North Atlantic response to the above-normal export of sea ice from the Arctic

Transcription

1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 108, NO. C7, 3224, doi: /2001jc001166, 2003 North Atlantic response to the above-normal export of sea ice from the Arctic Oleg A. Saenko, Edward C. Wiebe, and Andrew J. Weaver School of Earth and Ocean Sciences, University of Victoria, Victoria, British Columbia, Canada Received 9 October 2001; revised 24 April 2002; accepted 2 July 2002; published 11 July [1] The response of the thermohaline circulation (THC), as well as the freshwater and heat budgets of the northern North Atlantic, to above-normal sea ice export from the Arctic is examined using a global model. The model is not constrained by either open boundary conditions or prescribed atmospheric air temperature and humidity. Two sets of experiments are presented: the transient and the persistent above-normal ice export. In the transient case, ice export is increased by a factor of 2 for 1 5 years. Our century-long simulations do not support the notion that the simulated climate may switch to a new quasiequilibrium under such perturbations. Rather, in response to the transient positive ice export anomalies the overturning circulation first slows down but then almost completely recovers years after the perturbation is removed. However, the budgets of freshwater and heat continue to evolve for up to 40 years in this case. When the simulated North Atlantic freshening reaches a magnitude comparable to that during the Great Salinity Anomaly (GSA), the strength of overturning and heat transport from subtropical to subpolar North Atlantic reduce by no more then 5%. We show that in order to generate a previously reported decrease of overturning and heat transport of as much as 20% the doubled ice export must be sustained for at least 5 years. This would result in a North Atlantic freshening more than 3 times larger than that estimated for the GSA event. In the case of a persistent above-normal export of sea ice from the Arctic the THC also does not collapse, at least within the range of the ice export increase (1.5 3 times) used here. Rather, after about years the overturning begins to recover. The stability of the overturning to the persistent above-normal export of sea ice from the Arctic appears to be due to two processes operating on a decadal timescale of years. First, the internal (to the coupled system) redistribution of freshwater between the Arctic and North Atlantic, associated with the enhanced export of sea ice, makes the North Atlantic fresher and Arctic Ocean saltier. This, if persistent, decreases the amount of freshwater leaving the Arctic toward the North Atlantic in a liquid form. Second, because the overturning circulation does not collapse, the freshwater anomaly propagates downward within the region of deep water formation, removing the excess of buoyancy from the surface ocean. Also, the use of an active atmospheric component is important for stabilizing the overturning circulation. INDEX TERMS: 4215 Oceanography: General: Climate and interannual variability (3309); 4540 Oceanography: Physical: Ice mechanics and air/sea/ice exchange processes; 4207 Oceanography: General: Arctic and Antarctic oceanography; 4255 Oceanography: General: Numerical modeling; KEYWORDS: North Atlantic, Arctic, sea ice, ocean Citation: Saenko, O. A., E. C. Wiebe, and A. J. Weaver, North Atlantic response to the above-normal export of sea ice from the Arctic, J. Geophys. Res., 108(C7), 3224, doi: /2001jc001166, Introduction [2] The export of sea ice from the Arctic has long been recognized to play a role in the North Atlantic climate and deep water formation [Aagaard and Carmack, 1989]. The freshwater input to the ocean, resulting from sea ice melting, influences both the North Atlantic Ocean and the atmosphere through its impact on the rate of deep water Copyright 2003 by the American Geophysical Union /03/2001JC001166$09.00 production and on the associated heat transport and heat loss to the overlying atmosphere. It is thus of interest to quantify the connection between the amount of sea ice exported from the Arctic and the associated rate of deep water production in the North Atlantic. [3] The export of sea ice through Fram Strait exhibits strong interannual variability [e.g., Kwok and Rothrock, 1999; Hilmer et al., 1998]. Years of anomalous ice export are generally attributed to anomalies in atmospheric circulation [Serreze et al., 1992]. A modeling study by Hakkinen [1993] suggests that the Great Salinity Anomaly (GSA), 17-1

2 17-2 SAENKO ET AL.: NORTH ATLANTIC RESPONSE TO ARCTIC SEA ICE observed in the northern North Atlantic during [Dickson et al., 1988], could have been a result of an anomalously large export of ice from the Arctic, particularly in [4] Several recent modeling studies have examined the response of the North Atlantic thermohaline circulation (THC) to variations of ice export from the Arctic. For example, Hakkinen [1999] simulated a GSA-like event by forcing an ice-ocean model with the 1968 wind stress. The forcing was applied during the first year of model integration; for the second year and all following 10 or 15 years (depending on experiment) the wind stress was set to climatology. Using a limited area model with open boundaries, Hakkinen [1999] found that a single, 1 year long perturbation resulted in a reduction of meridional heat transport and overturning circulation by as much as 20% 5 years after the perturbation was applied. [5] In the present study we employ a global ocean-iceatmosphere model, thereby allowing us to analyze the model results many years after the applied perturbation. This is important because as we will show, the GSA event might have been felt in the deep Atlantic well beyond 1.5 decades. Another important difference between our model and that of Hakkinen [1999] is that we incorporate an interactive atmospheric component. Among other feedbacks (such as internally predicted precipitation, river runoff, etc.) the use of an interactive atmosphere means that the model air temperature and humidity are free to adjust and feed back upon the changes of North Atlantic sea surface temperature (SST) because of the extra ice inflow. This feedback has been shown to be an important factor in stabilizing the THC [e.g., Zhang et al., 1993; Rahmstorf and Willebrand, 1995]. [6] One of our objectives here is to investigate further the response of the simulated coupled climate to transient, pulse-like perturbations of ice export from the Arctic. One of our experiments is similar to those considered by Hakkinen [1999]. Specifically, we increase ice export (by increasing wind stress) from the Arctic for only 1 year and then set it back to climatology. Increasing the ice export by a factor of 2 in this case results in a freshening of the North Atlantic comparable in magnitude to the GSA event. However, it is also of interest to consider whether the THC would be able to survive larger freshwater anomalies than occurred during the GSA. This is done here by sustaining the above-normal ice export for 3 and 5 years before setting it back to normal. Although the main focus here is on the above-normal ice export from the Arctic, we also consider a response to below-normal ice export. [7] Another question we address is whether the North Atlantic THC is able to survive large persistent positive anomalies of ice export from the Arctic beyond several decades? If so, what are the negative feedbacks operating within the coupled system that tend to bring the THC back to its active state? Similar questions but on the timescale of no more than 1.5 decades have been addressed by Mauritzen and Hakkinen [1997]. They showed that their iceocean model produced a 2 3 Sv increase of the meridional circulation cell at 25 N after a decrease of sea ice export from the Arctic of 800 km 3 yr 1, sustained for 15 years. Here we apply persistent perturbations of larger magnitudes, ranging from 0.03 Sv (950 km 3 yr 1 ) to 0.1 Sv (3150 km 3 yr 1 ) above the model equilibrium value of Sv. Through our use of longer integration than was conducted by Mauritzen and Hakkinen [1997] we show that the timescale of years appears to be critical for the THC to pass its weakest state arising from a persistent above-normal inflow of sea ice from the Arctic and to begin recovering. [8] The rest of the paper is divided into four sections. We describe the global coupled model and the design of the numerical experiments in the following section. The model control simulation is analyzed, emphasizing a region of the North Atlantic where the model simulates most of the deep water formation in section 3. Results of the sensitivity experiments to the transient and persistent ice export anomalies are presented in subsections 4.1 and 4.2, respectively, of section 4. A discussion and conclusions follow in section Coupled Model and Experimental Design [9] In this work we use a climate model of intermediate complexity [Weaver et al., 2001]. The basic model components include the Geophysical Fluid Dynamics Laboratory (GFDL) (MOM) version 2 [Pacanowski, 1995], an updated version of a two-dimensional (2-D) energy-moisture balance model [Weaver et al., 2001], an energy-conserving thermodynamic sea ice model with a sub-grid-scale ice thickness distribution allowing for six ice categories within each model grid cell [Bitz et al., 2001], and an elasticviscous-plastic representation of sea ice dynamics [Hunke and Dukowicz, 1997]. The static stability criteria of ocean columns is checked separately under each ice category within a grid cell [Saenko et al., 2002]. [10] All model components employ a rotated grid geometry (see Weaver et al. [2001] for details) with horizontal resolution of in longitude and latitude, respectively. There are 19 vertical levels in the ocean model and up to 8 layers in the sea ice model. The ocean model uses isopycnal mixing after Gent and McWilliams [1990] with the coefficients of both thickness and isopycnal diffusivities set to cm 2 s 1 and with mixing surfaces limited to a maximum slope of The horizontal and vertical viscosities are set to and 10 cm 2 s 1, respectively. The coefficient of vertical diffusivity is set to 0.8 cm 2 s 1 ; the use of this relatively high value is dictated by a wellknown sensitivity of poleward heat transport and meridional overturning circulation to vertical diffusivity in low-resolution ocean models [e.g., Bryan, 1987]. The atmospheric model diffusively transports heat and moisture. Precipitation occurs when the relative humidity exceeds a threehold value of 85%. [11] Our model, like most other global climate models, does not resolve small-scale flows across the Greenland- Iceland-Faroes rise, which are important for proper simulation of North Atlantic Deep Water (NADW) formation. In order to account partly for the effect of these flows, as well as to improve the transport of heat in this ocean area, we removed the unresolved island of Iceland. Furthermore, flow through Bering Strait is also not resolved by the model, and here it is closed. [12] In the model version we use here the external constraints to the coupled system are the solar radiation at

3 SAENKO ET AL.: NORTH ATLANTIC RESPONSE TO ARCTIC SEA ICE 17-3 Figure 1. The region of analysis in the northern North Atlantic (NA region) and the average ventilation depth. The model grid has been rotated so that the north pole lies within Greenland. the top of the atmosphere and the wind stress. The wind stress data are taken from National Centers for Environmental Prediction/National Center for Atmospheric Research (NCEP/NCAR) reanalysis [Kalnay et al., 1996] averaged over the period to form an annual cycle from the monthly fields. Other than that, the model is fully interactive, with no restoring to any kind of temperature and/or salinity field. [13] The model was initially spun up for 2000 years. The final year of this spin-up is further referred to as the control state. The perturbation experiments were initialized from this control state. We mainly focus on two sets of experiments. In the first one the sea ice export from the Arctic was approximately doubled in a pulse-like manner for 1 (EXP1), 3 (EXP1a), and 5 (EXP1b) years and then set back to normal (control). In the second set of experiments the ice export was increased by approximately a factor of 2 (EXP2), 1.5 (EXP2a), and 3 (EXP2b) and was kept at that level during the model integration. Because the large-scale wind pattern over the Arctic favors an outflow of sea ice to the North Atlantic, we perturbed the wind stress over the Arctic sea ice cover in order to enhance the sea ice export to the North Atlantic. For example, in order approximately to double the export of ice for 1 year the wind stress over the Arctic sea ice cover was doubled for 1 year. The wind stress perturbation was applied to both the ice cover and the open ocean, provided that a grid cell in the model had at least 10% of sea ice concentration. The surface turbulent heat fluxes were not explicitly perturbed. [14] In addition, two experiments have been performed in which Arctic ice export was reduced for 1 year by about 50% (EXP3) and 100% (EXP3a) and then set back to normal. 3. Control Simulation [15] Here we discuss the control state of the coupled model. Because the main focus of this study is on the sensitivity of the meridional circulation and freshwater and heat budgets of the North Atlantic Ocean to the inflow of extra sea ice, we concentrate on the relevant diagnostics. In the next section some of these diagnostics are used as reference values for the sensitivity analysis. [16] The region of the North Atlantic that we analyze (hereinafter referred to as NA region or NA box) is shown in Figure 1. The NA region communicates with the rest of the world through sections A, B, and C (Figure 1) and through the surface of km 2. It is connected to the Arctic Ocean through sections A and B. Section A represents the approximate location of the Greenland-Scotland Ridge, with maximum depth of 1128 m allowed in the model. Section B connects the NA region with the Arctic Ocean through the channel, representing a combined flow through a number of shallow passages of the Canadian Archipelago. The annual average volume transport of water through the channel is 0.9 Sv, which is between the estimates of 0.5 and 1.7 Sv made by Steele et al. [1996] and Fissel et al. [1989], respectively, for the Canadian Archipelago; Timofeyev [1956] estimated this transport to be about 1 Sv. Section C connects the NA region to subtropical ocean. [17] In the NA region the model simulates most of the NADW formation. Spatial structure of the maximum depth of convection and vertical profile of downwelling in the region are shown in Figure 1 and Figure 2, respectively. The ventilation of the North Atlantic in the model occurs somewhat too far to the south, which is the case in many coarse-resolution coupled models. Very little deep water forms in the Greenland-Iceland-Norwegian (GIN) Sea region, which is ice covered in the winter (see also Holland et al. [2001] for more details). The maximum downwelling in the NA region reaches 11.5 Sv at about 1000 m (Figure 2). The maximum annual average rate of the North Atlantic meridional overturning (the overturning index), which indicates the strength of the THC, is 14.8 Sv. However, the depth of NADW overturn is somewhat underestimated by the model, which has also been found in other coarseresolution models employing the Gent and McWilliams [1990] mixing scheme [e.g., England and Rahmstorf, 1999]. [18] The annual mean rate of ice export from the Arctic, simulated by the model approximately at the latitude of the Fram Strait, is about Sv. This value is close to the estimate of Sv made by Hakkinen [1993] on the basis of a regional Arctic ice-ocean model. However, it is Figure 2. The vertical profile of net downwelling in Sv (1 Sv = 10 6 m 3 s 1 ) in the NA region.

4 17-4 SAENKO ET AL.: NORTH ATLANTIC RESPONSE TO ARCTIC SEA ICE Table 1. Freshwater (FW) and Heat Contents of the Different Layers of the North Atlantic Box a Layer FW Content, m 3 Heat Content, J Surface 228 m m bottom Total a See Figure 1 for the box location. The contents of FW and heat are calculated relative to salinity of 35 and temperature of 0 C, respectively. Table 2. Freshwater (FW) and Heat Budgets of the NA Region a Interface FW Fluxes, Sv Heat Fluxes, TW Surface Section A Section B Section C Net < a See in Figure 1. Positive transport of FW or heat are directed into the NA region. The lateral transports of FW and heat are calculated relative to salinity of 35 and temperature of 0 C, respectively (1 Sv = 10 6 m 3 s 1 ;1 TW = W). somewhat less than the observational estimate of Sv, recently reported by Kwok and Rothrock [1999], although it is greater than the value of 0.04 Sv calculated by Nazarov [1938]. [19] Table 1 summerizes the contents of freshwater and heat for different layers of the NA region. We will consider three layers in the NA box: the upper (0 228 m), the middle ( m), and the deep (1128 m-bottom) layers. This division is somewhat arbitrary. One of our intentions in section 4 is to track down the propagation of the freshwater anomalies within the NA box, as well as to separate the impact of ice export upon lateral transports of heat and freshwater in the upper, middle, and deep layers of the region. The contents of freshwater (FW) and heat (H) are defined as follows: Z S r S FW ¼ dv v S r ð1þ Z H ¼ C p r ðt T r ÞdV; ð2þ v where integration is done over the volume (V) of the NA region; C p and r are the heat capacity and density of water; and S and T are the model-simulated salinity and temperature, whereas S r and T r are the reference salinity and temperature. The latter were set to 35 and 0 C, respectively. The same reference values for salinity and temperature were used for calculating the transports of freshwater and heat, discussed below. [20] Table 2 gives a summary of annual mean budgets of freshwater and heat in the water column of the NA box. It should be noted that the magnitude of individual fluxes depends on the reference values. However, for any chosen reference the net divergence of horizontal fluxes in a water column must balance at equilibrium the associated flux through the surface. The net values in Table 2 indicate that the control state is not exactly in equilibrium, though the drift is small. Figure 3 provides a detailed look at the fluxes of freshwater and heat across different interfaces of the NA box. Again, the magnitudes of individual fluxes are given relative to the chosen reference values for the salinity and temperature. [21] Under these assumptions the following simple analysis can be done to the budgets shown in Figure 3 and in Table 2. For the whole water column of the NA box the positive freshwater fluxes through the surface and from the Arctic through sections A and B (see Figure 1) are balanced by the negative freshwater flux across section C, mainly because of the inflow of saline subtropical water (Figure 4a). The (liquid) freshwater flux of Sv from the Arctic into the upper layer consists of and Sv across sections A and B, respectively. The positive freshwater flux through the surface is not balanced by the lateral divergence of fluxes in the upper layer (Figure 3), with the rest of the fresh water moving into the middle layer. In the middle layer the freshwater flux from above is partly balanced by the lateral divergence, with the rest moving into the deep layer. The lateral flux of Sv into the middle layer consists almost entirely of that through the section A, without noticeable contribution from section B. In the deep layer the rest of the freshwater leaves the NA box mainly with the NADW (see Figures 3 and 4b). [22] The major components of the NA region heat budget include the transport of heat from subtropics with thermocline waters in the upper and middle layers (Figure 3; see also Figure 5a), which is approximately balanced by the heat losses through the surface, flow toward the Arctic, and the outflow of relatively warm NADW in the deep layer (Figure 3; see also Figure 5b). The heat flux into the middle box from the subtropics is mainly compensated by the vertical fluxes toward the layers above and below Figure 3. Annual mean budgets of (left) freshwater and (right) heat for the NA box after the 2000 year model spin-up. A + B denotes a combined flux across sections A and B (see Figure 1), whereas C denotes the flux through section C. The numbers 1, 2, and 3 denote the upper layer (0 228 m), the middle layer ( m), and the deep layer (1128 m-bottom). The lateral advective fluxes of freshwater and heat are calculated relative to the reference salinity of 35 and temperature of 0 C, respectively. The directions of internal fluxes are shown with dashed arrows. The solid arrows represent external fluxes, with numbers inside in sverdrups (freshwater) and in terawatts (heat) (1 Sv = 10 6 m 3 s 1 ;1TW=10 12 W).

5 SAENKO ET AL.: NORTH ATLANTIC RESPONSE TO ARCTIC SEA ICE 17-5 Figure 4. The salinity and velocity fields at (a) 177 and (b) 2148 m. In Figures 4a and 4b the contour interval for salinity is 0.1 and 0.01, respectively. Note the rotated grid geometry. Figure 5. The temperature field at (a) 177 and (b) 2148 m. In Figures 5a and 5b the contour interval is 1 and 0.1 C, respectively. Note the rotated grid geometry.

6 17-6 SAENKO ET AL.: NORTH ATLANTIC RESPONSE TO ARCTIC SEA ICE Figure 6. Time series of anomalies of (a) ice export through the Fram Strait-Barents Sea region and (b) overturning index in EXP1 (solid line), EXP1a (dashed line), and EXP1b (dash-dotted line). (Figure 3). In section 4 the sensitivity analysis is performed relative to the control state discussed above. 4. Results of the Sensitivity Experiments [23] In this section we present the results of the sensitivity experiments, concentrating mainly on EXP1 and EXP2. Hereinafter the response or anomaly mean a deviation of a diagnostic from its control state value, discussed in section 3. We begin our analysis with the case of the pulse-like (or transient) increase of sea ice export from the Arctic to the North Atlantic, which is followed by the case of persistent enhanced ice export. Experiments EXP3 and EXP3a, in which ice export is reduced, are also discussed Pulse-Like Increase of Ice Export [24] The responses of sea ice export and the overturning index in EXP1, EXP1a, and EXP1b are shown in Figure 6. The time series of ice export in each of the experiments closely follow the perturbation applied to the wind stress: doubling the wind stress over the Arctic ice for a specified period of time approximately doubles the rate of ice export for approximately the same period of time (Figure 6a). The increased freshwater fluxes, resulting from above-normal sea ice export from the Arctic, increase the buoyancy of North Atlantic near-surface waters and weaken the THC (Figure 6b). However, the THC does not collapse even when the doubled above-normal ice export is sustained for 5 years (EXP1b). Depending on the net inflow of extra sea ice, the overturning index decreases by about 5 (EXP1), 11 (EXP1a), and 15% (EXP1b). [25] The weakening of the THC is followed by its recovery. Two timescales can be identified, associated with the THC response. First, the minimum overturning strength lags the return of ice export to its control value by about 5 6 years (Figures 6a and 6b), consistent with results by Hakkinen [1999]. Holland et al. [2001] also found the largest anticorrelation between the ice export and overturning when the former leads overturning time series by 4 5 years. This timescale is associated with the initial capping of the NA region by the fresh water (see Figures 8a and 8c) due to sea ice melt. Second, the overturning index reaches its control value and even slightly overshoots it after about years from when the perturbation is removed or after about years after the perturbation is applied (Figure 6b). Holland et al. [2001] found enhanced spectral power of the THC variability at decadal timescales concentrated at approximately 20 years as a response to the stochastically forced wind stress over the Arctic sea ice cover. However, it was not clear what set the dominant 20 year timescale. As we will show, this decadal timescale seems to be associated with the downward propagation of the freshwater anomaly due to the still strong THC. [26] It takes years for the net freshwater contents of the NA box to return to their control values (Figure 7); then the contents drop somewhat below their control values. At its maximum the net freshening of the NA region is approximately 1500, 4200, and 7150 km 3 in EXP1, EXP1a, and EXP1b, respectively. For EXP1 the freshening at its maximum is of comparable magnitude to that passed through the Labrador Sea during the GSA event, which is estimated to be about 2200 km 3 [Dickson et al., 1988]. [27] In order to understand better how the North Atlantic might be affected by the inflow of extra ice we apply the freshwater/heat budget analysis. For the remainder of this subsection we mainly discuss EXP1. Similar conclusions can be drawn from the analysis of EXP1a and EXP1b. [28] The effect produced by the inflow of extra ice to the different components of freshwater and heat budgets of the NA region is shown in Figure 8. There are several important processes that follow the enhanced large-scale divergence of Arctic sea ice. The upper North Atlantic Ocean freshens, whereas the upper Arctic Ocean becomes saltier (Figure 8a). Figure 7. Time series of anomalies of freshwater content of the NA box in EXP1 (solid line), EXP1a (dashed line), and EXP1b (dotted line).

7 SAENKO ET AL.: NORTH ATLANTIC RESPONSE TO ARCTIC SEA ICE 17-7 Figure 8. Time series of anomalies of (a) freshwater (FW) content, (b) heat content, (c) transports of FW, and (d) transports of heat for the NA box in EXP1. In Figures 8a and 8b the solid, thin dashed, dash-dotted, and dotted lines denote the contents within the whole NA box, the upper 228 m, from 228 to 1128 m, and below 1128 m, respectively; in Figure 8a the heavy dashed line denotes the response of FW content of the Arctic Ocean above 228 m. In Figures 8c and 8d the solid, dashed, dash-dotted, and dotted lines denote the transports at the air-sea interface, through passages A, B, and C (see Figure 1), respectively. The latter is not spatially uniform (Figure 9a). One of the areas of large salinification is over the Arctic shelves, where the intensified ice motion reduces ice thickness (Figure 9c) and increases ice production. To the north of the Canada- Greenland region, on the other hand, ice production decreases, whereas ice convergence increases, making ice thicker. As we will show in section 4.2, the Arctic Ocean salinification, if persistent, imposes an important constraint on the net (liquid plus frozen) export of freshwater from the Arctic to the North Atlantic. [29] It takes approximately 5 years for the freshwater anomaly to cap the NA region. The major contributor to the upper ocean freshening is the flux through the surface (Figure 8c), which is almost totally due to ice melt; the change in precipitation minus evaporation accounts for less than 3% of the flux response. The initial rapid freshening of the upper ocean reduces the convection activity in the region, making the North Atlantic Ocean warmer at middle depths (Figure 8b). However, the freshwater content of the middle and deep layers is not much affected initially by the reduced convection because the salinity in the NA region has a relatively uniform vertical structure (not shown). Rather, the freshwater anomaly begins to propagate downward slowly (Figure 8a) with reduced, but not collapsed, meridional overturning (Figure 6b). This decreases the upper ocean freshwater content and makes first the middle and then the deep ocean fresher (Figure 8a). The downward propagation of the freshwater anomaly, as well as the increased heat content of the subsurface ocean (Figure 8b), tends to destabilize the North Atlantic water column and accelerates the THC relaxation back to its control state by years. In fact, as we will show in section 4.2, the system is able to return to a steady state with active THC even when the enhanced ice flux persists. [30] The fluxes of heat across different interfaces also exhibit a response to the changes of the North Atlantic freshwater budget, followed by their recovery to the control state (Figure 8d). The heat lost through the surface reaches its minimum at year 4 (note that the fluxes out of the NA box are negative so that the decrease (increase) of a negative flux means positive (negative) flux anomaly). The heat transport from the subtropics first decreases and then recovers at around year 17, closely following the THC. The correlation (at zero lag) between the response of heat transported to the NA region through section C and the overturning index is At its extreme low level the transport of heat from subtropics drops by about 6.5% in EXP1. Similar to the meridional overturning response, the impact of the ice export increases when its above-normal rate is sustained for longer. More specifically, in EXP1a and EXP1b the reduction of heat transport from the subtropics reaches 14 and 20%, respectively. The relationship between the overturning index and the heat flux across section C is shown in Figure 10 for each of the experiments. Our model

8 17-8 SAENKO ET AL.: NORTH ATLANTIC RESPONSE TO ARCTIC SEA ICE Figure 9. Annual mean anomalies of (a) salinity at 88 m depth, (b) near-surface air temperature, (c) ice thickness, all for EXP1; (d) annual mean anomaly of ice thickness in EXP3a. The salinity anomaly corresponds to year 5, whereas all the rest of the anomalies correspond to year 1 since the associated perturbation is set to zero. thus shows a weaker response than that used by Hakkinen [1999] to the GSA-like event. In order to generate in our model a reduction of the THC and heat transport by as much as 20%, reported by Hakkinen [1999], we had to sustain the 2 times above-normal export of ice from the Arctic for at least 5 years. This results in a freshening of the North Atlantic several times larger than that estimated for the GSA event. [31] The response of surface freshwater flux in Figure 8c attenuates the responses of the lateral fluxes. It is, however, of interest to take a closer look at the response of the lateral fluxes. At the initial stage of the response (up to about year 4) the (liquid) exchange with the Arctic (particularly through section A) tends to freshen the NA box, which is almost compensated by the tendency to make the NA box saltier across section C (Figure 11). The opposite occurs from year 4 to about year 14. During this time the outflow of freshwater from the Arctic decreases (particularly through section B, which is accompanied by the decrease of salinification through section C). The former is a result

9 SAENKO ET AL.: NORTH ATLANTIC RESPONSE TO ARCTIC SEA ICE 17-9 Figure 10. The evolution of heat transport across section C versus the overturning index in EXP1 (solid line), EXP1a (dashed line), and EXP1b (dotted line). Arrows show direction of time. of the above mentioned Arctic Ocean salinification. At around year 13 the downward propagation of the freshwater anomaly forces the outflowing NADW to be involved in the process of freshwater removal from the NA box (Figure 11b). [32] It thus takes up to decades for the freshwater signal to propagate from the surface downward to deeper layers within the NA box (Figure 8a), where it is being gradually removed with the deep flow (Figure 11b). About the same time is needed for the North Atlantic THC to recover back to its control state. However, the budgets of freshwater and heat in the subpolar North Atlantic continue to evolve for up to 40 years after the GSA-like perturbation (Figures 11a and 11b). [33] The use of an active atmospheric component in our model has an important effect on the THC recovery by reducing the feedback between the overturning circulation and heat lost to the atmosphere [Zhang et al., 1993]. More specifically, in response to the weakened THC and the reduced northward oceanic heat transport, both the North Atlantic SST (not shown) and the air temperature (Figure 9b) decrease. This causes the rate of heat lost to the atmosphere in the northern North Atlantic to be less affected by the THC weakening, increasing the THC s stability. It has been shown by Zhang et al. [1993] that the THC is less stable and is more sensitive to freshwater perturbations under a fixed, not interactive atmosphere. This feedback between the oceanic circulation, SST, and air temperature also likely plays an important role in preventing the THC from collapsing in the simulations with persistent above-normal export of ice (discussed in section 4.2). [34] In the Arctic the simulated ocean loses heat at a rate of 2.0 W m 2 on average, which is close to observed estimates. In a response to the pulse-like increase of ice export the heat loss increases by about 10, 25, and 45% in EXP1, EXP1a and EXP1b, respectively. The increased heat loss tends to warm up slightly the overlaying air (Figure 9b). [35] Although our main object here is to investigate a response to above-normal ice export, we also performed experiments with below-normal ice export from the Arctic. The reduction of ice inflow from the Arctic (EXP3 and EXP3a) has an opposite but less pronounced effect on the strength of THC (Figure 12). Both the SST and the air temperature increase (not shown), again working to reduce the response of the THC. The spatial structure of the ice thickness response is somewhat opposed to the case of the positive ice export anomaly (Figures 9c and 9d), with a major thickness reduction to the north of Greenland and its increase in the rest of the Arctic. Both fields, however, show a dipole-like structure. The response of ice thickness is largest just after the perturbation and then rapidly reduces when the perturbation is set to zero. The response of the oceanic heat loss in the central Arctic to the reduction of ice export is small. In section 4.2 we discuss the North Atlantic response to the persistent above-normal export of sea ice from the Arctic A Persistent Enhanced Ice Export [36] The response of the North Atlantic THC to the persistent anomalous export of sea ice from the Arctic was previously considered by Mauritzen and Hakkinen [1997]. In order to change the ice export they varied the viscosity of sea ice. Their model produced an increase (decrease) of about 2 Sv in the maximum annual overturning rate of the North Atlantic as a response to the reduced (increased) sea Figure 11. Time series of anomalies of (a) freshwater transports in EXP1 across sections A (dashed line), B (dash dotted line), and C (dotted line) and (b) individual components of the freshwater transport across section C in the upper and middle layers combined (dashed line) and in the deep layer (dash-dotted line); also shown in Figure 11b is the net freshwater transport across section C (dotted line).

10 17-10 SAENKO ET AL.: NORTH ATLANTIC RESPONSE TO ARCTIC SEA ICE Figure 12. Time series of the overturning index anomaly in EXP3 (thick solid line) and EXP3a (dashed line); also shown for comparison is the overturning index anomaly in EXP1 (thin solid line). ice export of 800 km 3 yr 1 (0.025 Sv) over 15 years. In our experiments the largest reduction of the overturning ranges between 2.5 and 5.5 Sv, depending on the magnitude of the excessive ice export from the Arctic (Figure 13). However, after about years the THC begins to recover (Figure 13b). Because the model we use does not have any restoring to observations, there must be an internal negative feedback in the system between the ice export from the Arctic and the North Atlantic THC, which operates on the decadal timescale. The feedback tends to bring the system back to its normal (control) state, even when the above-normal sea ice export persists. Below we apply the freshwater/heat budget analysis to understand what processes are involved in this ability of the THC to resist the persistent inflow of extra sea ice from the Arctic. [37] The timescale of years, needed for the THC to start recovering suggests, based on our results from section 4.1, that one of the processes involved in the THC recovery could be that associated with the downward propagation of the positive freshwater anomaly in the North Atlantic. Another contribution could come from the propagation of a negative freshwater anomaly laterally from the Arctic to the North Atlantic as a result of the Arctic Ocean s salinification. These notions are supported by the analysis below. Again, we mainly focus on one of the experiments, i.e., EXP2. Similar analysis and conclusions apply for EXP2a and EXP2b. [38] The adjustment of the NA box to the inflow of extra ice is shown in Figure 14. The freshwater content in the upper layer of the NA box rapidly increases during the first 15 years and then starts to decrease gradually. The middle layer keeps freshening for up to about 30 years (Figure 14a), tending to make the water column less stable. The (positive) freshwater anomaly reaches the deep layer of the NA box and then is being gradually removed (see Figure 15) with, though weakened, western boundary flow. By about year 20 the liquid freshwater flux from the Arctic approximately reaches its minimum (Figure 14c; see also Figure 16), indicating a propagation of the (negative) freshwater anomaly from the Arctic Ocean to the NA region. [39] The response of the heat content is similar (qualitatively) to that in EXP1: there is an initial increase of the amount of heat trapped in the middle layer, followed by its reduction back to and below the control value at around year 15 (Figure 14b). The upper ocean heat content falls below its control value mainly because of the decreased THC (through its effect on the transport of heat from the subtropics) and because of the decreased heat flux from the subsurface ocean. [40] The saturation of the NA upper layer with fresh water is a result of the decreased liquid freshwater transport from the Arctic Ocean (Figure 14c). This is mainly due to the Arctic Ocean salinification (Figure 14a). In EXP2 it is a reduction of the freshwater transport through the Canadian Archipelago that dominates the net decrease of lateral freshwater transport from the Arctic Ocean (Figure 16). However, the dominant pathway taken by the freshwater anomaly to leave the Arctic Ocean may depend on the magnitude of the ice flux anomaly. [41] Thus both the gradual downward propagation of the freshwater anomaly and the decrease of lateral liquid freshwater flux from the Arctic to the region of the NADW formation have about the same timescale of years. Their combined effect first increases a resistance of the THC to the persistent above-normal ice export from the Arctic and then turns the THC back toward its normal (control) state. As discussed in section 4.1, an additional factor stabilizing the THC is the atmospheric feedback. The air temperatures in the northern North Atlantic cool down by up to 2 C in EXP2 (not shown). The contents of Figure 13. Time series of anomalies of (a) ice export through the Fram Strait-Barents Sea region and (b) overturning index in EXP2 (solid line), EXP2a (dashed line), and EXP2b (dotted line).

11 SAENKO ET AL.: NORTH ATLANTIC RESPONSE TO ARCTIC SEA ICE Figure 14. Time series of anomalies of (a) freshwater (FW) content, (b) heat content, (c) transports of FW, and (d) of heat for the NA box in EXP2. In Figures 14a and 14b the solid, thin dashed, dash-dotted, and dotted lines denote the contents within the whole NA box, the upper 228 m, from 228 to 1128 m, and below 1128 m, respectively; in Figure 14a the heavy dashed line denotes the response of FW content of the Arctic Ocean above 228 m. In Figures 14c and 14d the solid, dashed, dash-dotted, and dotted lines denote the transports at the air-sea interface and through passages A, B, and C (see Figure 1), respectively. freshwater and heat follow the tendency of the THC to recover, turning back to their control values a few decades later (Figures 14a and 14b). Like in EXP1, the response of heat transported from the subtropics is highly correlated with the response of THC (see Figure 13b and Figure 14d). [42] Thus the internal redistribution of freshwater between the Arctic and North Atlantic due to the sustained ice export anomaly is not able to shut down the North Atlantic THC, at least for ice export of times above the normal. On timescales of more than 15 years the increase of ice flux to the North Atlantic tends to be compensated by the decrease of liquid freshwater transport from the Arctic Ocean to the North Atlantic (compare Figure 13a and Figure 16). It should be noted, however, that had the increase of freshwater flux to the North Atlantic not been associated with the ice export but specified externally, the North Atlantic THC could have collapsed. 5. Discussion and Conclusions [43] Pollard and Pu [1985] raised the question of If the surface layers of the North Atlantic freshened, where in the world did the oceans get saltier? Dickson et al. [1988] hypothesized that the GSA event was a result of increased freshwater export from the Iceland-Greenland Sea to the North Atlantic so that one should look into the northern Greenland Sea for a compensating salinification. However, a modeling study by Hakkinen [1993] seems to support a suggestion by Aagaard and Carmack [1989] that the GSA event could have been a result of enhanced export of freshwater from the Arctic, mainly in form of sea ice. Figure 15. Time series of anomalies of the individual components of the freshwater transport across section C in EXP2 in the upper and middle layers combined (thin solid line) and in the deep layer (dashed line); also shown is the net freshwater transport across section C (thick solid line).

12 17-12 SAENKO ET AL.: NORTH ATLANTIC RESPONSE TO ARCTIC SEA ICE Figure 16. Time series of anomalies of the lateral freshwater transports across (a) section A and (b) section B in EXP2 (solid line), EXP2a (dashed line), and EXP2b (dash-dotted line). [44] The export of sea ice from the Arctic through the Fram Strait has strong year-to-year variability, which is controlled by atmospheric circulation anomalies [e.g., Serreze et al., 1992; Kwok and Rothrock, 1999; Hilmer et al., 1998]. Because the general pattern of wind over the Arctic Ocean favors the export of ice to the North Atlantic with the Transpolar Drift Stream, the simplest way to accelerate the ice export is to increase wind speed over the Arctic ice cover. This has been attempted in this study using a global coupled model. Our results suggest that if the freshening of the North Atlantic is indeed associated with increased export of sea ice from the Arctic, then the answer to the above question by Pollard and Pu [1985] could be in the Arctic Ocean. [45] We have presented two sets of model experiments in an attempt to understand the response of the climate system, particularly in the North Atlantic, to the transient and persistent above-normal export of sea ice from the Arctic. Similar investigations have been performed before [e.g., Mauritzen and Hakkinen, 1997; Hakkinen, 1999; Goosse and Fichefet, 1999]. Our study addresses these issues by employing an interactive ocean-ice-atmosphere model, unconstrained by open boundary conditions or prescribed atmospheric air temperature and humidity. We presented a thorough discussion of the heat and freshwater budgets in the region of NADW formation in the model and how these respond to ice export anomalies from the Arctic. We showed that in the case of transient anomalies, it takes years for the THC to reduce and then recover. Moreover, the THC is able to survive the persistent increase of ice export from the Arctic and after about years begins to recover. It is shown that this decadal timescale is associated with the downward propagation of a positive freshwater anomaly within the northern North Atlantic and with horizontal propagation of a negative freshwater anomaly from the Arctic Ocean to the Atlantic. The main points can be summarized as follows. [46] 1. If the transient perturbation is sustained for 1 year, the overturning reaches its minimum approximately 5 years after, supporting the results by Hakkinen [1999]. However, our model simulates a change in overturning/heat transport of only 5 6% for this GSA-like case. In order to have both the overturning and heat transport changed by up to 20%, as found by Hakkinen [1999], the doubled ice export needs to be sustained for at least 5 years. This would result in the North Atlantic freshening by about 3 times more than that estimated for the GSA event. [47] 2. In the response to the subpolar North Atlantic freshening due to sea ice inflow and melt the Arctic Ocean gets saltier, in agreement with findings by Hakkinen [1999]. However, our century-long simulations do not support the notion that the simulated climate may switch to a new quasi-equilibrium [Hakkinen, 1999]. Rather, the climate gradually recovers after the perturbation is turned off. It takes years for the THC almost fully to recover. However, the budgets of freshwater and heat continue to evolve for up to 40 years after the perturbation is set to zero. [48] 3. Imposing a persistent above-normal ice export does not cause the North Atlantic THC to collapse. Rather, after years it begins to recover. Under these extreme conditions the contents of heat and freshwater in the subpolar North Atlantic do not show a tendency for recovery until a few decades after the THC begins to return back to normal. [49] 4. The stability of the THC to the persistent abovenormal export of sea ice from the Arctic appears to be due to two processes, operating on a timescale of years. The first process is associated with the vertical propagation of a positive freshwater anomaly within the subpolar North Atlantic with weakened but still active THC. The second process is associated with the horizontal propagation of a negative freshwater anomaly from the Arctic Ocean to Atlantic. Also, the use of active atmospheric component is important for stabilizing the overturning circulation. [50] Finally, this study suggests that at least some fraction of the North Atlantic THC variability on decadal timescale can be attributed to the variations of ice export from the Arctic. [51] Acknowledgments. We would like to thank reviewers for helpful comments. This research was supported by the Canadian Climate Change Action Fund, Meteorological Service of Canada/Canadian Institute for Climate Studies, the International Arctic Research Center, and NSERC. We thank A. Schmittner and M. Eby for interesting discussions about this work. References Aagaard, K., and E. C. Carmack, The role of sea ice and other fresh water in the Arctic circulation, J. Geophys. Res., 94, 14,485 14,498, Bitz, C. M., M. M. Holland, A. J. Weaver, and M. Eby, Simulating the icethickness distribution in a coupled climate model, J. Geophys. Res., 106, , Bryan, F., Parameter sensitivity of primitive equation ocean general circulation models, J. Phys. Oceanogr., 17, , Dickson, R. R., J. Meincke, S. A. Malmberg, and A. J. Lee, The Great Salinity Anomaly in the northern North Atlantic , Prog. Oceanogr., 20, , England, M. H., and S. 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