Beaufort Gyre energy sources and sinks

The mechanical energy budget is formulated by considering the depth evolution of the halocline to derive an expression for the rate of change of APE in the BG region (see “Methods”). The terms include wind energy input (i.e., the work done by the ocean surface stress on the surface geostrophic circulation), boundary thickness fluxes due to Ekman transport across open lateral boundaries, and eddy diffusion. To estimate these terms, we calculate the ocean surface stress from reanalysis wind, satellite-derived sea ice drift and concentration, and estimates of surface geostrophic currents produced by combining conventional radar altimetry data from the open ocean with specialized altimetry retrievals from openings within the sea ice pack11,13. This allows us to calculate the wind work and the cumulative wind energy input to the BG region (Fig. 2a, b), as well as the Ekman pumping (Fig. 2e, f). We make use of in situ hydrographic data to estimate the difference in density across the halocline, and use the relationship between the hydrography-derived halocline depth and DOT to estimate halocline depth across the BG region. This allows us to estimate the APE and boundary thickness fluxes (Fig. 2a, c) as well as the kinetic energy (KE) of the geostrophic circulation, which is negligible relative to the other terms. The residual of the wind energy input, Ekman boundary fluxes, and APE storage then represents the energy dissipation by eddies (Fig. 2c).

Fig. 2: Beaufort Gyre Energy and Freshwater Budget. a The atmosphere-ocean (green), ice–ocean (blue), and net (atmosphere-ocean + ice-ocean; gray) cumulative energy input to the Beaufort Gyre region (gray), and the boundary thickness fluxes due to Ekman transport (pink). b The monthly ice–ocean (blue) and atmosphere-ocean (green) seasonal cycle of power input before (solid lines) and after (dashed lines) 2007. c The available potential energy (APE, orange) and cumulative eddy dissipation (purple) in the Beaufort Gyre region, and d the Beaufort Gyre liquid freshwater content estimates3. e The Ekman upwelling (blue), downwelling (green), and net pumping (gray) in the Beaufort Gyre region, and f the seasonal cycle of Ekman pumping before (solid lines) and after (dashed lines) 2007. The gray box in c corresponds to the period when the relationship between DOT and halocline depth breaks down and our estimates of APE and eddy dissipation become unreliable (see “Methods”). The spread on the data in a, e represents the difference in the calculation of wind energy input and Ekman pumping using two different sea ice drift data sets, and the spread in c also incorporates the uncertainty of the fit between dynamic ocean topography and halocline depth (see “Methods”). Full size image

Wind energy input to the BG is highly seasonal, spatially inhomogeneous, and heavily dependent on the sea ice cover and atmospheric circulation (Figs. 2b and 3). The BG gains energy from the winds in the south, and loses energy in the north over a mean annual cycle (Fig. 1f). Strong atmospheric circulation in the autumn, combined with significant areas of open water, means that work done by the atmosphere directly on the surface geostrophic currents (W ao ) dominates energy input to the BG. Around 60% of the total wind power input occurs in September and October in the ice-free southern BG, where easterly along-shore winds do work on the westward flowing southern limb of the BG (Fig. 3). Averaged over the BG region, sea ice acts to dissipate energy in most months (Fig. 2b), dominated by negative ice–ocean surface stress work (W io ) in the northern BG region during winter (Fig. 3). This is the ice–ocean governor from the perspective of mechanical energy: the upper ocean flows beneath a relatively immobile ice pack and undergoes top-boundary drag, which dissipates energy while acting to spin down the gyre by inducing cyclonic stress and Ekman upwelling14,15,19. Ekman pumping is also at its seasonal maximum in the autumn (Fig. 2f), corresponding to the peak in the seasonal cycle of freshwater storage3,11.

Fig. 3: Seasonal wind energy input to the Beaufort Gyre. Seasonal climatologies of the atmosphere-ocean (a–d), ice–ocean (e–h), and total (i–l) wind energy input for winter (January–March), spring (April–May), summer (June–September), and autumn (October–December), as well as the mean seasonal cycles (m–o) before and after 2007. Full size image

Between 2003 and 2006, wind energy input was balanced by energy dissipation under sea ice over an annual cycle and there was not a significant net energy input to the BG region (Fig. 2a). At the same time, there was no significant change in APE during this period, implying that the role of eddies in dissipating energy was small. During this period, the BG can be thought of as existing in an “energetic governor regime”, where wind energy input was balanced by energy dissipation under sea ice and there was little change in APE storage. This corresponded to a period of net Ekman downwelling of ~170 mSv (Fig. 2e; 1 mSv ≡ 103 m3/s) and an increase in BG FWC of almost 2000 km3 over 4 years (Fig. 2d). However, this balance changed dramatically in late 2007, when a strong wind event injected ~30 PJ of wind energy into the BG region (Fig. 2a). Large open water areas associated with the then-record summer 2007 minimum sea ice extent7 (Fig. 4) meant that strong and persistent anticyclonic wind anomalies over the western Arctic did a large amount of work on surface currents in the southern BG that were flowing at least twice as fast as the climatological average13. FWC and APE in the BG increased significantly in response to this forcing (Fig. 2c, d) implying a potential for enhanced eddy generation by instabilities acting on the increased isopycnal slope. However the eddy response may lag the forcing by a few years24.

Fig. 4: Area of sea ice. Time series of sea ice area (106 km2) in the Beaufort Gyre region corresponding to the blue box in Fig. 1a. Full size image

After 2007, wind energy input and energy dissipation underneath sea ice both intensified and became more seasonal (Fig. 2a, b). Diminished summer and autumn ice cover (Fig. 4) led to higher wind energy input and generally less energy dissipation by sea ice between June and November; at the same time, increased surface currents in winter resulted in greater energy dissipation by sea ice between December and March (Fig. 2b). Between 2009 and 2014, there was a total wind energy input of ~9.5 PJ, the Ekman boundary flux added ~2.8 PJ, while APE decreased by ~3.2 PJ (Fig. 2a, c). We estimate ~15.7 PJ of energy dissipation by eddies between 2009 and 2014. After 2007, the ice–ocean governor mechanism was no longer able to fully dissipate the increased wind energy input, and the BG mechanical energy budget was out of equilibrium. Ekman upwelling and downwelling both increased after 2007, which will tend to produce a steeper halocline slope (Fig. 2e, f), however the net Ekman pumping actually decreased after 2007 (Fig. 2e, f) and the FWC remained relatively stable from 2008 onwards (Fig. 2d). We note that, as we use the 12-month smoothed DOT to estimate the halocline depth (see “Methods”), our estimates of APE, and hence eddy dissipation, break down during periods of rapid change (i.e., late 2007 ± ~1 year). This occurs because DOT varies faster than the halocline depth and the correlation between the two is weaker at short (~monthly) time scales.

Eddies role in Beaufort Gyre energy and freshwater balance

From energy budget considerations, the increased wind energy input to the BG after 2007, and the reduction in the APE reservoir after 2007, implies that energy dissipation associated with eddies must also have increased after 2007. The surface geostrophic currents data resolve ocean variability to a resolution of O (80 km), at monthly time scales13, leaving unresolved the energetic transient eddies with a characteristic scale of the Rossby radius, which is ~10–15 km in the Canada Basin29. There are several ways that transient surface eddies could contribute to increased energy dissipation. The cubic dependence of ice–ocean surface stress work on the characteristic eddy velocity, u eddy , means that energy dissipation by eddies is highly sensitive to small changes in u eddy . Changes in the ice–ocean drag coefficient, due to, e.g., changes in form drag30, or an increase in eddy density26, will also affect energy dissipation underneath sea ice. The spatial pattern of wind energy input to the BG (input along the southern edge, dissipation in the interior; Fig. 1f) implies a northward transport of energy from source to sink. Our method does not allow us to investigate this in detail and we can only speculate that the energy is transported by the mean geostrophic circulation as well as by preferential mesoscale eddy propagation away from the southern BG region, where strong baroclinic currents near the gyre boundaries facilitate eddy formation, towards the interior of the gyre where there is mechanical dissipation. As the eddies propagate to the interior of the gyre, they transport freshwater anomalies (halocline thickness anomalies) but also the eddy APE and eddy KE. Observations and theory of baroclinic ocean turbulence in the ice-free global ocean suggest that bottom drag is a key energy dissipation mechanism (e.g., Ferrari and Wunsch28). However, the Western Arctic is highly baroclinic, and currents at depth are an order of magnitude weaker than at the surface31, implying bottom drag work of order 10−3–10−2 GW, two or three orders of magnitude smaller than the surface energy input/dissipation (Fig. 2b). Hence, we conclude that an enhancement of upper-ocean turbulence, specifically the generation of mesoscale eddies, is critical to balance the increased wind-driven energy input into the BG.

Analogous to increased energy dissipation, the continued net Ekman downwelling after 2007 coupled with the stability of BG FWC after 2008 (Fig. 2d–f) implies a greater role for eddies in the FW budget. For the ice–ocean governor mechanism to fully compensate for increased Ekman pumping14,15,19, geostrophic currents must increase until Ekman upwelling, induced by ocean currents flowing beneath slower moving sea ice, fully compensates the Ekman downwelling. However, this is not the case in the Arctic after 2007. While the net Ekman downwelling does decrease after 2007, implying a role for an increased ice–ocean governor mechanism, increased halocline eddy activity must also play a role for the BG FWC to stabilize. From the expression for the eddy dissipation term (see “Methods”) we can estimate a value for K GM , the eddy diffusivity, of ~120 m2/s for the period after 2008 when eddies are found to play an important role in energy dissipation. This falls within the range of values estimated by Meneghello et al.27, and represents a gyre-wide mean that could be expected to vary locally and in time. The energy-budget approach together with considerations of Ekman-driven freshwater accumulation and release consistently reveal an increasing role for eddies in energy dissipation and FWC stabilization in the BG, particularly from 2008 onwards. Overall, the picture is one of a more energetic BG system since 2007, with increased energy sources and sinks and increased eddy diffusivity balancing and dissipating APE, and FWC stabilization by eddies required to balance net Ekman downwelling (Fig. 5).

Fig. 5: The changing components of the Beaufort Gyre energy budget. a Before and b after 2007, including the wind work, W (comprised of atmosphere-ocean, W ao , and ice–ocean, W io , components), available potential energy (APE), and eddy dissipation, W eddy . The atmosphere and ocean circulations are illustrated by u a and u g , respectively. The size of the arrows/vectors represents their relative strength. The loss of sea ice after 2007 led to increased wind energy input to the BG, increased APE, and increased energy dissipation and freshwater stabilization by eddies. Full size image

Implications for the changing Arctic

Arctic sea ice loss is projected to continue over the coming decades, with climate models predicting seasonally ice-free conditions (<106 km2) as early as the 2020s, but more likely towards the middle of the century7,32. Previous hypotheses suggested that the Arctic atmospheric circulation oscillates between predominantly cyclonic and anticyclonic circulation regimes over timescales of 5–7 years1, however, the Arctic has been in an anticyclonic regime since the late 1990s, coinciding with a period of increasing freshwater storage9. Were the Arctic to switch back to predominantly cyclonic atmospheric circulation we might expect a dissipation of mechanical energy by the atmosphere in summer and autumn months, as well as weaker (or reversed) net Ekman pumping and release of freshwater. Here, the results from 2012 provide a useful test case in contrast to the extreme of 2007. In 2012, as in 2007, there was an almost complete loss of sea ice in the BG region (Fig. 4), the cyclonic atmospheric circulation conditions actually dissipated energy in the summer (Fig. 2a) and were more favorable for upwelling (Fig. 2e) causing a (temporary) reduction in FWC (Fig. 2d). The difference in the BG response during extreme ice loss in 2012 compared to 2007 highlights the important interplay between atmospheric circulation and sea ice conditions that controls the state of the BG. A similar reversal of the prevailing anticyclonic atmospheric circulation was observed in wintertime 2016–17. This event was linked to increased intrusions of Atlantic cyclones due to thin ice in the Barents Sea region (and associated thermal anomalies), and a shift in the polar vortex33. Increased intrusions of cyclones into the western Arctic and further reversals in the wintertime atmospheric circulation due to declining sea ice in the Barents Sea is an intriguing, but highly uncertain, hypothesis34. These events suggest that a switch to more cyclonic circulation conditions would lead to a period of freshwater release and a spin down of the BG. However, regardless of the prevailing atmospheric circulation regime, we expect the Arctic Ocean to become more sensitive to atmospheric forcing as the sea ice cover continues to decline under climate change.

Our results show that as the BG region becomes increasingly ice-free earlier in summer and later into October and November, the current anticyclonic atmospheric circulation regime will do significantly more work on the ocean surface currents. Meanwhile, year-round dissipation of energy underneath sea ice will also increase as currents speed up (Fig. 2a, b), but our results suggest that this will not completely account for increased direct atmosphere-ocean energy input. Under this scenario, the Arctic Ocean will continue to become more energetic, and dissipation of additional energy and freshwater stabilization by eddies will be increasingly important. These are critical processes for the accurate representation of the BG system in models. However, currently, only the highest resolution numerical models are eddy-resolving in the Arctic Ocean, where the radius of deformation is 10–15 km in the deep basins and as small as 1 km in the shelf seas. Coupled climate models leave these important dissipative processes unresolved and it is unclear whether commonly used parameterizations of eddies, tuned for the global ocean, are representative in the Arctic. Increases in the eddy diffusivity and increased mixing by eddy activity in a more energetic Arctic Ocean is also expected to enhance vertical transport of warm Atlantic water, with consequences for sea ice growth and mixing of biogeochemical tracers.