simulation for the time period 1992–2009 minus an absolute offset of 0.4˚C. Left column is for off-shelf sites (triangles in Fig. 1a); right column is for on-shelf sites (diamonds in Fig. 1a). The size of the three-dimensional boxes used to calculate annual-mean subsurface temperatures is listed in Table 2. Left vertical axis indicates anomalies in subsurface temperature in reference to the average over the entire time period, between –1.5 and +1.5˚C, with vertical bars colored in red for positive anomalies and blue for negative anomalies. Right vertical axis in black is absolute subsurface temperature, with a scale that varies from site to site but not between on-shelf and off-shelf, except for K_O and K_S. Absolute temperature and anomaly calculations are insensitive to the particular trough or off-shelf location near each region (±0.2˚C) but are somewhat sensitive to the depth range (shelf) and offshore extent (off shelf) over which the calculation is made. Hence, both on- and off-shelf averaging volumes use the same depth range, chosen to be below the seasonal mixed layer and above on-shelf waters without significant temporal variability. The off-shelf volumes span the IW component of the boundary current where possible, or comparably sized regions off the shelf where not (P_S).

Waters within the SPG have significantly warmed since the mid-1990s owing to a shift to generally milder wintertime atmospheric conditions. With reduced wintertime air–sea heat losses, net lateral advective heat flux convergence into the subpolar gyre has dominated over atmospheric heat loss. The model closely reproduces the magnitude and timing of the observed interior warming of the Labrador and Irminger Seas: about +1.0˚C in the upper 500m between 1992 and 2004 and +0.6˚C between 500 and 2000 m. In the Labrador Sea, temperatures in the upper 500 m cease their dramatic warming by 2004 and thereafter exhibit interannual variability consistent with known atmospheric variability (e.g. the cooling/deep-convection event of 2008; Reference VageVage and others, 2009).

We expect the simulated SPG warming to be limited for several reasons. Reference Lohmann, Drange and BentsenLohmann and others (2009) showed that the response of the SPG to a shift from NAO+ to NAO– conditions is sensitive to the amount of preconditioning by the atmosphere. The model initial conditions are identical to the WOA05 and are therefore not preconditioned by the NAO+ forcing, which began in the late 1980s. In addition, because the lateral ocean open boundaries are taken from WOA05, the waters feeding the SPG do not exhibit realistic interannual hydrographic variability. Finally, model–data differences include errors in the JRA-25/JCDAS surface boundary conditions and inaccuracies in the model subgrid-scale parameterizations.

Comprehensive model–data comparison is considerably more difficult for the eastern Baffin Bay region than for the SPG owing to the paucity of in situ observations over the simulation period, especially during the first 4 years when no measurements are available. Assuming that the ocean state in 1992 is similar to measurements taken in 1988 and 1990, the data indicate a warming of 1.0˚C between 100 and 300m between 1992 and 1999. At these depths, the model is initially ~1.0˚C cooler than observations, possibly an artifact of the sparse data used to generate the WOA05. However, the model warms to temperatures comparable to initial conditions by 1998 and then to contemporaneous observations by 1999–2000. A dip in 2005 temperatures is reproduced, as well as the subsequent 0.5˚C warming. Simulated temperatures in this region for the 2000s are in broad agreement with narwhal-based observations reported by Reference Laidre, Heide-Jørgensen, Ermold and SteeleLaidre and others (2010) who found maximum wintertime temperatures up to 3.4˚C at 350 m in the same region.

In terms of total temperature change, the model exhibits an extra degree of warming in the eastern Baffin Bay region compared with that observed. Without additional experiments it is impossible to determine how much of the warming is due to a realistic model response to the specified atmospheric forcing, as opposed to the model evolving away from imperfect initial conditions or to systematic model errors. It also must be noted that the model analyses use all gridpoints and times, whereas the observation analyses rely on temporally and spatially sparse measurements. The impact of sample density on the observation analyses can be seen by comparing the temperature reconstructions before and after the widespread arrival of Argo floats in 2002.

3.3. Modeled regional changes

We examine changes in subsurface temperature at seven key locations, spanning different oceanic regimes (Fig. 3). The survey starts at Kangerdlugssuaq Gletscher on the east coast, closest to the source of IW, and continues clockwise around Greenland. We analyze both off-shelf (triangles in Fig. 1a) and on-shelf (diamonds) simulated subsurface ocean temperature data, averaged over 1 year periods and over the range of depths indicated in Table 2. Anomalies in subsurface temperature are calculated in reference to the average temperature for the entire period. The comparison of off-shelf versus on-shelf temperature documents whether changes in ocean temperature are transmitted across the shelf break. Off-shelf locations are selected at the edge of the continental shelf, near major troughs in the sea-floor; on-shelf locations are selected in front of major glacial fjords, within troughs. Troughs are important because they provide a natural pathway for subsurface warm waters to intrude onto the continental shelf and reach the glacial fjords.

Table 2.

Location of three-dimensional boxes used to plot average annual subsurface temperature in Figure 3. ‘Coordinates’ refers to the center of the volume used in the analysis while ‘Area’ and ‘Depth’ refer to the horizontal area and depth range of the averaging volumes, respectively. ‘Mean shelf depth’ indicates the mean bathymetric depth on the shelf for the shelf locations while ‘off-shelf depth range’ indicates the range of bathymetric depths beneath the off-shelf averaging locations. The off-shelf averaging regions generally span an area from shelf-break to deeper waters

At off-shelf location K_O (Fig. 1a), the initial ocean temperatures have the warmest (5.8˚C) IW of the seven selected locations. K_O is far offshore from K_S due to a broad continental shelf in that area. From 1994 to 2004, the modeled subsurface ocean temperature increased by 1.3˚C at K_O, and it did not significantly change thereafter. On the shelf, the temperature at K_S is initially 3.1 ˚C colder than off the shelf (2.7˚C at K_Svs 5.8˚C at K_O), and the warming is 1.2˚C from 1994 to 2004, with little warming in later years. At SE_O in the southeast, ocean temperature changes are synchronous with K_O, with a warming of 1.2˚C between 1994 and 2004. In contrast to K_O and K_S, both on- and off-shelf areas decline by 0.5˚C after 2005, yielding an overall warming for the entire period of ~1˚C.

On the west coast at N_O, simulated subsurface temperature changes are +1.8˚C in 1994–2004 and –0.6˚C after 2006. At N_S, both the warming (2˚C) and cooling (1.0˚C) are larger than at N_O. Larger temperature changes are modeled near the mouth of Disko Bay, with +2.5˚C between 1993 and 2003 at J_O. After 2004, J_S continues to warm by another 0.5˚C, but not J_O. Ocean waters at J_O are initially 4˚C colder than at K_O.

Farther north, in Melville Bay, modeled ocean waters at M_O are 2.6˚C colder than in Disko Bay at J_O (0.4˚C versus 3˚C for the entire period) and warmed by 3˚C in 1994–2009 for both the off- and on-shelf locations. The most significant warming takes place after 1997 and after 2004. This delayed warming compared to the south is expected from the northward spreading of warm waters. In north and northeast Greenland, where subsurface temperature anomalies are not directly linked to the SPG via the WGC pathway, warming is delayed until after 2004. The modeled change at 79N_S (+1˚C) is less than at 79N_O (+1.5˚C), indicating that only a fraction of the available ocean heat is transmitted up to the shelf at this location.

3.4. Modeled regional trends versus observations

The 1.6˚C modeled warming at K_S between 1995 and 1998 is consistent with the 1.8˚C warming reported by Reference ChristoffersenChristoffersen and others (2011). On the southwest coast at N_S and N_O, the 1.5–2˚C modeled warming for 1995–99 is consistent with the observations reported by Reference Myers, Kulan and RibergaardMyers and others (2007), followed by a cooling in 2007–09. In Disko Bay, Reference Holland, Thomas, de Young, Ribergaard and LyberthHolland and others (2008) reported an increase of 1.2˚C in 1991–96 to almost 3˚C in 1997–99 from fish trawl data, depth-averaged between 150 and 600m in their region 2 (Reference Wieland and KanneworffWieland and Kanneworff, 2002). Over the entire survey area, they report an increase from 1.7˚C to 3.5˚C until year 2002. Over the same period, the model indicates a 2˚C warming, which agrees with the fish trawl data, and a continuous warming after 2002 for a total increase of 2.5˚C. In terms of timing, the modeled change in ocean temperature at J_S is abrupt in 1995–98 and moderate in 1998–2002, again consistent with the fish trawl data.

Few observations exist in eastern Baffin Bay near M_S and M_O to evaluate the realism of the simulated trend. As discussed in Section 3.2, existing data suggest that the model started 1˚C too cold and the modeled warming is also 1˚C too large. At P_S, Reference Johnson, Münchow, Falkner and MellingJohnson and others (2011) report only 0.25˚C warming from 2003 to 2009 versus ~0.4˚C in the model over that time period. We conclude that the model results in the north tend to overestimate the warming compared to observations.

3.5. Potential for ocean heat to penetrate glacial fjords

Our analysis suggests that over the sites considered, changes in ocean temperature off the shelf are transmitted rapidly on the shelf. Whether the on-shelf changes reach the glacier fronts depends on winds, air–ocean heat exchange within the fjords and, most importantly, sea-floor topography, i.e. the presence of sills that may block the migration of warm, subsurface waters. Most sill depths are not well known (Reference BuchBuch, 2002). At Kangerdlugssuaq and Helheim (K_S and SE_S), the main sill depth of 550m should let warm subsurface waters penetrate the glacial fjord (Table 1). Sill depths farther south are unknown (Reference BuchBuch, 2002). At N_S, the main sill depth of 80 m should block warm-water inflow. At JKS, the 350 m deep sill is not sufficient to block warm waters from penetrating the fjord (Reference Holland, Thomas, de Young, Ribergaard and LyberthHolland and others, 2008), but at Torssukataq the sill depth is only 285 m. In northwest Greenland, there are no bathymetry data within 100 km of the coastline in the IBCAO database. At P_S, the 380 m depth sill should not block warm waters, but a shallower sill exists inside the fjord. At 79N_S, the sill depth is 20–200 m, so warm waters are blocked off.

3.6. Comparison of enhanced melt rates with glacier-front retreat and acceleration

Model simulations suggest that the subaqueous melt rate of a glacier calving face increases linearly with thermal forcing, defined as the difference between the water temperature and the in situ freezing point of sea water (Reference Xu, Rignot, Menemenlis and KoppesXu and others, 2012). We use this linear relationship herein to calculate the percentage increase in subaqueous melt rate resulting from an increase in ocean temperature (Table 1). We use a constant freezing point of sea water of –2.08˚C for a depth of 250 m and a salinity of 34.5 psu.

At K_S, Kangerdlugssuaq Gletscher doubled its speed in 2004–05, retreated 6 km, and then readvanced by 2 km in 2007–09 with a 50% reduction in speed (Reference Howat, Joughin and ScambosHowat and others, 2007). In the model, warm waters intruded at K_S from 1996 to 2004, followed by a slight cooling. The modeled temperatures translate into a 24% increase in melt rate over the entire time period and 36% between 1992 and 2004 (Table 1). At Helheim, the glacier accelerated in 2002–03 and doubled its speed by 2005; the ice front retreated by 7 km before readvancing by 3 km in 2005–06 and slowing down by 25% (Reference Howat, Joughin and ScambosHowat and others, 2007). The modeled changes in temperature at SE_S suggest a 14% increase in subaqueous melt for 1993–2009 and 21% for 1993–2004. The 1993–2004 increase in melt rate may have been sufficient to retreat the glacier front behind a shallow bathymetric feature (Reference JoughinJoughin and others, 2008) and trigger speed-up. The timing of the retreat and advance coincides with the warming and subsequent cooling of ocean waters at SE_S after 2004. A similar time evolution is found in southeast Greenland (Reference MurrayMurray and others, 2010).

Near Nuuk, glacier changes at Kangiata Nunata have been few despite the presence of warm waters on the continental shelf, probably because shallow sills block the entrance of warm subsurface waters. In Disko Bay, the ocean warming translates into a 68% increase in melt rate, the largest in Greenland. Farther north, in Torssukataq fjord, glacier changes have been less significant, possibly due to the presence of a shallower sill.

In northwest Greenland, glacier acceleration along the coast has been sporadic (e.g. Reference Moon and JoughinMoon and Joughin, 2008), but the glaciers have consistently lost mass since the 1990s and the mass loss accelerated after 2006 (Reference Khan, Wahr, Bevis, Velicogna and KendrickKhan and others, 2010). In fjords where subsurface warm waters intruded, our results suggest a potential increase in subaqueous melting of 65%, which is comparable to the record increase in Disko Bay.

At Petermann Fjord, no significant glaciological changes occurred until a major calving event took place in 2010 (Reference Johnson, Münchow, Falkner and MellingJohnson and others, 2011). Further observations and model evaluation are needed to determine whether ocean warming had an impact on the stability of the floating ice tongue. In the northeast, 79 North glacier has been stable since the 1990s, which is consistent with the presence of shallow sills at the mouth of the fjord.