Mechanisms for the variability of dense water pathways in the Nordic Seas

Transcription

1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 114,, doi: /2008jc004916, 2009 Mechanisms for the variability of dense water pathways in the Nordic Seas Rolf H. Käse, 1 Nuno Serra, 1 Armin Köhl, 1 and Detlef Stammer 1 Received 16 May 2008; revised 15 August 2008; accepted 30 October 2008; published 29 January [1] Interannual changes in simulated flow fields of the Nordic Seas are analyzed with respect to their dynamic causes and consequences regarding the flow of dense water from the Nordic Seas into the subpolar North Atlantic across the Greenland-Iceland-Scotland Ridge. The simple case of pure density-driven outflow with closed northern boundaries shows that dense water mainly originates in the northern Lofoten Basin and flows southward in three branches, namely along the Norwegian continental slope, along the Mohn and Jan Mayen Ridges, and a weak current along the east Greenland continental slope. Adding variable exchange through Fram Strait shows a strengthening of the most western branch and strong recirculations that may reverse the other two branches. For this case, we find in-phase modulation of the Denmark Strait overflow (DSO) by a changing Fram Strait supply and a Faroe-Shetland transport that is in opposite phase. The scaling of this relation provides a potential explanation of recently observed DSO changes. However, details of the changes in the simulated pathways suggest, in accord with the size of the prescribed varying Fram Strait supply, basin-wide wind stress curl and local convection, which feeds water from different source regions into the outflow pathways, as the primary cause for the upstream flow field reorganizations. Citation: Käse, R. H., N. Serra, A. Köhl, and D. Stammer (2009), Mechanisms for the variability of dense water pathways in the Nordic Seas, J. Geophys. Res., 114,, doi: /2008jc Introduction [2] The Nordic and Arctic Seas are important elements of the global circulation because warm and saline Atlantic surface waters are there transformed to cold dense water masses returning to the North Atlantic at depth. In this loop, the Greenland, Iceland and Norwegian (GIN) Seas constitute an effective buffer (storage). It is there where transformed water masses from the North Atlantic and water from the Arctic Mediterranean are finally mixed to form those water masses that represent the core of dense waters which, once entering the subpolar North Atlantic, become the lower limb of the global meridional overturning circulation (MOC [e.g., Swift, 1984; Schmitz and McCartney, 1993]). Any change in the dense water supply to the subpolar North Atlantic via the Greenland-Iceland-Scotland Ridge (GISR) system may therefore have a profound impact on the variability of the MOC [e.g., Bacon, 1998; Köhl and Stammer, 2009]. [3] The GISR consists of two main passages (Figure 1): the Denmark Strait (DS) and the Faroe Bank Channel (FBC). Water masses exported through either of those two passages have quite distinct water mass properties and therefore also distinct source regions. Early reports characterize the DS overflow as consisting of Greenland Sea Deep 1 Institut für Meereskunde, Zentrum für Meeres- und Klimaforschung, Universität Hamburg, Hamburg, Germany. Copyright 2009 by the American Geophysical Union /09/2008JC Water (GSDW) and the FBC overflow as consisting of Norwegian Sea Deep Water (NSDW [Swift et al., 1980]). However, there seems to be a consensus that, nowadays, returned (and cooled) Atlantic Water (RAW), as well as Arctic and Nordic Seas Intermediate Waters, form the core of the outflowing water masses [Crease, 1967; Borenäs and Lundberg, 1988; Mauritzen, 1996; Hansen and Osterhus, 2000; Borenäs et al., 2001]. This suggests that significant shifts in the properties and possible creation mechanisms of outflowing water masses have occurred over the last decades. [4] To understand the impact of the GISR overflow water and its related fluctuations on a changing MOC, it is important to consider the total export of the Nordic Seas across the entire GISR, because transport changes in individual passages may compensate each other. Moreover, we have to understand changes in the pathways of the dense waters and variations in their source mechanisms if we want to understand the cause of GISR overflow changes and their predictive skill on the MOC variability. In this context one may consider the entire Nordic Seas-Subpolar North Atlantic system as modulated by a large-scale wind-driven cyclonic gyre surrounding Iceland, that, if modified, can enhance the DS overflow, while simultaneously reducing the FBC overflow, or vice versa [Biastoch et al., 2003]. [5] Macrander et al. [2005] observed for the first time a reduction of the DS overflow by 20 30% during 4 years and argued that during this time frame the FBC overflow was increasing while the net overflow seemed to be almost unchanged. In the context of Nordic Seas variability, the establishment of a new additional DS overflow path from 1of14

2 Figure 1. Model domain showing the main topographic and geographic features of the Nordic Seas and subpolar North Atlantic. The inset shows in detail the Faroe-Bank Channel region and gives an impression of the model resolution. the Iceland Sea to DS by Jonsson and Valdimarsson [2004] raises the additional question where this flow has its origin. Is it a result of stronger localized convection in certain basins or increased export of dense water from the Arctic basins that produces temporarily enhanced flows on certain pathways? Are decadal/interdecadal changes in the curl of the wind stress responsible for the variability? Clearly, to explain changes in the properties of the GISR overflow water, we need to understand not only the processes changing the water properties in the vicinity of the ridge, but also those changing the route of dense waters toward the pathways and in addition also the water mass formation properties on a basin scale. [6] It has been shown that the hydraulic character of the outflow through the passages of the GISR theoretically smoothes the transport variations [Käse, 2006]. According to this, variability on subdecadal scales are predominantly seen in local storage variations and less so in the outflow. Considering changes in the overflow properties, one therefore has to think about water mass formation changes upstream of the passages. [7] This paper is concerned with temporal variability of dense water pathways in the Nordic Seas and its impact on overflow transports. Accordingly, our focus will be on the upstream origin of changes that have been observed and modeled for the DS by Köhl et al. [2007]. We will show that circulation changes occurring within the Nordic Seas, feeding different water masses to the DS and the FBC, are correlated with the basin-wide wind stress curl, the Arctic export of dense water through Fram Strait and with changes in convection regimes in the Nordic Seas. [8] The study is based on and follows from the numerical simulations reported by Köhl et al. [2007], who investigated causes of changes in the DS overflow. The time averaged circulation of dense water (denser than s q = 27.8 kg m 3 )in the high-resolution Köhl et al. [2007] model is depicted in Figure 2. Large-scale cyclonic gyres are present in the Greenland and Iceland Seas and in the Norwegian and Lofoten Basins, the latter featuring an embedded strong quasipermanent anticyclonic eddy [Köhl, 2007]. The most important pathways of dense water upstream of DS and FBC reside along the east Greenland slope and along the Mohn and Jan Mayen Ridges. Two other pathways are significant, the first along the north Iceland slope and the second along the Norwegian slope. In this work we will investigate changes in these dense water pathways and explain them according to forcing agents. 2of14

3 Figure 2. Time-averaged vertically integrated volume fluxes (m 2 s 1 ) of the dense water (s q 27.8 kg m 3 )in the Köhl et al. [2007] numerical simulations. Vectors are colored according to transport magnitude: black, between 5 and 15 m 2 s 1 ; red, between 15 and 100 m 2 s 1 ; blue, larger than 100 m 2 s 1. The thick black line corresponds to the 350-m isobath. [9] The structure of the remaining paper is as follows. Section 2 will describe the methodology used in our study and section 3 discusses changes of the dense water circulation of the Nordic Seas produced by different forcing mechanisms. Section 4 analyzes transport balances and water mass conversions. Section 5 finally provides a discussion and concluding remarks. 2. Methodology [10] Our study is based on the regional model described in detail by Köhl et al. [2007]. The model uses the MITgcm code and is implemented for the eastern subpolar North Atlantic and Nordic Seas from 51 Nto78 N and from 46 W to17 E (see Figure 1). The realistic bottom topography is extracted from the Smith and Sandwell [1997] data set. In the vertical, the model has 30 levels with thickness varying from 10 m near the surface to 500 m at the largest depth. Between m (which is the range of the deepest GISR passages), the vertical resolution is about 100 m. The model includes a parameterization for vertical mixing by the KPP scheme of Large et al. [1994] and a dynamic/thermodynamic sea ice model [Zhang and Rothrock, 2000]. [11] In our study, the above mentioned model is used for several new experiments but now run in idealized settings with 1/4 horizontal resolution and with idealized initial and boundary conditions. Background coefficients of vertical diffusion and viscosity are 10 6 m 2 s 1 and 10 4 m 2 s 1, respectively. Horizontally, biharmonic diffusion and viscosity represent unresolved eddy mixing. Coefficients of horizontal diffusion and viscosity are both m 4 s 1. [12] The idealized experiments, which are listed in Table 1, have no surface heat and freshwater forcing and were performed to separate the effects of different forcing mechanisms upon the dense water pathways and overflow variability in the Nordic Seas. Specifically, we investigate the role of (1) pure density-driven cross-ridge exchange, (2) reaction to Arctic export, and (3) varying wind stress curl. The lower resolution is justified since the focus is not on overflow hydraulic effects, but on large-scale upstream circulation pathways. [13] The initial and boundary conditions for our experiments are specified as follows. In the first experiment (DAM BREAK), we used an idealized two water mass distribution as specified by Käse and Oschlies [2000]. In this case, spatially uniform density values south and north of the GISR are associated, respectively, with temperatures of 5 C and 1 C at constant salinity of 35. In a second experiment (LEVITUS) initial conditions were taken from the January state of the WOA2001 climatology, i.e., spatially varying. In LEVITUS WIND the mean wind stress of either or are additionally applied. For these experiments, the lateral boundaries are closed. In experiments FRAM DEEP and FRAM SHALLOW, the model is driven laterally by idealized settings of prescribed climatological values of temperature, salinity and velocity. Temperature and salinity are from the WOA2001 climatology and velocity is specified as barotropic inflow and outflow. Figure 3b (top) shows the resulting cumulative transports at the northern boundary split into northward and southward transports. Figure 3b (bottom), showing the average temperature across Fram Strait, motivated the selection of the transport distribution. Clearly visible is the warm Atlantic water inflow into the Arctic on the right hand side and the cold return southward flow over the Greenland shelf on the western side. In FRAM DEEP and FRAM SHALLOW experiments, the eastern and western boundaries are closed, while at the southern boundary, an inflow-ouflow of 8 Sv is imposed with temperature and salinity prescribed (at inflow points) from the WOA2001 climatology. 3. Sensitivity of Dense Water Pathways [14] A central question for our analysis is the cause of changes in the characteristics of water masses leaving both straits of the GISR. To that end, we will use the above sensitivity experiments to investigate potential causes for changes in the dense water pathways bringing different Table 1. Summary of Model Configurations for the Idealized Experiments Experiment Northern Boundary Condition Initial Condition Forcing DAM BREAK Closed Käse and Oschlies [2000] None LEVITUS Closed WOA02001 None LEVITUS WIND Closed WOA , Wind FRAM DEEP Steady deep exchange WOA02001 None FRAM SHALLOW Varying shallow exchange WOA02001 None 3of14

4 Figure 3. (a) Model bottom topography (in meters) for the study area. Dense water pathways are highlighted by white areas. Specific transport sections shown are discussed in the text: FSW, Fram Strait West; FS, Fram Strait; FSE, Fram Strait East; DS, Denmark Strait; EGC, East Greenland Current; IS, Iceland Sea; MOHN, Mohn/Jan Mayen Ridge Current; LOF, Lofoten Basin; FSC, Faroe Shetland Channel. (b, top) Cumulative total transport for two different inflow scenarios with Levitus initial conditions. The deep inflow case (FRAM DEEP, black) has all southward inflow in region FS while the second case (FRAM SHALLOW, red) has additional inflow in the shallower western Fram Strait (FSW). Note that there is no net transport through Fram Strait in all the idealized runs. (bottom) Model temperature section (in C) at the northern boundary (78 N). water masses toward the sill. We will first look at the results of experiment DAM BREAK, addressing a purely densitydriven exchange in an idealized closed basin with only two distinct water masses north and south of the GISR that are representative of the density contrast between the Nordic Seas and the subpolar North Atlantic Density-Driven Exchange [15] The DAM BREAK experiment consisted of a run which starts from two homogeneous densities north and south of the GISR and was integrated for 10 years. During this period the near surface density front situated initially 4of14

5 over the GISR, separating the Nordic Seas dense water from the warm Atlantic water, moves northward. The water mass front, shown in Figure 4a for the initial time as the transition between the shaded and unshaded areas, is captured subsequently by the 4.0 C and 4.1 C isotherms. After 1 year the front still resides in the vicinity of the GISR; however, the enhanced inflow from the Atlantic is already indicated by a more poleward position in the eastern part. This shape continues during the following years. The fact that the front moved northward results from the baroclinicity of the flow, moving light water northward near the surface, while the deep dense water is flowing, steered by the topographic slopes, in the opposite direction. The front reaches the northern wall in 6 years on the Norwegian side (around 10 E in longitude). [16] At depth, the colder dense water (specified here by water colder than 0 C) is propagating southward through the channels (Figure 5a). We note that one part of the flow along the Jan Mayen Ridge hits the Iceland slope and continues directly toward Denmark Strait while the other part joins the Lofoten flow and feeds the Faroe-Shetland channel. Apparently, there is some loss of water colder than 0 C, seen from the continuously diminishing values flowing through both channels. We also note that the largest southward transport close to the northern wall agrees with the most extreme poleward position of the 4 C isotherm in the upper layer, suggesting that water mass formation is favored in that region due to pathways of deep flow (pull of upper layer water). [17] The volume of warm water that has moved northward in 5 years (when the northern wall is reached) is about 0.8 millions of cubic kilometers corresponding to a net northward flow of 5 Sv. In more detail, we seem to find the dense water mainly originating in the northern Lofoten Basin. It then splits into three branches: (1) one that continues southward along the Mohn Ridge, (2) the second which is stronger and flows along the Norwegian continental slope, and (3) a smaller one in the west flowing toward and further downstream along the east Greenland continental slope. [18] After 10 years, at 69.5 N, the Lofoten branch carries 2.7 Sv, the Mohn/Jan Mayen branch 2.2 Sv and the east Greenland branch only 0.1 Sv of volume transport (Figure 6), adding up to the 5 Sv previously estimated from the volume flux of warm water. We note that while individual pathways are adjusted after about 5 years, the total GISR overflow needs at least 10 years to reach some equilibrium for a predefined density level. In fact, there will be no equilibrium without production and the exchange will spin down continuously Influence of Arctic Export Through Fram Strait [19] During the second experiment (LEVITUS), a realistic and spatially varying initial density distribution based on 5of14 Figure 4. The adjustment of the surface front defined by the isotherms 4.0 C and 4.1 C as function of time for three experiments with different initial and/or boundary conditions: (a) idealized water masses (DAM BREAK experiment), (b) Levitus initial condition with inflow prescribed at the northern boundary through sections FS (FRAM DEEP experiment), and (c) same as Figure 4b but with an additional inflow at section FSW (FRAM SHALLOW experiment). Shaded area in Figure 4a corresponds to the initial conditions; in Figures 4b and 4c, the shaded area corresponds to the respective water mass front in the Levitus data set.

6 Figure 5. Instantaneous vertically integrated fluxes of the water layer with temperatures colder than 0 C after 10 years from (a) the DAM BREAK, (b) the FRAM DEEP, and (c) the FRAM SHALLOW experiments (in this case, the northern inflow is at its maximum phase). Thick vectors correspond to transports larger than 20 m 2 s 1 in Figure 5a and larger than 10 m 2 s 1 in Figures 5b and 5c. Vertical and horizontal lines correspond to the transport sections introduced in Figure 3. the Levitus climatology was used instead of the two homogeneous water masses of the DAM BREAK experiment. Results lead to findings comparable to the DAM BREAK experiment, except for a smaller overflow magnitude caused by the weaker initial density contrast in the upper layers, and are therefore not shown here. However, adding inflow from the Arctic had a significant impact as described next. [20] Current meter measurements obtained in Fram Strait over a 3-year period [Schauer et al., 2004] established a view of the variability of exchange of mass and heat between the Nordic Seas and the Arctic. Over the 3-year period reported, southward transports varied between 12 and 13 Sv and northward transports between 9 and 10 Sv. Since the exact value of net throughflow is not different from zero within error bars, we have postulated a balanced flow at Fram Strait and investigated the response to a steady or varying inflow/outflow of 15 Sv amplitude with a period of 2 years. Two experiments were conducted (FRAM DEEP and FRAM SHALLOW experiments), with deep and shallow inflows, respectively (see Figure 3b). [21] The FRAM DEEP experiment is similar to LEVITUS, but has prescribed steady inflow of dense water from the Arctic through Fram Strait (through section FS in Figure 3a) and a mass balanced export of light water in the east (through section FSE). The FRAM SHALLOW experiment has in addition a shallow inflow (in section FSW of Figure 3a), which is compensated again in section FSE. This flow is modulated by a sinusoidal oscillatory signal that brings down the southward flow to zero or to double amplitude with a 2-year period. [22] Figure 4b shows the time development of the warm surface front for the FRAM DEEP case. Instead of reaching the Greenland Sea within 5 years as in Figure 4a, the front is still south of Jan Mayen. This is due to the blocking by the increase of deep circulation in the Greenland Sea generated by the deep inflow in section FS. [23] Figure 4c shows the time development of the warm surface front for the FRAM SHALLOW case. As can be inferred from Figure 4, the thermal front does not move poleward, especially on the western side, but stays close to its initial climatological position because warm water is allowed to exit Fram Strait and cold water spreads along the east Greenland shelf. Since this exchange is periodically changing, the frontal response is mainly an east/west undulation close to Iceland and west of Spitsbergen. Years 1 and 5, labeled in Figure 4c, correspond to a minimum Fram Strait inflow while Year 2 correspond to a maximum phase. The general frontal displacement from the initial position in the northwest results from our simplification of a closed Barents Sea entrance. [24] The deep flow from both experiments with inflow through Fram Strait is shown in Figures 5b and 5c, respectively. In experiment FRAM DEEP (Figure 5b), the mean flow of dense water reveals a more pronounced East Greenland Current than found during the density-driven cases. However, most of the deep inflow is recirculating in the Greenland Sea basin. Qualitatively, the dense water branches are the same, only their strength is different. This is not surprising because the topographic steering favors regions of strong slope currents. Adding the shelf inflow in section FSW (FRAM SHALLOW), markedly leads to a 6of14

7 Figure 6. (top) Dense water volume transports on specific upstream sections (see locations in Figure 3). (bottom) The time dependence of the overflow at Denmark Strait (DSO) through the Faroe-Shetland Channel (FSCO) and the sum of both for the density-driven exchange DAM BREAK experiment. strong shallow western boundary current (Figure 5c) that is modulated periodically and that progresses toward Denmark Strait. [25] Transports on the individual sections for the FRAM SHALLOW experiment are depicted in Figure 7. Noteworthy is the direct in-phase modulation of the DSO by the changing Fram Strait supply. The FSCO is in opposite phase resulting from the behavior of the Lofoten and Mohn branches that act to compensate the supply in order to maintain the zero throughflow condition at the northern boundary. The mean transport through DS is about 3 Sv in this case and the amplitude of change is 2 Sv. If we scale the inflow according to the observed magnitude by Schauer et al. [2004], which has an amplitude of roughly 1.5 Sv, the resulting DSO change would be 0.5 Sv. This compares well to an observed variability of amplitude 0.4 Sv at DS [Macrander et al., 2005]. Note that the EGC has a larger amplitude (about 4 Sv) than the total Fram Strait flow, since part of it does not exit through DS. Instead, it recirculates via the IS branch north of Iceland. If there was no mixing at all, the sum of EGC, MOHN, and LOF would exactly equal FRAM. This is only approximately true during the first few years. After 10 years there is a difference of about 1 Sv that is lost from the dense to lighter water. This is not surprising since there is no renewal due to local surface buoyancy fluxes in this idealized setting Influence of Wind on Dense Water Pathways Variability [26] The simulated maximum GISR overflow reported in Köhl et al. [2007] occurred at a particularly high North Atlantic Oscillation (NAO) state in (NAO index and +1.83) and turned to lower values until (NAO index 0.5 and +0.79, see Osborn [2006]). However, other than the former sections would suggest, the prescribed inflow through Fram Strait are lowest during high NAO while the signal of the shallow inflow is less than 0.5 Sv and is thus not able to explain the simulated transport changes. Due to the northward shift of the westward storm tracks over the North Atlantic during high NAO conditions, the strength of the wind stress curl over the Nordic Seas is changing with the NAO index. The high NAO phase coincides also with observed large DSO [Macrander et al., 2005]. The circulation of dense water in the high resolution model is illustrated in Figure 8, which shows the vertically integrated transport in the layer with potential density s q larger than kg m 3. Strong southward flow directed to the GISR system is present in Figure 8 along the east Greenland shelf break and also along the eastern flank of the Jan Mayen Ridge. North of 70 N, northward flow can be found west of the Mohn Ridge and along the Norwegian shelf break. Superimposed is the dense layer topostrophy, T, which is a parameter useful to diagnose the existence of rim currents (either cyclonic or anticyclonic). The topostrophy is here defined as T = k (~v h rh), where ~v h is the vertically integrated (in our case below s q = kg m 3 ) flow field and rh is the gradient of the bottom topography (compare definition with Holloway et al. [2007]). T is positive for cyclonic flow (i.e., the deep basin is to the left of the flow) and negative vice versa. Figure 8 therefore suggests that during high NAO conditions, the circulation of dense water in the 7of14

8 Figure 7. As in Figure 6, but for the case with Levitus initial fields and a varying inflow at Fram Strait with a 2-year period (FRAM SHALLOW experiment). Nordic Seas feeding the GISR overflow is dominated by two cyclonic gyres separated by the Mohn and Jan Mayen Ridges and superimposed to local circulation features that presumably are topographically steered. Although the cyclonic circulation (reddish color in the topostrophy) dominates the Nordic Seas, there is a remarkable deviation from this cyclonic dominance near the entrance of the two main GISR passages. This holds near the DS, north of Iceland, and in the Faroe-Shetland Channel, north of Scotland. Both regions are seen as negative (or blue tone colored) topostrophy. [27] The circulation changes associated with the transition from high to low NAO are displayed in Figure 9, which shows the topostrophy difference ( minus ) of the dense water flow for the high-resolution experiment. Negative topostrophy difference is clearly visible near the eastern and western margins of the Nordic Seas, indicating a reduced East Greenland Current that feeds the DSO in association with larger transports along the Lofoten branch feeding the FBC overflow. The overall structure is that of a basin wide gyre change which was already described by Köhl et al. [2007] in the context of the associated sea level changes. To investigate further their hypothesis of a mainly wind driven change and to investigate the influence of this change on the dense water pathways in the Nordic Seas, the experiment LEVITUS WIND was performed, which adds wind-forcing to the previous LEVITUS experiment. Here two cases were considered, namely the time averaged wind stress forcing from 1998 to 1999 and from 2001 to 2002, representing relatively high and low NAO conditions, respectively. We then compare the flow differences between low and high NAO states in the LEVITUS WIND experiment with the corresponding high resolution model results. Figure 9b displays the topostrophy difference (case forcing minus case forcing) of the dense water flow in the idealized run. Note that the absolute value of topostrophy is resolutiondependent due to larger topographic slope in a model with higher horizontal resolution. Qualitatively, the difference plots are comparable with one exception: the flow along the Jan Mayen Ridge at about 8 W and between 65 N and 70 N shows an opposite (positive) sign in the realistic run. This corresponds to an enhanced southward flow during the phase of decreased DS overflow in the realistic simulation. We therefore cannot explain this feature by wind stress changes alone. However, wind stress was shown to be an active agent in determining large-scale changes in the flow field Changes on Pathways as a Result of Convection [28] Finally, convection, mixing and air-sea interactions in the Nordic Seas have to be considered. Several mechanisms participate in the water mass formation process occurring in the Nordic Seas. Among those are the changes in the inflow properties of water entering from the Atlantic. It was shown by Orvik and Skagseth [2003] that fluctuations in the warm Norwegian Atlantic Current (NwAC) correlated well with wind stress changes near 55 N 15 months earlier. Fluctuations in the water mass exchange between the Nordic Seas and the Arctic Ocean likewise impact the water mass formation in the Nordic Seas [Blindheim and Rey, 2004]. 8of14

9 Figure 8. Dense water topostrophy (units, m 2 s 1 ; see text for explanation) for the period of high DS overflow transport ( ) from the high-resolution model of Köhl et al. [2007]. Dense water (s q kg m 3 ) transport vectors are superimposed. [29] According to Gerdes et al. [2003], strong convection in the Greenland Sea appears to be correlated with negative phases of the NAO. Over the last decades, the NAO systematically increased from negative values in the sixties and early seventies to positive values in the nineties, resulting in a substantial warming of the Arctic and a reduced deep convection in the Greenland Sea. Nevertheless, Dickson et al. [1999] found no indication of a changing strength of the DS overflow in the Irminger Basin, suggesting that substantial inflow of Arctic dense water into the Nordic Seas through Fram Strait (FS) must have occurred to compensate for the reduced water mass formation in the Nordic Seas. [30] In contrast, Hansen et al. [2001] reported a decreasing thickness of NSDW near the ocean weather ship M for several decades and reported a decrease of the FBC overflow for almost 5 years from direct current measurements. This finding is not really plausible because the weather ship M location is not by itself an indication of the dense water reservoir. It has been shown [Borenäs et al., 2001; Mauritzen, 1996] that the FBC outflow includes North Iceland winter water that flows along the slope of the Iceland-Faroe Ridge toward the Shetland Channel. Interestingly, Olsen et al. [2008] do not support the Hansen et al. [2001] results in a new model analysis and by comparisons with current meter observations. [31] NSDW is regarded as a mixture of GSDW with Eurasian Basin Deep Water (EBDW) which is exported via Fram Strait into the Greenland Sea along the East Greenland continental slope but also along the Greenland Fracture Zone [Blindheim and Rey, 2004]. This mixture leaves the Greenland Basin via the Jan Mayen channel and fills the deep Norwegian and Lofoten Basins. One might speculate, therefore, that the diminished Greenland Sea convection is partly responsible for the reduced outflow of NSDW across the Iceland-Scotland Ridge. On the other hand, Fahrbach et al. [2001] reported also a significant change in heat transport through Fram Strait during the 1990s. [32] From recent model results, Eldevik et al. (personal communication) argue, however, that the commonly presumed causality between convective mixing in the Greenland Sea and changes in the overflows is neither evident in the observations nor in the model. Here we take a look at the outcrop region of the density surface s q = kg m 3, which is shown in Figure 10a for the strong outflow phase in DS ( , shaded) and the strong outflow phase in FBC ( , hatched). The reduced cyclonic vorticity in the Greenland Sea leads to a spin-down of the interior gyre, thereby reducing the doming of isopycnals toward the surface. This process leads to less convection as seen in Figure 10 of the outcrop area. The extension of the outcrop area is not only smaller in , but also shifted from the central Greenland Sea to the region southwest of Spitsbergen. [33] A rough estimate of the volume produced by convection within the Nordic Seas can be obtained by determining the summer depth of the isopycnal that defines the upper limit of the dense water (s q = kg m 3 ) and then multiplying it with the winter (here February) outcrop area. The respective time series is shown in Figure 10b (thick line). The time series shows a significant correlation (r 2 = 0.71 with r 2 = 0.55 being significant on the 95% significance level) with the annual mean of the DS overflow (dashed line). This is compatible with the view that the temporal variability of the dense water formed in the Nordic Seas may control the overflow through the GISR straits via changing the interface height upstream of the sills. [34] In our experiment, density variations are mainly controlled by (or at least correlated with) temperature variations and the region of temperatures between 1.8 C and 3 C is found to be colocated with the outcropping area. 9of14

10 KA SE ET AL.: NORDIC SEAS DENSE WATER PATHWAYS Figure 9. (a) Topostrophy difference minus (units, m2 s 1) as results from the high-resolution model of Ko hl et al. [2007]. (b) Topostrophy difference produced by applying the mean wind stress of and that of to the Levitus initial condition case (LEVITUS experiment). Temporal variations of the area defined by the temperature criteria (asterisks) are therefore highly correlated (r2 = 0.87) with the outcrop area time series. If the temperature is prominently controlled by local heat fluxes, the heat flux over the outcrop area should also be a good proxy for the creation of dense water. Indeed, the time series of the spatially averaged heat flux (Figure 10b, triangles) is found to be significantly (r2 = 0.62) correlated with our estimate of dense water production. 4. Volume and Heat Balances in the Nordic Seas [35] Oliver and Heywood [2003] applied an inverse model to determine transports of mass, heat and freshwater in nine different water mass classes from a hydrographic section across the Nordic Seas. Based on data from summer 1999, they estimated the southward volume transport across a line from Norway to Greenland to be 6.7 Sv for their dense water classes 4 9 (potential densities larger than sq = kg m 3). The transport of the light waters in classes 1 3 (sq < kg m 3) added up to 5.4 Sv, but all estimates are associated with large uncertainties. The authors also estimated a net heat transport, relative to their section mean, of 0.2 ± 0.08 PW toward the Arctic. We note, however, that the section was made during the high DS overflow reported by Macrander et al. [2005] and that it is still unclear how variable in time such estimates would be. [36] We use results from the sensitivity experiment FRAM DEEP and the high-resolution experiment to compute meridional volume transports within the same den- 10 of 14

11 Figure 10. (a) The outcrop region of the density surface s q = kg m 3 for the strong outflow phase in Denmark Strait ( , shaded) and the strong outflow in Faroe Bank Channel ( , hatched). (b) Time series of (thick line) the preceding summer depth of the s q = kg m 3 isopycnic multiplied by the February outcropping area in Figure 10a, (asterisks) the outcropping area diagnosed from the area of the January SST between 1.8 C and 3 C, (triangles) the February heat flux integrated between 70 N and 78 N, and (dashed line) the transport of water denser (s q = 27.8 kg m 3 ) through Denmark Strait filtered with a 1-year running mean. All curves are normalized by their standard deviation after removing the mean. sity classes as specified by Oliver and Heywood [2003, Figure 11]. The meridional transports shown were integrated from the bottom up to three different density levels: s q = kg m 3, s q = kg m 3 and s q = kg m 3. Although Oliver and Heywood [2003] did not consider s q = kg m 3 a water mass boundary in the Nordic Seas, we nevertheless include it here because many authors define the DS overflow as having densities larger than this value. [37] As can be seen in Figure 11a, the mean southward transport between Fram Strait and the DS is 4 Sv in the idealized experiment. South of the latitude of DS, mixing entrains surrounding water and the transport increases almost by a factor of 2, but looses its identity further downstream. The layer between s q = kg m 3 and the bottom has a slightly larger transport because water up to s q = kg m 3 is also part of the northward flowing 11 of 14

12 a transition of the overturning maximum to lighter water masses due to entrainment. However, the behavior downstream is rather different. While the entrainment of water denser than s q = kg m 3 is 6 8 Sv at different periods (3 Sv at s q = kg m 3 ), the water denser than s q = kg m 3 disappears completely reaching the latitude of 60 N. [39] The variability of the water mass divergence can be analyzed in more detail by looking at three different 2-year periods, , and for the three density limits (Figure 11b): s q = kg m 3, s q = 27.8 kg m 3 and s q = kg m 3. All lines closely intersect at 66 N, the location of the DS sill. We note that between 72 N and 74 N there is only little temporal volume change for all chosen density classes, an indication of the importance of processes south of the latitude of Jan Mayen. A remarkable feature is further that stronger flow north of DS leads to lower transport downstream in the same density class, suggesting that mixing is proportional to the strength of the flow. [40] Figure 12a presents the heat balance between 61 N and 69 N for the FRAM SHALLOW experiment during one full inflow/outflow cycle. In the absence of surface heat fluxes, the heat storage (i.e., the time derivative of the box integrated heat content) is balanced by the net heat flux through the northern and southern boundaries. The heat budget follows from H stor þr~f ¼ D; where H stor is the heat storage, D is the diffusion and Figure 11. Time-averaged meridional transport integrated from the bottom up to three different isopycnics (s q = kg m 3, s q = 27.8 kg m 3, and s q = kg m 3 ) from (a) the FRAM DEEP experiment and (b) the highresolution model. In Figure 11b, colors correspond to different time periods and line styles to different boundary isopycnics as follows: solid lines, s q = kg m 3 ; dashed lines, s q = 27.8 kg m 3 ; long dashed lines, s q = kg m 3. Atlantic water. This water mass is, however, diluted by entrainment so that at 60 N it is almost absent. The layer denser than s q = kg m 3 is less exported through Fram Strait and has a southward transport of only less than 3 Sv. This is due to northward flowing light water reducing the net transport. The large increase south of 66 N again emphasizes the role of further entrainment. [38] Concerning the results from the eddy-resolving run, we encounter a main difference to our sensitivity result in the region north of the sill. In the idealized run no local water mass production is happening due to missing surface heat and freshwater fluxes, so the transports are almost unchanged from Fram Strait to DS. On the other hand, the transport of water denser than s q = 27.8 kg m 3 in the high resolution run increases by 4 Sv from Fram Strait to Denmark Strait at 66 N. It is noteworthy that all curves for the three different density limits intersect at 66 N, indicating r~f ¼ F north F south þ F east F west þ F top F bot is the heat flux divergence. [41] In the idealized run, there is no flux through eastern and western walls and through the top or bottom. As we see in Figure 12a, the balance holds within reasonable error margins. The meridional distribution of zonally integrated heat flux for the extreme inflow/outflow phases of the FRAM SHALLOW experiment are shown in Figure 12b. The extreme positive and negative deviations from the mean north of 69 N are rather uniform (0.18 PW and 0.06 PW, respectively). A large divergence occurs southward of 69 N and is mainly balanced by local storage. [42] The corresponding heat balance between 61 N and 69 N for the case of the high-resolution model can be seen in Figure 13. We show the sum of storage and surface heat flux between 61 N and 69 N as a dashed line and the meridional heat flux at the northern boundary (69 N) as a solid line in Figure 13 (bottom). The heat transport at the southern boundary (61 N) is the dashed line in Figure 13 (top) and the dotted line represents the sum of storage, surface heat flux and the transport at 69 N. Without diffusion out of the regarded open box, the latter two lines should perfectly match. This is not the case, but the mismatch is on the order of 0.02 PW and therefore in agreement with the diffusive transport shown in Figure 12a. It is interesting to note that on longer timescales the DS overflow (drawn multiplied by 1/19 as a solid line in Figure 12a, top) is highly correlated with the heat transport 12 of 14

13 Figure 12. (a) Heat balance between 61 N and 69 N during a full cycle of periodic forcing at Fram Strait (FRAM SHALLOW experiment): thin, flux at 69 N; thin dashed, flux at 61 N; thick, storage; crosses and dots, diffusion at each latitude; squares, residual. (b) The meridional heat flux as function of latitude for the extreme inflow/outflow phases of the experiment. delivers water along preferred pathways east of the Mohn and Jan Mayen Ridges. This pathway lacks direct confirmation from current meter measurements, however, it has been predicted based on dynamical reasoning by Orvik [2004]. There it is suggested that the deep counterflow underneath the Arctic front amounts to about 1 Sv. This is in agreement with our model results. Additional supply from the Arctic Ocean through Fram Strait takes different routes, depending on the location of the prescribed inflow to the Nordic Seas, enhancing or reducing the transport found in the unforced case. Noteworthy is this: without convection and surface momentum stress, the density difference across the GISR drives a baroclinic circulation north of the sills that transports warm and saline Norwegian Atlantic Water into the Norwegian and Lofoten Basins, replacing the southward flowing dense water. The bulk of the overflow passes primarily through the Faroe Shetland Channel and exits via the Faroe Bank Channel. About 1/3 is carried via the Iceland Sea and exits through Denmark Strait. [44] Adding inflow to the Greenland Sea at the deep parts of the East Greenland slope, without affecting the transports on the shelf and shelf break, changes the throughflow only marginally but leads to a stronger cyclonic gyre in the central Greenland Sea. The situation changes markedly if inflow at the shallower regions is included. Then, the East Greenland Current carries a large amount of dense water. Its part above the DS sill depth flows into the Irminger Sea and its deeper part is topographically steered along the Iceland branch and follows then back north along the Jan Mayen Ridge. [45] By adding wind-forcing typical of NAO+ and NAO regimes, we find that its effect is similar to the shallow Fram Strait inflow. Comparing the idealized case with the fully forced run of Köhl et al. [2007], changes in dense water production suggest that this difference is the result of convection being shifted toward the Lofoten Basin during the early years of The effect of this enhanced convection in the Lofoten/Spitsbergen area is to increase the dense transport along the Mohn and Jan Mayen Ridge path and consequently the westward flow in the Iceland branch. However, this cannot compensate for the larger reduction of at 61 N. If, in agreement with hydraulic control, the heat transport is seen as a consequence of the DSO, the shown sum has also to be, because of the smallness of the diffusive transport, regarded as a consequence of DSO. The possibility of a following northward propagation of this heat anomaly is then indicated by multiyear delay. Figure 13 indicates a 1.5-year lag of the heat transport at 69 N tothe transport at 61 N which is in agreement with the lag between these components of about 14 months from FRAM SHALLOW shown in Figure 12a. This is remarkable and shows a generic response of the system, since the former is driven by wind stress changes and the latter by changes in transport through Fram Strait. 5. Discussion and Concluding Remarks [43] We have run a numerical ocean model with several idealized forcing scenarios to look into the question of changes in the transport of dense water upstream of the GISR. It was shown that the density-driven exchange Figure 13. Heat balance between 61 N and 69 N from the high-resolution model (see enclosed explanation of each curve). 13 of 14

14 the EGC which is due to the reduced cyclonicity of the wind field during NAO. [46] We conclude that the pattern of changes in the dense water circulation in the Nordic Seas can be explained by a combination of varying inflow from the Arctic, wind stress and heat flux changes superimposed on the basic basin-wide density-driven circulation. The full model of Köhl et al. [2007] can be partly regarded as a hindcast of the Nordic Seas circulation, whereby the part of circulation change due to changed transports though FRAM Strait is not simulated but represented by in-phase changes due to wind stress changes. Dickson et al. [2008] have discussed the observed fluctuations in the DS overflow [Macrander et al., 2005] and found a general slowdown after 2000 further downstream in the VEINS Angmassalik current meter array as well as in the model simulations of Olsen and Schmith [2007]. 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