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    Climatology of Arctic freshwater content integrated vertically from the surface to the depth of the 34.8 isohaline from the PHC (Steele et al. 2001). The Beaufort Gyre region as defined by Proshutinsky et al. (2009) and Giles et al. (2012) is highlighted by the black box, and clearly stands out as the region of greatest freshwater content in the Arctic.

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    Reduced gravity model of the Canada Basin with an active surface halocline layer of density ρ − Δρ initialized to be 400 m thick, overlying a stationary (u = 0) Atlantic layer of density ρ. The model domain is ~1500 km in diameter and is connected to a sponge region by a channel ~150 km wide. The sponge region is designed to absorb any waves propagating out of the domain, and to allow an inflow or outflow transport to develop in response to changes in layer thickness within the domain. The model is forced with an anticyclonic ocean surface stress centered over the domain. The magnitude of the curl of the stress field (which is proportional to the strength of the Ekman pumping) is at a maximum in the center of the domain, and is zero at the boundaries and in the outflow.

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    Summer (July–October) and winter (November–June) averages of (a) mean sea level pressure from the ERA-Interim reanalysis and (b) sea ice concentration from the SSM/I satellite instrument over the Arctic Ocean between 1979 and 1990. (c) The annual cycle in ocean surface stress used to force the control run is based on the occurrence of stronger and more anticyclonic winds over the Beaufort Gyre during the 8 months of winter (green) compared to the four months of summer (red). The reduction in sea ice cover between July and October defines our summer period.

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    (a) Mean layer thickness (colors) and velocity field (vectors), (b) cross section of the mean layer thickness through the center of the domain (i.e., along the white line), (c) the annual cycle in layer thickness in the center of the domain (i.e., at the white dot), and (d) the annual cycle in transport through the channel (i.e., across the red line) all calculated over the last 5 years of the control run. Positive transports indicate an inflow and negative transports indicate an outflow. Integer year numbers correspond to the end of our summer period.

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    Modifications made to the annual cycle in ocean surface stress with respect to the control run (thick blue line) for (a),(b) the modified seasonality experiments and (c),(d) the modified net forcing experiments. In (a),(c), the ocean surface stress is incrementally increased in summer, representing the increased efficiency of momentum transfer. In (b),(d), the period over which high values of summer stress is applied is then incrementally increased, representing a longer melt season. For just the modified seasonality experiments in (a),(b), the stress in winter in each model run is reduced, to ensure that when integrated over the entire year the net forcing remains constant at 0.02 N m−2 (gray dashed line).

  • View in gallery

    The annual cycle in (a),(d) ocean surface stress, (b),(e) the layer thickness in the center of the domain (white dot in Fig. 4a), and (c),(f) the transport through the channel (red line in Fig. 4a) from the last 5 years of each of the modified seasonality experiments. Positive transports indicate an inflow and negative transports indicate an outflow. Note that (a)–(c) are for model runs where the ocean surface stress in summer has been increased, while (d)–(f) are for model runs where the melt season length has been increased. Line styles and colors correspond to those used in Figs. 5a and 5b. The thick blue line represents the control run. Integer year numbers correspond to the end of our summer period.

  • View in gallery

    Results of the modified net forcing experiments. (a) The increase in mean layer thickness in the center of the domain plotted as a function of the increase in net forcing. The blue dot represents the control run, the red dots represent model runs with an increased ocean surface stress in summer, and the green dots represent model runs with a longer melt season (as per the colors in Figs. 5c and 5d). (b) The cumulative volume and freshwater transports into the domain over the full 40 years of each model run as a function of the increase in net forcing.

  • View in gallery

    The blue lines show the mean annual cycle in liquid freshwater content anomaly integrated between the surface and 25 m (dotted line) and between the surface and the depth of the 34.8 isohaline (solid line) over the Beaufort Gyre region from the MIMOC. The red line shows the mean annual cycle in the solid (ice) freshwater content anomaly from upward-looking sonar deployed as part of the Beaufort Gyre Exploration Project. The black line is the mean annual cycle in Ekman pumping velocity at the base of the Ekman layer over the Beaufort Gyre region, calculated from the wind stress curl. Positive velocities indicate a downwelling.

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On the Link between Arctic Sea Ice Decline and the Freshwater Content of the Beaufort Gyre: Insights from a Simple Process Model

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  • 1 Department of Earth Sciences, University of Oxford, Oxford, United Kingdom
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Abstract

Recent satellite and hydrographic observations have shown that the rate of freshwater accumulation in the Beaufort Gyre of the Arctic Ocean has accelerated over the past decade. This acceleration has coincided with the dramatic decline observed in Arctic sea ice cover, which is expected to modify the efficiency of momentum transfer into the upper ocean. Here, a simple process model is used to investigate the dynamical response of the Beaufort Gyre to the changing efficiency of momentum transfer, and its link with the enhanced accumulation of freshwater. A linear relationship is found between the annual mean momentum flux and the amount of freshwater accumulated in the Beaufort Gyre. In the model, both the response time scale and the total quantity of freshwater accumulated are determined by a balance between Ekman pumping and an eddy-induced volume flux toward the boundary, highlighting the importance of eddies in the adjustment of the Arctic Ocean to a change in forcing. When the seasonal cycle in the efficiency of momentum transfer is modified (but the annual mean momentum flux is held constant), it has no effect on the accumulation of freshwater, although it does impact the timing and amplitude of the annual cycle in Beaufort Gyre freshwater content. This suggests that the decline in Arctic sea ice cover may have an impact on the magnitude and seasonality of the freshwater export into the North Atlantic.

Corresponding author address: Peter Davis, Department of Earth Sciences, University of Oxford, South Parks Road, Oxford, OX1 3AN, United Kingdom. E-mail: peterd@earth.ox.ac.uk

Abstract

Recent satellite and hydrographic observations have shown that the rate of freshwater accumulation in the Beaufort Gyre of the Arctic Ocean has accelerated over the past decade. This acceleration has coincided with the dramatic decline observed in Arctic sea ice cover, which is expected to modify the efficiency of momentum transfer into the upper ocean. Here, a simple process model is used to investigate the dynamical response of the Beaufort Gyre to the changing efficiency of momentum transfer, and its link with the enhanced accumulation of freshwater. A linear relationship is found between the annual mean momentum flux and the amount of freshwater accumulated in the Beaufort Gyre. In the model, both the response time scale and the total quantity of freshwater accumulated are determined by a balance between Ekman pumping and an eddy-induced volume flux toward the boundary, highlighting the importance of eddies in the adjustment of the Arctic Ocean to a change in forcing. When the seasonal cycle in the efficiency of momentum transfer is modified (but the annual mean momentum flux is held constant), it has no effect on the accumulation of freshwater, although it does impact the timing and amplitude of the annual cycle in Beaufort Gyre freshwater content. This suggests that the decline in Arctic sea ice cover may have an impact on the magnitude and seasonality of the freshwater export into the North Atlantic.

Corresponding author address: Peter Davis, Department of Earth Sciences, University of Oxford, South Parks Road, Oxford, OX1 3AN, United Kingdom. E-mail: peterd@earth.ox.ac.uk

1. Introduction

The Arctic Ocean receives an input of freshwater from Eurasian and North American river runoff, net precipitation over evaporation (positive PE), and sea ice melt. The Atlantic contributes relatively salty water via Fram Strait and the Barents Sea, and the Pacific relatively freshwater via Bering Strait. After significant watermass transformation, freshwater is exported in both liquid and solid (ice) form into the Nordic and Labrador Seas, through Fram Strait and the Canadian Arctic Archipelago, respectively.(Coachman and Aagaard 1974; Aagaard and Carmack 1989; Carmack 2000; Serreze et al. 2006; Dickson et al. 2007). A moderate freshening of the surface waters in these exit regions may impact the formation of North Atlantic Deep Water, and thus affect both the global thermohaline circulation and the global climate (Dickson et al. 1988; Aagaard and Carmack 1989; Vellinga and Wood 2002; Jones and Anderson 2008).

Within the Arctic Ocean itself, more than 70 000 km3 of freshwater is stored within a very fresh surface layer (Aagaard and Carmack 1989; Serreze et al. 2006; Fig. 1), which is separated by a strong halocline from the relatively warm and saline Atlantic-derived layer beneath (Aagaard et al. 1981). Over the past few decades, observations have shown that the freshwater content has been increasing (Proshutinsky et al. 2009; McPhee et al. 2009; Rabe et al. 2011; Krishfield et al. 2014), with the largest freshening observed in the Beaufort Gyre in the Canada Basin (Fig. 1). Recently, Giles et al. (2012) have suggested that the rate of freshening in the Beaufort Gyre may be accelerating, and have proposed that a dynamical adjustment in response to the decline in Arctic sea ice cover may be the cause. Understanding this dynamical response is the focus of this study.

Fig. 1.
Fig. 1.

Climatology of Arctic freshwater content integrated vertically from the surface to the depth of the 34.8 isohaline from the PHC (Steele et al. 2001). The Beaufort Gyre region as defined by Proshutinsky et al. (2009) and Giles et al. (2012) is highlighted by the black box, and clearly stands out as the region of greatest freshwater content in the Arctic.

Citation: Journal of Climate 27, 21; 10.1175/JCLI-D-14-00090.1

The Beaufort Gyre is a permanent anticyclonic circulation driven by the winds associated with the atmospheric Beaufort high (Proshutinsky et al. 2009). These winds cause water to converge in the center of the gyre, and the resulting downwelling (Ekman pumping) leads to an accumulation of freshwater through the mechanical deformation of the salinity field (Proshutinsky et al. 2002; Yang 2009; Proshutinsky et al. 2009). On annual time scales, stronger anticyclonic winds in winter result in an accumulation of freshwater, whereas weaker (or possibly cyclonic) winds in summer relax the salinity gradient resulting in a release of freshwater (Proshutinsky et al. 2002). In addition, ice growth in winter and melt in summer will decrease or increase the freshwater content of the Beaufort Gyre respectively, and any changes in the advection of surface Pacific and Atlantic waters are also likely to have an impact (Proshutinsky et al. 2009). Consequently, the mean annual cycle in freshwater content reflects the complex interplay among several different forcings.

On interannual time scales, Proshutinsky and Johnson (1997) and Proshutinsky et al. (2002) have suggested that two different wind regimes exist within the Arctic: cyclonic and anticyclonic. During the anticyclonic regime, freshwater is accumulated within the Beaufort Gyre over several years because of a strengthened atmospheric Beaufort high. In contrast, during the cyclonic regime the atmospheric Beaufort high weakens, and freshwater is released to the shelves where it may be exported into the North Atlantic. As a result, there is a strong linear relationship between the freshwater content of the Beaufort Gyre and the wind stress curl on interannual time scales (Proshutinsky et al. 2002, 2009). Since 1997, the Arctic has been in the longest anticyclonic regime on record, leading to a large excess accumulation of freshwater within the Beaufort Gyre. Using hydrographic data, Rabe et al. (2011) have shown that the Beaufort Gyre has gained 8400 ± 2000 km3 of freshwater between the periods 1992–99 and 2006–08, and Giles et al. (2012) used satellite observations of sea surface height between 1995 and 2012 to infer a similar increase of 8000 ± 2000 km3.

The results of Giles et al. (2012) show a clear spatial correlation between the positive trend in the sea surface height associated with the Beaufort Gyre, and the negative trend in the curl of the wind field (i.e., a trend toward more anticyclonic winds). However, despite this clear spatial correlation, the temporal correlation is less clear. While the trend in the curl of the wind field remained constant over the whole period, the trend in the sea surface height changed sign after 2002. Consequently, Giles et al. (2012) have suggested that since 2002 the winds may have become more effective at driving an accumulation of freshwater within the Beaufort Gyre, because of an increase in the efficiency of momentum transfer into the upper ocean. On the other hand, Morison et al. (2012) argue that a shift in the advection pathways of Eurasian river water, associated with a positive phase of the Arctic Oscillation during 2005–08, can account for the increased accumulation of freshwater, with no role played by the Beaufort Gyre wind-driven circulation. In reality, it is likely that both processes are important (Mauritzen 2012).

In the absence of sea ice, the magnitude of the momentum transfer into the upper ocean (i.e., the ocean surface stress) is determined simply by the surface wind stress. However, in the partly ice-covered Arctic, the magnitude of the momentum transfer is determined by both the surface wind stress (τAir-Water) and an ice-water stress component (τIce-Water), with their relative magnitudes scaled by the ice concentration (α; Yang 2009):
e1
In the past, the thick and extensive Arctic sea ice cover has acted to reduce the momentum transfer, due to the large internal ice stresses reducing the ice-water stress component, and shielding of the ocean from direct wind forcing. However, as the sea ice cover has begun to break up and retreat farther and for longer each year (Stroeve et al. 2007; Maslanik et al. 2007; Cavalieri and Parkinson 2012), it has become both weaker and thinner (Kwok and Rothrock 2009; Zhang et al. 2012; Laxon et al. 2013), and the number of leads, melt ponds, and ice floe edges has increased, changing the shape of the ice pack (Flocco et al. 2012; Tsamados et al. 2014). Consequently, the efficiency of momentum transfer into the upper ocean has increased, as a thinner and weaker sea ice cover is more easily forced by the winds, and the changing shape of the ice pack is providing more and more near-vertical faces for the wind to push against (Andreas et al. 2010). As a result, not only is the annual mean ocean surface stress (τOceanSurface) increasing (i.e., the net forcing), but its seasonality is also changing. For example, Martin et al. (2014) have shown using the Pan-Arctic Ice–Ocean Modeling and Assimilation System (PIOMAS) model that the Arctic-wide annual mean ocean surface stress has increased by 0.006 N m−2 decade−1 between 2000 and 2012, corresponding to an increase of approximately 9% decade−1 based on the long-term mean stress of 0.064 N m−2. Furthermore, Tsamados et al. (2014) have shown using the Los Alamos sea ice model (CICE) and a new parameterization for form drag, that the drag coefficient (which can be taken as a measure of the efficiency of momentum transfer into the upper ocean) exhibits a small positive trend over the Beaufort Gyre in summer between 1990 and 2012.

The accelerated accumulation of freshwater in the Beaufort Gyre has coincided with the dramatic decline observed in Arctic sea ice cover. Has the corresponding increase in the annual mean ocean surface stress, or its changing seasonality, contributed to the accelerated accumulation and thus modified the linear relationship observed between the freshwater content and the wind stress curl (Proshutinsky et al. 2002, 2009)? The changing seasonality may have important implications for the adjustment of the Arctic Ocean to a change in forcing. If the dominant adjustment time scale is on the order of a season, modifications to the length of each season, or an asymmetry between spinup and spindown time scales over the annual cycle, may result in a multiyear trend in freshwater accumulation.

The dynamical response of the Beaufort Gyre to the thinning, weakening, and changing shape of the Arctic sea ice cover will depend upon exactly how much more stress is transferred from the surface of the ice pack to the ocean below (i.e., the change to the annual mean ocean surface stress) and how the seasonal distribution of this transfer may change (i.e., the change to the seasonality in the ocean surface stress). However, both remain poorly constrained and understood, and many different processes such as stratification, atmospheric boundary layer stability, ocean circulation, and sea ice conditions are important in determining the transfer of momentum through sea ice (McPhee 2012). Indeed, two state-of-the-art studies by Martin et al. (2014) and Tsamados et al. (2014), which were designed to investigate how the magnitude and seasonality of the ocean surface stress have changed with the recent decline in Arctic sea ice cover, show seasonal cycles in ocean surface stress that are 180° out of phase. Here we adopt the simplest possible approach and use an idealized process model to investigate the dynamical response of the Beaufort Gyre to the changing efficiency of momentum transfer, and its link to the accelerated accumulation of freshwater, by idealistically perturbing the annual cycle in ocean surface stress through a wide parameter space. We do not explicitly represent sea ice in the model, but rather account for the effect of its decline via perturbations made to the magnitude and seasonality of the ocean surface stress. In section 2 we describe the model setup, before presenting the results of a control run in section 3. In sections 4 and 5 the setup and results of two different sets of idealized experiments designed to investigate the response of the Beaufort Gyre to the decline in Arctic sea ice cover are presented, with a discussion of the results and their implications in section 6. Our conclusions are presented in section 7.

2. Model setup

A nonlinear 1.5-layer reduced gravity model (e.g., Johnson and Marshall 2002; Allison et al. 2011) is used to simulate the surface halocline layer of the Canada Basin in the Arctic Ocean. The model is governed by the following equations:
e2
e3
where u is the velocity, t is time, f is the Coriolis parameter equal to 2Ω, where Ω is the angular velocity of the earth (i.e., the model is an f plane centered on the pole), ξ = (∂υ/∂x) − (∂u/∂y) is the relative vorticity, B = gh + (u2 + υ2)/2 is the Bernoulli potential, g′ is the reduced gravity of 0.02 m s−2, A is the lateral friction coefficient of 150 m2 s−1, τ is the ocean surface stress [i.e., τOceanSurface in Eq. (1)], ρ0 is the average density of 1026 kg m−3, h is the surface layer thickness, ⋅ (κh) is an advective term arising from the Gent–McWilliams eddy parameterization (Gent 2011), and κ is the eddy diffusivity. The equations are discretized on a C grid with a resolution of 15 km × 15 km.

Following the idealized approach of Spall (2003, 2004, 2013), the Canada Basin is represented as a circular domain approximately 1500 km in diameter, connected to a sponge region by a channel approximately 150 km wide (Fig. 2). The sponge region is constantly restored to the initial layer thickness over the period of a day, and is designed to absorb any waves propagating out of the domain, as well as allow an inflow or outflow transport to develop through the channel in response to changes in layer thickness inside the domain. In the vertical, the model consists of an active halocline layer overlying a stationary Atlantic layer. This approach is similar to that of Proshutinsky and Johnson (1997), who used a barotropic model to investigate the wind-driven dynamics of the upper layer of the Arctic Ocean. The halocline layer is initialized with a thickness of 400 m, matching the average depth of the 34.8 isohaline outside of the Beaufort Gyre in the Polar Science Center Hydrographic Climatology (PHC; Steele et al. 2001). The reduced gravity of 0.02 m s−2 is based on an average salinity of 32.3 in the halocline layer and 35.0 in the Atlantic layer from the PHC, resulting in an internal Rossby deformation radius of approximately 20 km.

Fig. 2.
Fig. 2.

Reduced gravity model of the Canada Basin with an active surface halocline layer of density ρ − Δρ initialized to be 400 m thick, overlying a stationary (u = 0) Atlantic layer of density ρ. The model domain is ~1500 km in diameter and is connected to a sponge region by a channel ~150 km wide. The sponge region is designed to absorb any waves propagating out of the domain, and to allow an inflow or outflow transport to develop in response to changes in layer thickness within the domain. The model is forced with an anticyclonic ocean surface stress centered over the domain. The magnitude of the curl of the stress field (which is proportional to the strength of the Ekman pumping) is at a maximum in the center of the domain, and is zero at the boundaries and in the outflow.

Citation: Journal of Climate 27, 21; 10.1175/JCLI-D-14-00090.1

The model is forced with an anticyclonic ocean surface stress [τOceanSurface in Eq. (1)] centered over the domain (Fig. 2). Along any diameter within the domain, the ocean surface stress in the x and y direction is described by
e4
e5
respectively, where θ is the angle that lines connecting each grid point with the center of the domain make with the positive x axis, and r is the radial distance between each grid point and the center of the domain (i.e., r2 = x2 + y2). Within the channel and sponge region, the stresses are described by
e6
e7
where x and y are the distances to each grid point from the center of the domain along the x and y axes, respectively, and C is a scale factor to ensure that the curl of the stress field is continuous at the domain/channel boundary. The stress fields are normalized between 0 and 1, and during each model run are multiplied by an idealized annual cycle in ocean surface stress to set the magnitude of the forcing. This stress formulation ensures that the curl of the ocean stress field reaches a maximum in the center of the domain, and decreases to zero at the boundaries and in the outflow region (Fig. 2).
To balance the input of vorticity from the winds (thus allowing the model to reach a steady state, or to spin down if the forcing is reduced to zero), the effect of eddies and diapycnal mixing have been incorporated into the model using the Gent–McWilliams parameterization (e.g., Gent 2011). Following the approach of Allison et al. (2011), the magnitude of the eddy diffusivity κ can be determined by considering the total transport T across any closed layer thickness contour around the domain. Ekman transport driven by the anticyclonic ocean surface stress will cause water to accumulate in the center of the domain, steepening the pressure gradient and driving an anticyclonic geostrophic current. Baroclinic instability associated with this current will result in an eddy-induced bolus transport toward the boundary of the domain. Consequently, the total transport across any closed thickness contour is the residual between the Ekman and eddy-induced transport velocities (T = TEkTEddies) and can be obtained by integrating the cross-contour components of the eddy (Eddies) and Ekman (Ek) transport velocities along the entire thickness contour:
e8
where λ is the distance in the azimuthal direction, τ(λ) is the ocean surface stress aligned parallel to the thickness contour, and r is the radial direction; it has been assumed that the thickness contours form concentric circles inside the domain. Integrating Eq. (8) between the boundary and the center of the domain and dividing by the radius of the domain gives
e9
where hc and hb are the layer thickness in the center and at the boundary of the domain respectively. If we assume that at steady state T = 0, τ(λ) = 0.02 N m−2 (discussed in more detail later) and hchb = 60 m based on the deepening of the 34.8 isohaline across the Beaufort Gyre in the PHC, then the eddy diffusivity (κ) must equal approximately 1300 m2 s−1.

While eddies are likely to be the first-order process responsible for balancing the Ekman pumping (Marshall et al. 2002), other processes such as lateral friction against the Chukchi Cap may play a role, and in our model their effect has been incorporated into the strength of the eddy flux. Consequently, it must be noted that the strength of the eddy flux may be an overestimate.

3. Control run

Initially the model is run for 40 years from rest in order to reach an idealized equilibrium state. The magnitude of the annual cycle in ocean surface stress used to force this control run (Fig. 3c) is based on the average conditions that existed between 1979 and 1990 (i.e., before the dramatic decline in Arctic sea ice cover; Stroeve et al. 2012). Figures 3a and 3b show the summer (4 months: July–October) and winter (8 months: November–June) averages of the mean sea level pressure from the monthly European Centre for Medium-Range Weather Forecasts (ECMWF) Interim Re-Analysis (ERA-Interim; Dee et al. 2011) and sea ice concentration from the Special Sensor Microwave Imager (SSM/I) satellite instrument (Maslanik and Stroeve 1999) over this period. The winds associated with the atmospheric Beaufort high are stronger and more anticyclonic during winter than during summer, and there is a decline in the sea ice concentration over the Beaufort Gyre during the four months of summer (although this seasonal change is much smaller than that seen in more recent observations). The results of Yang (2009) suggest that over this period the ocean surface stress is dominated by atmospheric forcing rather than by the sea ice concentration, and therefore the annual cycle in ocean surface stress used to force the control run reflects only the changing wind conditions, with 8 months of strong anticyclonic winds or high stress (0.023 75 N m−2) followed by 4 months of weak anticyclonic winds or low stress (0.0125 N m−2; Fig. 3c). Because of the difficulties in estimating the ocean surface stress directly from wind speed in the ice-covered Arctic, the value of the ocean surface stress is taken from the Nucleus for European Modelling of the Ocean (NEMO)–Louvain-la-Neuve Sea Ice Model (LIM) global coupled sea ice–ocean model, which has been used in previous Arctic studies (Lique et al. 2009, 2010; Lique and Steele 2012). Using a drag coefficient of 2.25 × 10−3 (Timmermann et al. 2005), the model annual mean ice–ocean stress of 0.018 N m−2 between 1979 and 1990 corresponds to an “effective” wind speed of 2.0 m s−1. This is lower than the ERA-Interim annual mean of 2.5 m s−1 over the same period, because of the sea ice cover reducing the transfer of momentum into the upper ocean.

Fig. 3.
Fig. 3.

Summer (July–October) and winter (November–June) averages of (a) mean sea level pressure from the ERA-Interim reanalysis and (b) sea ice concentration from the SSM/I satellite instrument over the Arctic Ocean between 1979 and 1990. (c) The annual cycle in ocean surface stress used to force the control run is based on the occurrence of stronger and more anticyclonic winds over the Beaufort Gyre during the 8 months of winter (green) compared to the four months of summer (red). The reduction in sea ice cover between July and October defines our summer period.

Citation: Journal of Climate 27, 21; 10.1175/JCLI-D-14-00090.1

Figure 4a shows the mean layer thickness and velocity fields over the last 5 years of the 40-yr control run, while Fig. 4b shows a cross section of the mean layer thickness through the center of the domain. As expected from the scaling of the eddy diffusivity, the halocline layer has thickened on average by 60 m across the domain, with the maximum thickening in the center. Around the boundaries and in the outflow region little thickening has occurred, which makes us confident that any forcing in the outflow region is not affecting the response within the circular domain. Associated with the deepening of the halocline layer is an anticyclonic geostrophic circulation centered over the domain. The time-averaged maximum velocity of approximately 2 cm s−1 is consistent with the climatological values reported in McPhee et al. (2009) and McPhee (2013). Figures 4c and 4d show the annual cycles in layer thickness in the center of the domain, and the transport through the channel over the last 5 years of the control run. The annual cycle in layer thickness, which can be interpreted as a proxy for the annual cycle in freshwater content, oscillates around a mean of 458.7 m and has a peak-to-peak amplitude of 3.4 m, consistent with the results of Proshutinsky et al. (2009). The annual cycle in transport through the channel has a peak-to-peak amplitude of 0.2 Sv (1 Sv ≡ 106 m3 s−1), oscillating around zero. The maximum freshwater content and maximum inflow into the domain occurs at the end of our winter period, whereas the minimum freshwater content and maximum outflow occurs at the end of our summer period, reflecting the integrated forcing throughout the year.

Fig. 4.
Fig. 4.

(a) Mean layer thickness (colors) and velocity field (vectors), (b) cross section of the mean layer thickness through the center of the domain (i.e., along the white line), (c) the annual cycle in layer thickness in the center of the domain (i.e., at the white dot), and (d) the annual cycle in transport through the channel (i.e., across the red line) all calculated over the last 5 years of the control run. Positive transports indicate an inflow and negative transports indicate an outflow. Integer year numbers correspond to the end of our summer period.

Citation: Journal of Climate 27, 21; 10.1175/JCLI-D-14-00090.1

Note that sensitivity analysis has shown that the results in the following sections are not dependent upon the exact magnitude of the annual cycle in ocean surface stress used to force the control run or on the realism of the steady state.

4. Experiment 1: Modified seasonality

Because the Arctic sea ice cover has begun to break up and retreat farther and for longer each year, the annual mean ocean surface stress (i.e., the net forcing) is increasing and also its seasonal cycle (i.e., the seasonality) is changing. Each may have an effect on the accumulation of freshwater in the Beaufort Gyre by triggering a different dynamical response, and therefore two sets of experiments were designed in order to investigate their effects separately.

In the first set of experiments, the effect of modifying the seasonality in ocean surface stress while holding the net forcing constant was investigated. Figures 5a and 5b show the idealized modifications made to the annual cycle in ocean surface stress. Over a number of model runs, the ocean surface stress during the 4 months of summer was incrementally increased with respect to the control run (Fig. 5a), idealistically representing an increased efficiency of momentum transfer into the ocean. For high levels of stress in summer, the period over which this stress was applied was then incrementally increased from 4 to 6 months (Fig. 5b), representing a longer melt season and therefore a longer period of enhanced momentum transfer. At the same time, however, the ocean surface stress in winter for each model run was reduced, to ensure that the net forcing remained constant at 0.02 N m−2 (gray dotted line in Figs. 5a and 5b). Each model run was initialized from the layer thickness and velocity fields at the end of the control run and then run for an additional 40 years. Note that, while we have used the terminology “summer” and “winter” here, our seasons are defined entirely in terms of periods of different ocean surface stress and need not relate to specific times of the year.

Fig. 5.
Fig. 5.

Modifications made to the annual cycle in ocean surface stress with respect to the control run (thick blue line) for (a),(b) the modified seasonality experiments and (c),(d) the modified net forcing experiments. In (a),(c), the ocean surface stress is incrementally increased in summer, representing the increased efficiency of momentum transfer. In (b),(d), the period over which high values of summer stress is applied is then incrementally increased, representing a longer melt season. For just the modified seasonality experiments in (a),(b), the stress in winter in each model run is reduced, to ensure that when integrated over the entire year the net forcing remains constant at 0.02 N m−2 (gray dashed line).

Citation: Journal of Climate 27, 21; 10.1175/JCLI-D-14-00090.1

Figure 6 shows the results of these experiments in terms of both the annual cycle in layer thickness in the center of the domain, and the transport through the channel over the last 5 years of each model run. The key result is that if the net forcing is held constant, the changing seasonality in ocean surface stress has no effect on the freshwater content of the Beaufort Gyre. The annual cycle in layer thickness (and thus the freshwater content) in each of the model runs oscillates around the control run mean of 458.7 m (Figs. 6b,e), and when the annual cycle in transport through the channel is integrated over the full 40 years of each model run, it is clear that there is no net inflow into the domain (Figs. 6c,f).

Fig. 6.
Fig. 6.

The annual cycle in (a),(d) ocean surface stress, (b),(e) the layer thickness in the center of the domain (white dot in Fig. 4a), and (c),(f) the transport through the channel (red line in Fig. 4a) from the last 5 years of each of the modified seasonality experiments. Positive transports indicate an inflow and negative transports indicate an outflow. Note that (a)–(c) are for model runs where the ocean surface stress in summer has been increased, while (d)–(f) are for model runs where the melt season length has been increased. Line styles and colors correspond to those used in Figs. 5a and 5b. The thick blue line represents the control run. Integer year numbers correspond to the end of our summer period.

Citation: Journal of Climate 27, 21; 10.1175/JCLI-D-14-00090.1

Nevertheless, by changing the seasonality in ocean surface stress we have significantly altered the shape of the annual cycle in both the layer thickness in the center of the domain, and the transport through the channel. With respect to the control run, as the ocean surface stress is steadily increased in summer (cf. the blue line with the dashed, dotted, and dot-dashed red lines in Fig. 6a), the phase of the annual cycle in both the layer thickness and the transport through the channel eventually reverses, as the stress in summer becomes greater than the stress in winter, and the amplitude increases, reflecting the increased amplitude of the annual cycle in ocean surface stress. The peak-to-peak amplitudes increase from 3.4 m and 0.2 Sv to a maximum (over these model runs) of 5.7 m and 0.3 Sv for the layer thickness in the center of the domain (dot-dashed red line in Fig. 6b) and the transport through the channel (dot-dashed red line in Fig. 6c), respectively.

When the length of the melt season is increased for high values of ocean surface stress in summer (cf. the blue line with the dot-dashed, dashed, and dotted green lines in Fig. 6d), the amplitudes of the annual cycle in layer thickness and the transport through the channel increase further to 8.6 m and 0.5 Sv, respectively (dotted lines in Figs. 6e and 6f). The timing of the minimum freshwater content also occurs earlier because of the shorter winter period, whereas the timing of maximum freshwater content remains unchanged.

5. Experiment 2: Modified net forcing

To explore the dynamical response of the Beaufort Gyre to an increased net forcing, and its effect on the freshwater content, a second set of experiments was performed. We applied the same modifications to the annual cycle in ocean surface stress as seen in the modified seasonality experiments (Figs. 5a,b) but, in this case, the ocean surface stress in winter was not reduced to offset the increase in summer (Figs. 5c,d). Consequently, the net forcing increased from the control run value of 0.02 N m−2 to a maximum of 0.028 N m−2. Each model run was again initialized from the layer thickness and velocity fields at the end of the control run and then integrated for an additional 40 years.

Figure 7a shows that the mean layer thickness in the center of the domain, averaged over the last 5 years of each model run, increases linearly with respect to the increased ocean surface stress. The linear relationship is described by the equation , where and are the increase in mean layer thickness and net forcing, respectively, from the control run. To estimate the magnitude of the corresponding freshwater accumulation, the annual cycle in transport through the channel can be integrated in time to produce a cumulative volume transport into the domain over the full 40 years of each model run. This cumulative volume transport can then be multiplied by a salinity anomaly:
e10
where Sref is the reference freshwater salinity and S is the in situ salinity, to calculate the total freshwater accumulation. As the model only solves the dynamical equations and thus there is no salinity variable, S has been set to a salinity of 27.7 in accordance with Giles et al. (2012). This value is based on the salinity measured in the mixed layer by an Ice Tethered Profiler deployed in the Canada Basin during April 2007 (Toole et al. 2010), and as such the freshwater anomaly estimate that results must be viewed as an upper bound. The reference value (Sref) equals 34.7 [again in accordance with Giles et al. (2012)]. Figure 7b shows the cumulative volume and freshwater transports into the domain as a function of both time and the increase in net forcing. All model runs show a net flux of freshwater into the domain with respect to the control run. The lack of seasonal variation when is approximately 4 × 10−3 N m−2 is due to the lack of seasonal variability in the forcing (i.e., the ocean surface stress in summer has been increased to a point where it equals the ocean surface stress in winter). Fitting an asymptotic exponential growth model to each of the model runs
e11
where is the cumulative freshwater transport into the domain as a function of time t, is the total freshwater accumulated in the domain for each model run, and k is the growth constant that reveals an e-folding time scale of approximately 3.3 yr for the accumulation of freshwater in the domain. This is slightly longer than the e-folding time scale of approximately 2.7 yr for the increase in the mean layer thickness in the center of the domain. As with the mean layer thickness in the center of the domain, the total accumulation of freshwater depends linearly on the increase in net forcing, and for the greatest increase in net forcing in our suite of model runs (0.008 N m−2), the total freshwater accumulated in the domain is approximately 4000 km3. The results of Martin et al. (2014) suggest that between 2000 and 2012, the Arctic-wide annual mean ocean surface stress has increased by 0.007 N m−2, which corresponds to an accumulation of freshwater in our model of approximately 3500 km3. This is around 50% of the increase observed by Giles et al. (2012) over broadly the same period.
Fig. 7.
Fig. 7.

Results of the modified net forcing experiments. (a) The increase in mean layer thickness in the center of the domain plotted as a function of the increase in net forcing. The blue dot represents the control run, the red dots represent model runs with an increased ocean surface stress in summer, and the green dots represent model runs with a longer melt season (as per the colors in Figs. 5c and 5d). (b) The cumulative volume and freshwater transports into the domain over the full 40 years of each model run as a function of the increase in net forcing.

Citation: Journal of Climate 27, 21; 10.1175/JCLI-D-14-00090.1

6. Discussion

Using a simple process model and idealized perturbations to the ocean surface stress, we have demonstrated that the dynamical response of the Beaufort Gyre to the decline in Arctic sea ice cover, and the related increase in the efficiency of momentum transfer into the upper ocean, is a plausible mechanism behind the accelerated accumulation of freshwater inferred by Giles et al. (2012). Despite the simple nature of our process model, in this section we will discuss some of the broader implications of our results and their relevance to the wider Arctic system.

a. Adjustment timescales

The results of both the modified seasonality experiments and the modified net forcing experiments shed some light on the time scales over which the Arctic Ocean adjusts to a change in forcing, as well as on which adjustment processes are responsible. Figures 6b and 6e show that over each season the layer thickness in the center of the domain continues to increase or decrease in response to the change in ocean surface stress, and never fully adjusts to a new equilibrium state. In contrast, Figs. 6c and 6f suggest that the magnitude of the transport through the channel adjusts more quickly to the seasonal changes in ocean surface stress, asymptoting toward a new equilibrium value over the period of each season. Consequently, it appears the processes that dominate the adjustment to a change in forcing around the boundary of the domain are very different from those that determine the adjustment in the center. The response time scale (e-folding time scale) of the transport through the channel is approximately 1 month, which is in good agreement with the time required for a boundary trapped Kelvin wave to propagate around the domain at a speed given by . This suggests that boundary waves are the dominant process behind the fast adjustment on the boundary. Waves propagating out of the sponge region (where we apply a strong restoring to the initial layer thickness of 400 m) effectively fix the layer thickness around the boundary in our model on time scales longer than 1 month.

Away from the boundaries, however, Kelvin waves have little effect. Here, the layer thickness as a function of time is determined by the balance between wind-induced Ekman pumping and the eddy-induced volume flux toward the boundary [i.e., Eq. (8)]:
e12
where the first term on the right-hand side represents wind-induced Ekman pumping and the second term represents the relaxation effect of the eddy flux. Scale analysis suggests that the relaxation time scale due to the eddy term is given by rABG/κλBG, where ABG and λBG are the area and circumference of the Beaufort Gyre respectively, and is approximately equal to 14 yr. As a result, we can conclude that eddies play a limited role in balancing Ekman pumping on the seasonal time scale. Consequently, the layer thickness in the center of the domain continually adjusts to the change in ocean surface stress over each season, and does not tend toward a new equilibrium state. Furthermore, this long adjustment time scale ensures that the annual cycle in layer thickness in the center of the domain (e.g., 3.4 m in the control run) is only a small fraction of the change in mean layer thickness across the Beaufort Gyre (60 m), as the annual cycle represents only the very beginning of a much longer-term adjustment to the changing ocean surface stress. Indeed, when the model is run from rest to a fully adjusted equilibrium state with a constant ocean surface stress of either 0.023 75 or 0.0125 N m−2 (i.e., the winter or summer value of ocean surface stress from the control run, respectively), the difference in equilibrium layer thickness between the two runs is approximately 33 m.

In the modified net forcing experiments, however, the eddy relaxation term eventually balances the wind-induced Ekman pumping, and it is the eddy relaxation time scale that determines the total quantity of freshwater that is accumulated in the domain (see Fig. 7b). This highlights the importance of eddies in the Arctic Ocean’s adjustment to a change in forcing, and provides a strong motivation to use eddy-resolving models when investigating the transient response of the Arctic Ocean to future changes in the forcing.

b. Effect of sea ice concentration

The results of both Martin et al. (2014) and Tsamados et al. (2014) show a long-term positive trend in the annually averaged momentum flux into the upper ocean in response to the decline in Arctic sea ice. However, Martin et al. (2014) find that the trend is not positive over all seasons. Instead, their model suggests that the efficiency of momentum transfer into the upper ocean is at an optimum when the sea ice concentration is approximately 80% (i.e., in fall and spring). Above this point, a thick and extensive sea ice cover damps the transfer of momentum due to the large internal ice stresses and shielding of the ocean from direct wind forcing. Below this point, the momentum transfer into the upper ocean decreases, as the drag associated with drifting sea ice is greater than that of open water. Consequently, between the 1980s and 2000s the ocean surface stress in summer decreases by 7% in their model, reflecting the increasing length of time that the sea ice concentration is below 80%. It must be noted however, that this negative trend in summer was not observed in the results of Tsamados et al. (2014). It is possible that the form drag parameterization used by Tsamados et al., which was not implemented in the model used by Martin et al., allows the ice concentration to decline further before the efficiency of momentum transfer decreases due to the fraction of open water. For example, the observations of Andreas et al. (2010) show that over summer sea ice and the marginal ice zone (where form drag is important), the drag coefficient peaks at an ice concentration of approximately 50%.

Irrespective of the exact details, the Arctic is currently undergoing a transition from the presence of a thick and extensive sea ice cover year round to a seasonally ice free state. If the concept of an optimal ice concentration for the efficiency of momentum transfer into the upper ocean holds, then the current positive trend in the momentum flux will begin to slow and eventually reverse, as the length of time when the sea ice concentration is below the optimum value increases (Martin et al. 2014). At this point, the accelerated accumulation of freshwater in the Beaufort Gyre will stop, and the excess freshwater that has built up will no longer be supported by Ekman pumping. It is unclear how long it will take for the positive trend in the momentum flux to reverse, and this may have important implications for the future evolution of the Arctic’s freshwater budget and for freshwater release into the North Atlantic.

Throughout this study, we have assumed in the construction of our idealized annual cycles that the momentum transfer will increase during summer. It should be noted, however, that the seasons have been arbitrarily defined, and the increased momentum transfer in summer could equally be interpreted as an increased momentum transfer in winter. Furthermore, given the results of the modified seasonality experiments, it is clear that whether the momentum transfer is increased in winter or summer is not important for the accelerated accumulation of freshwater in the Beaufort Gyre, as all that matters is the integrated net forcing over the entire year.

c. Relevance of our simple model

Despite having no effect on the freshwater content of the Beaufort Gyre, modifications made to the seasonality in ocean surface stress while the net forcing is held constant do impact the annual cycle in freshwater content and transport through the channel in our model (see Figs. 6b,c,e,f). Given the limitations of our simple process model (including, but not limited to, the parameterized effects of eddies and diapycnal mixing, and the lack of vertical resolution and a sea ice component), it is possible there are processes missing from our model that may be important in determining the response of the Beaufort Gyre to a change in the seasonality in ocean surface stress. To explore this further, we examine the driving mechanisms behind the annual cycle of freshwater content in the Beaufort Gyre using Arctic climatological datasets and data collected in the Beaufort Gyre.

Averaged over the Beaufort Gyre region, Fig. 8 shows the annual cycle in liquid freshwater content anomaly integrated between both the surface and the 34.8 isohaline (blue solid line) and between the surface and 25 m (blue dashed line) from the Monthly Isopycnal/Mixed-Layer Ocean Climatology (MIMOC; Schmidtko et al. 2013), the annual cycle in solid (ice) freshwater content anomaly (red line) from upward-looking sonar deployed as part of the Beaufort Gyre Exploration Project, and the magnitude of the annual cycle in Ekman pumping (black line) from the ERA-Interim reanalysis. We have used MIMOC instead of PHC because it resolves the annual cycle and has been updated with the most recent surface data collected by ice-tethered profilers. In the upper 25 m of the water column, the annual cycle in liquid freshwater content (blue dashed line) is dominated by thermal forcing (sea ice formation and melt; red line). It reaches a minimum in April–May because of brine rejection from sea ice formation, and a maximum in October–November because of sea ice melt. On the other hand, over the full depth range of the Beaufort Gyre (i.e., from the surface to the depth of the 34.8 isohaline), the annual variation in Ekman pumping (black line) appears to be the dominant forcing behind the annual cycle in total liquid freshwater content (solid blue line), with thermal forcing having little effect. The liquid freshwater content peaks in December–January due to stronger Ekman pumping during winter (although there appears to be a short time lag) and is at a minimum during August–September due to weaker Ekman pumping and a relaxation of the salinity field. An intermediary peak is observed in May, which is most likely associated with a seasonal reintensification of the atmospheric Beaufort high (Yang 2006, 2009).

Fig. 8.
Fig. 8.

The blue lines show the mean annual cycle in liquid freshwater content anomaly integrated between the surface and 25 m (dotted line) and between the surface and the depth of the 34.8 isohaline (solid line) over the Beaufort Gyre region from the MIMOC. The red line shows the mean annual cycle in the solid (ice) freshwater content anomaly from upward-looking sonar deployed as part of the Beaufort Gyre Exploration Project. The black line is the mean annual cycle in Ekman pumping velocity at the base of the Ekman layer over the Beaufort Gyre region, calculated from the wind stress curl. Positive velocities indicate a downwelling.

Citation: Journal of Climate 27, 21; 10.1175/JCLI-D-14-00090.1

Together, these results suggest that Ekman pumping is the dominant process in setting the annual cycle in liquid freshwater content within the Beaufort Gyre (although locally the balance may be different; Proshutinsky et al. 2009). Consequently, we can conclude that our simple process model does resolve the important dynamical processes behind the response of the Beaufort Gyre to a change in forcing, and that the decline in Arctic sea ice cover will affect the seasonality of the Beaufort Gyre and its freshwater content in a manner similar to the results of our model. We would expect the amplitude of the annual cycle in freshwater content (solid blue line in Fig. 8) to increase as the amplitude of the annual cycle in ocean surface stress increases, and the timing of the peaks in the annual cycle to change as the efficiency of momentum transfer increases and decreases at different times throughout the year.

d. Implications for the Arctic freshwater budget

In the modified seasonality experiments, the amplitude of the annual cycle in transport through the channel averages 0.3–0.4 Sv. Compared to the studies of de Steur et al. (2009) and Curry et al. (2014), this is smaller than the amplitude of the annual cycle in the volume export of surface layer waters through both Fram Strait (~4 Sv) and Davis Strait (~0.6–0.8 Sv). The small amplitude in our model directly reflects the limited annual variability seen in the layer thickness, which is both consistent with observational data (e.g., Proshutinsky et al. 2009) and is a direct consequence of the long time scale over which the model adjusts to the change in forcing. Consequently, while having some impact, the annual variability in the export of freshwater from the Beaufort Gyre cannot be the main driver behind the annual variability seen in the exports to either side of Greenland, and variability driven by local conditions is likely to be more important. For example, Lique et al. (2009) have shown that the variability in the liquid freshwater flux through Fram Strait depends upon variability in both the volume flux and the in situ salinity, which varies strongly because of ice formation and melt north of Greenland. Given the far-reaching consequences that changes in the export of freshwater to either side of Greenland may have on the circulation of the Atlantic Ocean, future studies should aim to quantify the contribution that the seasonal release of freshwater from the Beaufort Gyre has on the total freshwater export from the Arctic Ocean.

Independent of any accumulation of freshwater due to sea ice decline, the freshwater content of the Beaufort Gyre has been steadily increasing since the mid-1990s due to the large-scale anticyclonic wind regime (Proshutinsky et al. 2009). If, in the future, the wind regime becomes cyclonic, then the excess freshwater stored in the Beaufort Gyre will be released to the shelves (Häkkinen and Proshutinsky 2004; Condron et al. 2009; Jahn et al. 2010), where it may be exported into the North Atlantic (Stewart and Haine 2013). Such a release of freshwater has been proposed as a mechanism to explain some of the Great Salinity Anomaly (GSA) events, during which the subpolar North Atlantic and Nordic seas underwent decadal periods of freshening (e.g., Dickson et al. 1988; Belkin et al. 1998; Belkin 2004), with associated effects on the formation of North Atlantic Deep Water (Dickson et al. 2000; Jahn et al. 2010). Consequently, the accelerated accumulation of freshwater in the Beaufort Gyre may exacerbate the effects of a switch to a cyclonic wind regime, by making the Arctic even more anomalously fresh beforehand, and thus increasing the quantity of freshwater that may be exported into the North Atlantic.

e. Effect of changing meteorological conditions

Throughout this study, we have focused on the effect of changing sea ice conditions on the freshwater content of the Beaufort Gyre, without considering the potential impacts of changes in the wind field. However, as the strength of the atmospheric Beaufort high projects onto several different modes of atmospheric variability (Serreze and Barrett 2011; Mauritzen 2012), large-scale changes in atmospheric conditions may lead to a further accumulation or release of freshwater from the Beaufort Gyre that has not been accounted for in our model. For example, since the mid-1990s a negative trend in the phase of the Arctic Oscillation has strengthened the atmospheric Beaufort high in winter (i.e., the anticyclonic wind regime) and, more recently, a positive trend in the strength of the Arctic dipole between 2007 and 2012 may have acted to strengthen the Beaufort high in summer (Overland et al. 2012; Ogi and Wallace 2012). Furthermore, the study of Simmonds and Keay (2009) has shown that the strength of Arctic summer cyclones over the Eurasian Arctic has been increasing over the past 30 years, and if this were to continue, then the momentum flux into the ocean during summer may be reinforced by an enhanced wind forcing over the Beaufort Gyre region.

7. Conclusions

The overall aim of this study was to investigate the dynamical response of the Beaufort Gyre to the changing efficiency of momentum transfer associated with the decline in Arctic sea ice cover, and its link to the accelerated accumulation of freshwater observed by Giles et al. (2012). From two sets of experiments where the annual cycle in ocean surface stress was idealistically perturbed to reflect the effect of diminishing Arctic sea ice, we have shown that when the annual mean stress (i.e., the net forcing) is held constant, changes to its seasonal cycle (i.e., the seasonality) have no net effect on the freshwater content of the Beaufort Gyre in our model. We find that the Beaufort Gyre continually adjusts to the seasonal changes in ocean surface stress since the adjustment time scale, which is set by eddies, is considerably longer than seasonal.

On the other hand, we find a linear relationship between the increase in the annual mean ocean surface stress and the amount of freshwater accumulated in the Beaufort Gyre, regardless of the seasonal cycle in the forcing. This linear relationship is similar to that found by Proshutinsky et al. (2002, 2009) between the freshwater content of the Beaufort Gyre and the wind stress curl over the region, and it suggests that as the Arctic sea ice cover continues to decline, the linear accumulation of freshwater in the Beaufort Gyre for a given wind stress curl will increase (e.g., Giles et al. 2012), because of the enhancement in the magnitude of the stress that is transferred from the surface of the ice pack to the ocean below. This confirms that the increased efficiency of momentum transfer into the upper ocean associated with the decline in Arctic sea ice cover is a plausible mechanism to explain the accelerated accumulation of freshwater in the Beaufort Gyre. The total quantity of freshwater accumulated for a given change in ocean surface stress depends on the eddy diffusivity (i.e., on the length of time it takes for the eddy field to balance the change in Ekman pumping). The key dynamical process at play is as follows: as Arctic sea ice is becoming weaker, thinner, and more broken up, the annually averaged momentum flux into the upper ocean is increasing for the same wind speed, resulting in an accelerated linear accumulation of freshwater through an enhanced mechanical deformation of the salinity field.

Our results have implications for both the net amount of freshwater stored in the Beaufort Gyre and the annual cycle in its freshwater content. This may have an effect on the annual cycle in freshwater export from the Arctic Ocean, and future observational studies aimed at detecting changes in the seasonality of the Beaufort Gyre freshwater content would be worthwhile. At the same time, our results have highlighted the importance of eddies in the transient response of the Arctic to a change in forcing, and more should be done to investigate the role they play using eddy resolving models.

Acknowledgments

We thank Michel Tsamados, Daniel Feltham, and David Marshall for useful discussions during this study. The upward-looking sonar data were collected and made available by the Beaufort Gyre Exploration Project based at the Woods Hole Oceanographic Institution (http://www.whoi.edu/beaufortgyre) in collaboration with researchers from Fisheries and Oceans Canada at the Institute of Ocean Sciences. The comments and suggestions of three anonymous reviewers are also gratefully acknowledged. This work was funded by the Natural Environment Research Council (NERC).

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