Warm subtropical-origin Atlantic water flows northward across the Greenland-Scotland Ridge into the Nordic Seas, where it relinquishes heat to the atmosphere and gradually transforms into dense Atlantic-origin water. Returning southward along east Greenland, this water mass is situated beneath a layer of cold, fresh surface water and sea ice. Here we show, using measurements from autonomous ocean gliders, that the Atlantic-origin water was re-ventilated while transiting the western Iceland Sea during winter. This re-ventilation is a recent phenomenon made possible by the retreat of the ice edge toward Greenland. The fresh surface layer that characterises this region in summer is diverted onto the Greenland shelf by enhanced onshore Ekman transport induced by stronger northerly winds in fall and winter. Severe heat loss from the ocean offshore of the ice edge subsequently triggers convection, which further transforms the Atlantic-origin water. This re-ventilation is a counterintuitive occurrence in a warming climate, and highlights the difficulties inherent in predicting the behaviour of the complex coupled climate system.
Without redistribution of heat by the atmospheric circulation and ocean currents, only a small portion of the Earth’s surface would be habitable. In the Atlantic Ocean the poleward transport of heat is largely accomplished by the Atlantic meridional overturning circulation (AMOC), in which warm waters are transported northward near the surface and cold waters are returned to the south at depth1. Most of the North Atlantic deep waters originate from the Nordic Seas2, as result of a warm-to-cold transformation that takes place primarily in the eastern part of that region3,4,5. The resulting product, referred to as Atlantic-origin water6, is returned southward by the East Greenland Current7 (Fig. 1). This is a main source of dense water to the overflow plume that passes between Iceland and Greenland through Denmark Strait8 and provides the largest and densest contribution to the lower limb of the AMOC9.
The other major contribution to the Denmark Strait overflow is Arctic-origin water formed in the interior Iceland and Greenland Seas10. Proximity to the ice edge, where the highest ocean to atmosphere fluxes of heat occur due to frequent intense cold air outbreaks in which frigid, dry air masses are advected over comparatively warm surface water, is an important factor11,12,13,14. Diminished heat loss from this region owing to a retreat of the ice edge and different rates of warming of the ocean and the atmosphere will likely reduce the formation of Arctic-origin water15. In the Iceland Sea this water mass is primarily produced on the north-western outskirts of the cyclonic gyre, where the atmospheric forcing is more intense13,16. Farther to the west (within the region outlined in black in Fig. 1), the influence of fresh, low-density surface water becomes dominant17,18. The presence of this water mass is known to inhibit convection19,20, but thus far a lack of wintertime measurements from the westernmost part of the Iceland Sea has prevented verification13. Here we show that this fresh, low-density surface layer is not present in the western Iceland Sea in winter, which permits the formation of deep mixed layers offshore of the ice edge. As the ice edge retreats toward Greenland, water masses including the Atlantic-origin water transported by the East Greenland Current that were previously insulated from the atmosphere underneath sea ice are now being ventilated.
Re-ventilation of Atlantic-origin water along east Greenland
In the summer and fall of 2015 three autonomous gliders (see the Methods section for details of the processing and calibration of the glider data) were deployed with the purpose of mapping the extent of convection in the Iceland Sea during the subsequent winter. Here we focus on measurements obtained near the ice edge in the western Iceland Sea. Vertical sections of potential temperature and salinity from a cruise in summer 2012 and the two glider transects that approached the ice edge in winter 2015–2016 are shown in Figs. 2 and 3. The summer section has a pronounced 50–100 m thick fresh, low-density layer at the surface and a well-defined Atlantic-origin water mass beneath characterised by intermediate maxima in temperature and salinity (Figs. 2a and 3a). By late winter the surface layer had vanished and only a trace of the intermediate temperature maximum is visible (θ ≥ 0 °C). Instead, the western end of the February 2016 glider transect has an approximately 400 m deep mixed layer with a density of σ θ = 28.01–28.02 kg/m3, more than sufficient to supply the dense Denmark Strait overflow plume (Figs. 2b and 3b). As this was still relatively early in the convective season, most likely the depth and density of the mixed layer increased further into winter due to continued heat loss to the atmosphere13. The second glider encountered a re-stratifying water column farther to the south in late April, at the end of the convective season, that also shows similarly modified Atlantic-origin water (Figs. 2c and 3c). Both gliders were turned around due to the proximity of sea ice prior to reaching the core of the East Greenland Current. Hence, the extent to which the bulk of the Atlantic-origin water mass was further transformed by convection and how readily the modified waters were transported toward Denmark Strait remain uncertain. However, the glider transects clearly demonstrate that the fresh surface layer characteristic of the western Iceland Sea in summer does not prevent wintertime convection off the ice edge and indicate that the Atlantic-origin water transported by the East Greenland Current is re-ventilated by direct contact with the atmosphere while transiting this region in winter.
The western Iceland Sea is a sparsely sampled region, in particular in winter. Nonetheless, using newly available data from instrumented seals21 in combination with historical measurements13, the monthly mean hydrographic properties averaged over the upper 50 m of the water column demonstrated a pattern that is consistent with the shipboard and glider measurements (Fig. 4). The fresh surface layer starts developing in May or June and is most pronounced between July and August with a salinity well below 34. In October the surface layer becomes substantially more saline, and by November a well-defined surface layer is no longer present. At that time the water column has an ~80 m deep mixed layer and is well preconditioned for convection.
Local formation of deep mixed layers
In order to investigate whether the deep mixed layers observed at the western turnaround point of the February glider transect may have formed locally, a one-dimensional model (see the Methods section and ref. 22 for details of the model) was employed. The August 2012 profile closest to the turnaround point was used as initial conditions and a constant heat loss of 120 W/m2, typical for winter 2015–2016 (Fig. 5b), was applied from November to mid-February. As expected from previous work in this region19,20, the fresh surface layer prevented convection, resulting in effect instead in the formation of sea ice as soon as the cooling commenced. This demonstrates that the disappearance of the fresh water is not the result of vertical mixing caused by heat loss to the atmosphere. Another mechanism is instead required to remove the freshwater prior to the onset of convection.
Strong northerly winds are prevalent along the east coast of Greenland in fall and winter23. We estimate the onshore Ekman transport induced by these northerly winds by assuming an Ekman layer depth of 50 m, and find a pronounced seasonal cycle (Fig. 5a). While there is no substantial Ekman transport in summer, it increases considerably through fall. We estimate that the Ekman transport is sufficient to flush the fresh, low-density layer onto the Greenland shelf, thereby preconditioning the western Iceland Sea region for convection before the atmospheric forcing peaks between December and March (Fig. 5b). The integrated Ekman transport distance through the strong forcing season of order 300 km sets the horizontal scale of the band that is preconditioned for convection. We note that such northerly barrier winds, which often occur along the entire east coast of Greenland in winter24,25, inhibit the offshore diversion of freshwater and hence limit its impact on convection in the interior Nordic and Irminger Seas. In 2015–2016 the non-linear combination of unusually strong northerly winds and high heat fluxes at the beginning of winter, along with nearly ice-free conditions, were particularly conducive for convection (Fig. 5). Melting of sea ice in summer, when onshore Ekman transport is negligible, helps replenish this fresh surface layer26.
A new simulation with the one-dimensional model was initialised in which the fresh surface layer was replaced by a mixed layer extending from 80 m depth to the surface, in agreement with the mean November profile from the combined seal and historical hydrographic data set. The resulting simulated profile (red trace in Fig. 6) corresponds very closely to the profile recorded by the glider at the turnaround point in February 2016 (blue trace in Fig. 6), even though the initial conditions were obtained from summertime measurements 4 years prior to the glider transect. This implies that the deep and dense mixed layer recorded by the glider in February 2016 is consistent with local formation as the result of convection offshore of the ice edge.
A consequence of the re-ventilation of Atlantic-origin water east of Greenland may have been detected farther downstream. At the Denmark Strait sill there is no discernible seasonal variability in the overflow transport, but a significant seasonal signal in temperature. In particular, the temperature of the Denmark Strait overflow is minimum in September27, which, assuming an advective speed of 4–5 cm/s, is consistent with the transport of colder Atlantic-origin water that was re-ventilated the previous winter in the Iceland Sea. Taking into account that the re-ventilation may also have taken place farther north along the ice edge east of Greenland, this is in reasonable agreement with observed velocities of the Atlantic-origin water in the East Greenland Current7,8.
The sea ice concentration in the western Iceland Sea has diminished substantially over the last few decades15. There was, in particular, very little sea ice in winter 2015–2016 relative to climatological values (Fig. 5c). As a consequence, increasing areas along the Greenland continental slope, and hence also of the East Greenland Current, are no longer insulated by sea ice from interaction with the atmosphere in winter. If the ice edge continues to retreat toward Greenland, we may expect to see more re-ventilation of Atlantic-origin water and higher heat fluxes into the atmosphere along the pathway of the East Greenland Current from Fram Strait to Denmark Strait. If such re-ventilation becomes more pronounced, this could further modify the properties of the Denmark Strait overflow and thereby impact the lower limb of the AMOC. While large-scale sea ice loss may lead to a weakening of the AMOC28, the ice edge retreat toward Greenland causing increased ventilation of Atlantic water in the Nordic Seas during a warming climate is unexpected and contrary to the tendency of reduced convection generally predicted by climate models29.
The wintertime transects were obtained by autonomous, buoyancy-driven Seagliders that are capable of diving to 1000 m depth30. The gliders profile in a sawtooth pattern with a typical vertical to horizontal glide ratio of 1:3 that results in a speed through water of about 20 cm/s. Conductivity and temperature were sampled at 20–80 s intervals, resulting in a vertical resolution of 1–5 m, enhanced within the surface mixed layer. Following a tuning of the flight model parameters for each glider, the data were reprocessed and corrections were applied to compensate for thermal-inertia and flushing speed issues arising from usage of unpumped sensors31. Temperature and salinity measurements outside the expected range of values in the Nordic Seas (−2–20 °C and 20–36, respectively) were discarded. Each dive/climb cycle was subsequently inspected for density inversions, and measurements causing inversions exceeding 0.05 kg/m3 were excluded18,32,33. The sensors were laboratory calibrated prior to deployment. In addition, the glider data were calibrated against shipboard measurements at the times of deployment and recovery. A rendezvous partway through the deployment was used for data intercomparison. Corrections were estimated in potential conductivity–potential temperature space34. Offsets corresponding to 0.006 and 0.016 in salinity were applied to two of the gliders (sg559 and sg562, respectively), while a drift over the entire deployment period corresponding to 0.0008 in salinity was corrected for one glider (sg559). No corrections were applied to the conductivity sensor of the third glider (sg564) or any of the temperature sensors.
Wind stress, surface sensible and latent heat fluxes, as well as sea ice concentration for the climatological analyses (Fig. 5) were taken from the European Centre for Medium-Range Weather Forecasts interim reanalysis (ERA-Interim) for the period 1979–2016. Estimates of wind stress and heat fluxes were derived from short-range forecasts averaged over 6-hourly intervals (using forecast steps for 9–21 h). Only ice-free grid cells, with a sea ice concentration of less than 0.5, were considered.
The Ekman transport distance was estimated from
where τ(t) is the 6-hourly average mean wind stress in the ice-free portion of the western Iceland Sea region (Fig. 1), ρ = 1025 kg/m3 is the reference density of sea water, f is the coriolis parameter, and n⊥ is the box mean unit vector perpendicular to Greenland’s shelf. For the depth of the Ekman layer a constant value of h = 50 m was assumed, which corresponds approximately to the typical depth of the sharp pycnocline that is associated with the summertime fresh surface layer (Fig. 6). We note that the precise depth of the Ekman layer cannot be determined based on the available observations but the value of 50 m is in line with depth estimates based on the similarity height (e.g., ref. 36). Since the Ekman transport distance is inversely proportional to Ekman depth, a decrease of h by 50% would result in an increase of the estimated distance by a factor of 2, whereas an equivalent increase of h would result in a lower distance by one-third.
To implement the one-dimensional PWP22 mixed-layer model, turbulent heat fluxes, which provide the dominant contribution to mixed-layer deepening37, were imposed at the surface at each time step. Initial conditions were obtained from the August 2012 shipboard and February 2016 glider measurements. The depth and properties of the mixed layer were then adjusted until three stability criteria, chief among which is static stability, were satisfied.
The glider data can be accessed at the Pangaea repository (https://doi.org/10.1594/PANGAEA.884339). The shipboard data can be obtained from http://kogur.whoi.edu. The seal data can be found in the MEOP database (www.meop.net). The ERA-Interim reanalysis data were obtained from the European Centre for Medium-Range Weather Forecasts and the passive microwave sea ice product38 was obtained from the National Snow and Ice Data Center.
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We are deeply indebted to the NACO team at the University of Bergen for the successful deployment and operation of the gliders, in particular Erik Magnus Bruvik, Idar Hessevik, Karsten Kvalsund, and Tor de Lange. We thank Héðinn Valdimarsson for assistance during deployment and recovery of the gliders, and James Bennett and Bastien Queste for a very helpful tutorial on glider data processing. Dallas Murphy and Pål Erik Isachsen kindly provided valuable comments to the manuscript. We thank Kit Kovacs and Christian Lydersen, the Norwegian Polar Institute and MEOP (NFR grant number 176477), and Pål Erik Isachsen for providing access to the calibrated seal data set. Support for this work was provided by the Norwegian Research Council under Grant agreement no. 231647 (L.H. and K.V.), the Bergen Research Foundation under Grant BFS2016REK01 (K.V.), and the Centre for Climate Dynamics at the Bjerknes Centre through the FRESHWATER project (K.V.). Additional funding was provided by the Swiss National Science Foundation grants P2EZP2162267 and P300P2174307 (L.P.), the National Science Foundation grant OCE-1558742 (M.A.S.), the Norway Fulbright Foundation (K.V.), the Canada Fulbright Foundation (G.W.K.M.), and the Natural Sciences and Engineering Research Council of Canada (G.W.K.M.).
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