Exchange of mass, heat and carbon across the Barents Sea Opening

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2 Exchange of mass, heat and carbon across the Barents Sea Opening Peter M. Haugan University Courses on Svalbard, Longyearbyen, Norway, and Geophysical Institute, University of Bergen, Norway January 13, 1999 Abstract Atlantic Water (AW) flows through the Barents Sea Opening (BSO) from the Norwegian Sea into the Barents Sea where some ofit recirculates back to the Norwegian Sea while another part, strongly cooled, escapes to the Arctic Ocean. Vessel mounted Acoustic Doppler Current Profiler (ADCP) observations from show that flow across the BSO, corrected for tides, has a dominating barotropic component and is highly variable on daily time scales with velocities of order 20 cms- 1 Inflow of AW to the Barents Sea on the southern slope of the Bear Island Trough occurs predominantly in two cores. On the steep slope of Svalbardbanken dose to the Polar Front, A W sometimes flows into the Barents Sea as a retrograde warm core jet while at other times there is outflow in the entire water column. A review of published model and observational studies of the region is combined with analysis of the ADCP data to produce an estimate of the recirculation and throughflow of AW across the BSO. The associated net exchange of heat and carbon is discussed in relation to other studies of heat and carbon budgets. The sensitivity of the annual mean exchanges to large scale forcing anomalies is discussed. 1

3 1 INTRODUCTION 1 Introduction The Barents Sea is an area of intense water mass formation (Hiikkinen and Cavalieri (1989), Hopkins (1991), Simonsen and Haugan (1996)). Through its western flank along the continental slope between Norway and Spitsbergen, and named the Barents Sea Opening by Hopkins (1991), some fraction of the warm and saline water originating from the North Atlantic and flowing along the Norwegian continental slope as the Norwegian Atlantic Current (NAC), turns eastward bringing heat, salt and nutrients into the Barents Sea. This fuels a large productivity providing the basis for food chains including important commercial fish species as well as marine mammals and birds. In the northern parts of the Barents Sea, along the boundary to the Arctic Ocean, perennial sea ice cover prevails. Between these extremes, the shelf sea experiences a seasonal ice zone and hydrographic conditions noted for large seasonal and interannual variability. The Barents Sea has been studied by many investigators during the last 100 years or so, motivated both by the special role this marginal sea may play in the physical climate system and by the close links between the physical climate state and the biological systems. The Institute of Marine Research (IMR) in Bergen, Norway, has for a number of years performed surveys of the climatic conditions in the Barents Sea and upstream in the Norwegian Sea along the Norwegian coast providing the basis for annual nowcasts and forecasts (Aure, 1998). Since 1988, IMR ships have been using vessel mounted Acoustic Doppler Current Profilers (ADCP) providing underway measurements of the vertical profile of horizontal current vectors. In the present paper such data obtained in the Barents Sea Opening from the years are analyzed and used in combination with a review of available published information from observations and modelling studies to shed some light on the following questions: What is the magnitude, geographical distribution and seasonality of the Atlantic Water (AW) inflow through the BSO? How much of the inflow returns to the Norwegian Sea rather than flowing through to the Kara Sea and the Arctic Ocean? Which physical processes and external forcings govern the exchange across the Barents Sea Opening? What are the transport averaged residence times of the recirculation and throughflow? What are the implications of the residence time for heat and carbon exchange with the atmosphere while in the Barents Sea? It will be clear that adequate answers to all of these questions can not be given here, and we will have to be content with preliminary estimates based on the available information. Yet, the present study is intended as a contribution towards improved understanding of these issues which all have a bearing on determining the role of the Barents Sea in the physical climate system and the carbon cycle. 2

4 1 INTRODUCTION The study is to a large extent motivated by recent work indicating that the water mass formation rate in the Barents Sea may constitute a major fraction of that of the entire Arctic Mediterranean (the Nordic Seas and the Arctic Ocean). In particular, in the study by Simonsen and Haugan (1996) based on large scale advective ocean heat budgets and sea surface heat flux estimates, it was found that the best estimate of the annual mean surface heat loss from the Barents Sea was about a factor of two larger than the net advective heat exchange obtained from classical estimates based on volume transports and temperatures of water flowing into and out of the Barents Sea. Although there are many uncertainties in the approach of Simonsen and Haugan (1996) for estimating sea surface heat loss, the large mismatch and the availability of new observations from the BSO, suggested a detailed study of the advective exchanges as attempted in the present paper. Arctic. Ocean Fr&rt& Jout Land HO = Hopen Deep, H " Hopen. E " E~ya, S" Stortjorden. 100m depth contouili. Figure 1: Study area The study area is shown in figure 1. Note that we use the name Svalbardbanken for the shallow area around Bear Island and Hopen. This bank has also been called Spitsbergenbanken in the literature. We prefer the former since Svalbard is the official name of the entire group of islands including (West) Spitsbergen, Bear Island, Hopen, Edge0ya etc. i.e. including islands which are actually situated on the bank. The name Spitsbergen is reserved for the largest of the islands in Svalbard, formerly often called West Spitsbergen. We refer to the deepest part of the Barents Sea 3

5 1 INTRODUCTION to the south of Bear Island as the Bear Island Trough. This area has sometimes been denoted as the Bear Island Channel. The continuation of the Bear Island Trough east of Hopen is called the Hopen Deep (in keeping with the Norwegian name "Hopendjupet"), although the Hopen Trench has also been used in the literature. For all banks we will consistently use Norwegian spelling, e.g. Sentralbanken rather than the Central Bank and Storbanken instead of the Great Bank. The Arctic Mediterranean (Aagaard et al., 1985) is the customary name for the area consisting of the Arctic Ocean and the Nordic Seas, where the latter include the Greenland-Iceland-Norwegian (GIN) Sea as well as the Barents Sea (Hurdle, 1986). There is also an inflow of Atlantic derived water to the northern Barents Sea from the continuation of the West Spitsbergen Current (WSC) north of Svalbard along the slope of the Eurasian basin of the Arctic Ocean. This inflow is small relative to that occurring in the BSO (Mosby (1938), Rudels (1987), Loeng et al. (1997), Schauer et al. (1997)) and furthermore seems to be largely isolated from the branch flowing in through the BSO in the sense that water overspilling the ridges between the southwestern Atlantic domain and the depressions cutting into the Barents Sea from the north seems to always be cooled and modified to such an extent that the heat transport is very small and the throughflow, if any, can hardly be classified as Atlantic Water. Pfirman et al. (1994) distinguish between northern Barents Atlantic derived Water and southern Barents Atlantic derived water. It will be assumed here that the contributions of AW inflow to the northern Barents Sea from the Arctic Ocean between Svalbard and Frans Josef Land can be neglected in lowest order mass, heat and carbon budgets for AW flowing in southwest of the sill or ridges, and rather can be viewed as transport back and forth between the Arctic Ocean and the northern Barents Sea (Rudels, 1987). Admittedly AW entering from the north is expected to be cooled during its residence in the northern Barents Sea, and thus the cycle of water transport back and forth implies a net import of heat from the Arctic Ocean to the Barents Sea. For the heat budget of the Barents Sea, there is a partially cancelling impact of cold Arctic surface water imported to the Barents Sea from the north and northeast. In this study we concentrate on the fate of AW flowing in through the BSO south of Bear Island, including a discussion of the possibility of this water to actually cross the ridges and flow out to the Arctic Ocean between Svalbard and Frans Josef Land. The area between the shallow Svalbardbanken and Spitsbergen, including the entrance to Storfjorden, is excluded since no Atlantic derived water is expected to cross the bank. It is not the objective to present a complete volume or salt budget for the Barents Sea or to study ice distribution or flow of Arctic surface water or coastal water. These aspects will be touched upon only to the extent that they directly influence AW entering through the BSO. A recent summary of these wider aspects including a complete mean volume budget for the Barents Sea may be found in Loeng et al. (1997). The interest in the carbon cycle in this area is relatively recent, and there are few measurements and studies available (Anderson et al. (1998), Fransson et al. (1998)). The strong surface cooling in the Barents Sea provides a laboratory for investigating air-sea exchange and the functioning of the solubility pump which is associated with the increase of solubility of CO 2 in sea water when the 4

6 2 THE ROLE OF THE BARENTS SEA IN THE ARCTIC MEDITERRANEAN CLIMATE SYSTEM temperature is decreased. Effects of the solubility pump are therefore closely linked to the physical processes affecting water flow and transformation. Whether biological pumping due to temporary storage of carbon in biological pools unavailable for air-sea exchange is significant beyond the annual time scale, depends on vertical transport to near bottom layers or subtle mechanisms related to seasonal successions of ice cover and vertical mixing. A brief discussion of some of these processes is given in section 6.3 on carbon cycling together with estimates of carbon transports associated with the water transports. In section 2, a brief overview is given of the role of the Barents Sea in the large scale ocean circulation. Only some selected studies are referenced in section 2, while many other studies are discussed in more detail in later sections in connection with specific aspects of the exchange. The ADCP data set is described and analyzed in section 3. In section 4, the broad questions in the bullet points above are discussed. Published observational data and studies from the Bear Island Trough and the Hopen Deep are reviewed in section 4.1 in view of the insights provided by the ADCP data set. The evidence for throughflow is discussed in section 4.2, and a discussion of the effects of densification by surface heat loss is included in section 4.3, taking into account also some recent progress from published theoretical and modeling studies of convection. A large number of regional and large scale modeling studies covering the Barents Sea have been performed in the last decade or so, and we include in section 5 an overview of such studies, noting in particular how the different models can provide insight on throughflow, recirculation and the geographical and seasonal structure of the flow across the BSO. A synthesis is given in section 6 with a conceptual description of the mechanisms of BSO mass exchange and a budget intended to be applicable to the time period in section 6.1. Implications for heat and carbon exchange are discussed in sections 6.2 and 6.3, respectively. Section 7 includes a discussion of the findings in relation to decadal and longer time scale climatic variability as recently reported both in the ~orth Atlantic and Norwegian Sea (section 7.1) and in the Arctic (section 7.2), and some conclusions and final remarks are given in section The role of the Barents Sea in the Arctic Mediterranean climate system Already Nansen (1906) had observations of cold, high salinity water from the eastern Barents Sea and drew attention to a potential role for such locally formed water to renew bottom water in the Arctic via the St. Anna Trough. A modern view is presented by Rudels et al. (1994) who sketched the branching of the Norwegian Atlantic Current at the BSO and discussed the impacts of the two branches on the Arctic Ocean, primarily based on data from the ODEN-91 expedition. Rudels et al. (1994) showed that the Barents Sea branch with its expanded density range, when it reaches the Arctic Ocean, not only interacts with the WSC branch affecting the Atlantic layer of the Eurasian Basin, but also contributes more strongly to the intermediate water below. It also affects the deep water of the Canadian Basin (Jones et al., 1995). Thus, processes in the Barents 5

7 2 THE ROLE OF THE BARENTS SEA IN THE ARCTIC MEDITERRANEAN CLIMATE SYSTEM Sea may be of importance for a wide region, and not only for the climatic state of the shelf sea itself. The two branches of inflow to the Arctic Ocean were judged by Rudels et al. (1994) to be of apparent equal strength (~ 2 Sv, where 1 Sv 10 6 m 3 s- 1 ). Interleavings and inversions are created where the two branches meet, and the WSC branch is displaced from the slope by the Barents Sea branch which feeds more of the overflow to the Canada basin while the WSC branch feeds cyclonic recirculation in the Eurasian basin. The magnitude and geographical distribution of the components of the surface heat flux over the Arctic Mediterranean was studied by Simonsen and Haugan (1996) based on sea surface heat flux parameterizations which gave total heat flux divergence in the Arctic Mediterranean in agreement with estimates of net heat advection based on oceanographic data. The advective estimates were based on a comprehensive review of published budgets using all available data back to early in this century, with the commonly applied steady state hypothesis. The atmospheric data were monthly means produced by blending data from several sources including the Comprehensive Ocean Atmosphere Data Set, mean data from European Centre for Medium Range Weather Forecasts and coastal weather station data from different periods (Simonsen and Haugan, 1996). The study of Simonsen and Haugan (1996) showed that the Iceland Sea gives annual mean surface heat loss of order 10 Wm- 2, similar to that of the Arctic Ocean, while the Greenland and Norwegian Seas have a heat loss close to 50 Wm- 2. The Barents Sea has around 100 Wm- 2, integrated to 136 TW, which is more than the heat loss integrated over the entire GIN Sea. The shallow depth of the Barents Sea implies that the heat loss there is distributed over a thin water layer, and therefore that the water mass properties change rapidly. In the Atlantic domain of the Norwegian Sea, the stratification near the base of the AW limits vertical mixing, giving substantial water mass formation also there (Mauritzen, 1996a). In contrast, in the Greenland Sea, where monthly mean surface heat loss in the marginal ice zone in winter time may become similar to that of the Barents Sea, of order 500 Wm- 2 (Hiikkinen and Cavalieri, 1989), the surface heat loss is distributed over a deep, weakly stratified water column so that the water mass properties change only slowly. Simonsen and Haugan (1996) noted the discrepancy between the estimated surface heat loss in the Barents Sea and available estimates of advective import of heat which primarily takes place with AW through the BSO. Based primarily on data from the GIN Sea during the 1980s, Mauritzen (1996a) proposed that the Barents Sea branch of A W, after recirculation in the Arctic and returning southwards through the Fram Strait, passes the Greenland Sea without much interaction with the Greenland Sea Gyre, and finally becomes the primary source of upper Norwegian Sea Deep Water (NSDW), which again supplies the most dense fraction of the overflow across the Greenland-Scotland ridge feeding North Alantic Deep Water (NADW) formation. This role ofthe Barents Sea branch in the NADW formation proposed by Mauritzen (1996a) was presented as a hypothesis derived mainly from the lack of other sources or pathways within the GIN Sea itself, capable of sustaining the relatively large (order Sv or more) and apparently seasonally and interannuajly stable upper NSDW overflow to the North Atlantic (see however (Turrell ct al., 1999) for evidence that the NSDW contribution to overflow has reduced in recent decades). The part of the qualitative hypothesis 6

8 2 THE ROLE OF THE BARENTS SEA IN THE ARCTIC MEDITERRANEAN CLIMATE SYSTEM of Mauritzen (1996a) which is concerned with the role of water mass formation in the NAC and the Barents Sea, and Barents Sea throughflow to the Arctic, gains some support from studies of the heat budget (Simonsen and Haugan, 1996), studies of water mass formation processes in the Barents Sea (Midttun (1985), Quadfasel et al. (1992)) and direct and indirect documentation of the outflow of modified water from the northeastern Barents Sea to the Arctic Ocean (Rudels et al. (1994), Schauer et al. (1997), Loeng et al. (1997)). However, the quantitative aspects are not so simple. In a companion paper, Mauritzen (1996b) constructed an inverse 15-box model of the Arctic Mediterranean based on conservation of mass, heat and salt. Initial surface heat fluxes were set to 70 W m- 2 everywhere in ice-free surface boxes, and inflow from the North Atlantic to the Norwegian Sea was initialized at 7 Sv. The standard solution of the model had an almost unchanged inflow of AW from the North Atlantic to the Norwegian Sea (down from 7 Sv to 6.8 Sv) and surface heat flux over the NAC box in the Norwegian Sea (down from 70 to 68 Wm- 2 ). The flow of AW across the BSO became 1.6 Sv (up from 1.5 Sv initially), but the surface heat flux of the Barents Sea box dropped to 25 Wm- 2 with a formal uncertainty of order only 5 Wm- 2. The surface heat flux in the Iceland Sea became almost 50 Wm- 2, but the formal uncertainty of that flux was almost as large as the flux itself. These surface fluxes differ substantially from those of Simonsen and Haugan (1996), who get a lower surface heat loss in the Iceland Sea and a much higher surface heat loss in the Barents Sea consistent with estimates of Hiikkinen and Cavalieri (1989). The work of Rudels (1987) also deserves mentioning here. Rudels (1987) assumed that the surface heat loss in the southwestern Barents Sea was 75 Wm- 2 which over an area of 0.6x1012m2 gave a total heat loss of 45 TW, while heat loss over the remainder of the Barents Sea was neglected (The entire Barents Sea has a total area of 1.32xl0 12 m 2 in the calculations of Simonsen and Haugan (1996)). The surface heat loss estimate of Rudels (1987) was balanced by a coastal branch of AW throughflow (0.8 Sv, temperature reduction from 7 to -2 C, contributing 30 TW) and an offshore branch of AW (1.44 Sv of which 0.4 Sv was throughflow and 1.04 Sv was recirculation, both experiencing a temperature drop from 3.5 to 1 C, contributing 15 TW). Other potential contributors to the heat budget were discussed but found to be negligible. The mass transport of the coastal branch was based on geostrophic estimates based on density difference between the coastal and offshore branches (0.4 kgm- 3 ) and the thickness of the layer at the coast, but the transport of the offshore branch was calculated so as to match the given surface heat loss. Rudels (1987) mentions that higher surface heat loss of order 100 Wm- 2 may be more appropriate than the 75 Wm- 2 base case he used, and that this can be accommodated by a corresponding increase in the volume transports, presumably his offshore branch. Given the more recent surface heat loss estimates of Hiikkinen and Cavalieri (1989) and Simonsen and Haugan (1996) which essentially agree with the highest of the heat loss intensities mentioned by Rudels (1987), but over an area more than twice as large, it is timely to investigate if these sea surface heat flux estimates can be combined with physically acceptable mass transports and cooling rates. While we do not attempt to test or evaluate the entire scheme proposed by Mauritzen (1996a, 1996b), the aspects concerning the Barents Sea branch and surface heat 7

9 2 THE ROLE OF THE BARENTS SEA IN THE ARCTIC MEDITERRANEAN CLIMATE SYSTEM loss there and the apparent quantitative inconsistency with Simonsen and Haugan (1996) will be addressed. We will use the ADCP data as well as information from a wide range of modelling and observational studies with the objective to get a better qualitative and quantitative understanding of this branch and its variability. It is intriguing that the model of Mauritzen (1996b) with its small surface heat loss from the Barents Sea, is used to argue for a stronger role of the Barents Sea in global thermohaline circulation (Mauritzen, 1996a) than hitherto assumed. The surface heat loss estimates of Simonsen and Haugan (1996) are supported by independent measurements (Hiikkinen and Cavalieri, 1989) and older climatologies as reviewed by Hopkins (1991). Can these estimates be reconciled with advective heat transport, and if so, what are the implications for the role of the Barents Sea in water mass formation and large scale circulation? Zooming in on the Barents Sea itself, Loeng (1991) gives an overview of the physical climate system of the Barents Sea and Loeng et al. (1992) discuss climate variability. References to earlier literature can be found in those two as well as other papers and will not be repeated here. Quite early it became evident that the heat import with the inflowing AW across the BSO is a decisive factor for the climatic state of the Barents Sea including its ice extent. Time series of ice extent from the turn of the century are clearly negatively correlated with the heat content in repeat sections crossing the AW. It was also evident to early investigators that signals in the Norwegian Atlantic Current (NAC) upstream in the Norwegian Sea may be useful predictors ofthe climatic state of the Barents Sea with lead times up to seasons and years depending on the distance from the predictor section to the Barents Sea. E.g. Loeng et al. (1992) showed that the mean temperature and salinity in the m depth range along parts of the Barents Sea Opening is closely correlated with the same quantities in a section further east in the Barents Sea, and that time lagged correlations also exist between similar mean quantities from several sections crossing the path of the NAC in the Norwegian Sea. It is unclear however, to what extent the anomalies detected in mean temperature and salinity are propagating with a relatively constant effective volume transport, or are due to variations in the volume transport with accompanying changes in the time available for modification of properties, or due to anomalies in the regional surface fluxes, or due to changes in the structure of the currents e.g. the vertical or lateral distribution of the transport along the mean current path, or the distribution of the total transport among the Barents Sea and WSC branches. Evidence, particularly from the WSC (Dickson et al., 1988), but also from the Norwegian and Barents Seas, discussed later, indicate that the amplitude of interannual anomalies may increase downstream, demonstrating that simple advedive propagation of anomalies is definitely not the whole story. To disentangle the various causes of variability may be difficult, particularly since several of the external forces are correlated and their effects may be felt simultaneously. E.g. increase in the strength of southwesterly winds over the North Atlantic, conveniently characterized by high values of the North Atlantic Oscillation (NAO) index (Hurrell and van Loon (1997), Dickson et al. (1996)), may increase transport into the Barents Sea, but will also bring in warm, moist air reducing the regional loss of sensible and latent heat, and increased cloudiness, which may influence the radiative balance, in particular the long wave net surface heat loss (Simonsen and 8

10 2 THE ROLE OF THE BARENTS SEA IN THE ARCTIC MEDITERRANEAN CLIMATE SYSTEM 2.1 Conditions in the Norwegian Sea Haugan, 1996). Correlations with precipitation and runoff are also significant (Dickson et al., 1996). Hurrell (1995) shows that December to March evaporation minus precipitation decreases (essentially due to precipitation increase) by more than 1 mm/day between winters with NAO index less than -1 and winters with NAO index greater than +1 over the entire path of the NAC in the Norwegian Sea and over Scandinavia. Influence on Arctic surface water (Blindheim et al., 1998) and intermediate water (Blindheim, 1990) may indirectly affect the transport and properties of AW reaching the Barents Sea. Since there are several distinctly different mechanisms possibly relating AW mass or heat transport with the NAO state, correlations between different observed time series of upstream and internal climate indicators must be treated with caution when the objective is to better understand causal relationships. Blindheim (1989) presented a set of widely quoted current measurements and associated transport estimates for the BSO exchange. Before we enter into a detailed discussion of the structure of that exchange, an overview is given of the conditions in the Norwegian Sea through which the NAC passes and where its properties and strength may be set. Then we discuss the climatic state and internal processes in the Barents Sea, and finally the outflow towards the northeast. 2.1 Conditions in the Norwegian Sea Mysak and Schott (1977) analysed current measurements from the Norwegian continental slope between 62 on and 63 on, near the inner part of what will be referred to as the Svin0Y section. This area was studied also by several of the early investigators including Helland-Hansen and Nansen (1909) who further out over the deep Norwegian Sea discovered what they called "puzzling waves", which we now denote mesoscale varibility. The current measurements used by Mysak and Schott (1977) were obtained during a 6 week period from late July to early September 1969 over a depth interval on the slope from about 500 m to 900 m with good vertical resolution at each of the four moorings. The mean velocity within the core of AW, which leaned onto the slope at depths between 200 and 500 m, was cm/s. Current shear was associated with the transition to Norwegian Sea Deep Water (NSDW) underneath. A dominating feature in the time series was oscillations with a period of 2-3 days, and Mysak and Schott (1977) investigated theoretical models to explain such waves. They proposed a model with baroclinic instability of the rapidly flowing AW overlying a slower moving and more dense NSDW over sloping topography. The most unstable waves of the model had structure and period matching the observed time series, and wavelengths of order 40 km. However, Schott and Bock (1980) later showed that barotropic instability transferred energy back to the mean flow at a greater rate. Blindheim (1993) estimated from hydrography that the monthly mean transport in 1990 through the Svin0Y section extending northwest from the Norwegian coast at 62 N, had a maximum of 7.9 Sv in January, a minimum of 2.9 Sv in September and an annual mean of 5.5 Sv. Orvik and Mork (1996) deployed current meters over the slope in the same section during April 1995 to August 1996 and found that the inner branch of the N AC over m depth had a very stable transport of 5.3 Sv during all months of the investigation period. Mork and Blindheim (1998) calculated the 9

11 2 THE ROLE OF THE BARENTS SEA IN THE ARCTIC MEDITERRANEAN CLIMATE SYSTEM 2.1 Conditions in the Norwegian Sea mean geostrophic transport of water warmer than 2 C through the same section relative to a level of no motion at 1000 m based on spring and summer data for and found it to be around 5 Sv. This includes an offshore branch over m depth in addition to the one over the shelf slope which is expected to have an additional barotropic component. Of particular interest for our study is the clear seasonal difference in location of the maximum current of the shelf slope branch. In March-April the maximum geostrophic current occurs over 500 m depth whereas in July-August it is displaced westwards to the 800 m depth contour (Mork and Blindheim, 1998). Blindheim et al. (1998) showed a striking negative correlation between the NAO index and the westward extent (varying over more than 300 km) of high salinity (S?:: 35) AW in May-June in the Norwegian Sea at I N, capturing both a long term increase and three oscillations in annual observations since the late 1960s. It is apparent that the structure of the NAC is variable and responding to large scale atmospheric forcing anomalies. The variations of the westward extent of AW in the Norwegian Sea give no direct information about the strength of the branch of the current along the continental slope. However Blindheim et al. (1998) suggest that reduced westward extent of AW in the Norwegian Sea should correlate with increased supply of AW both to the Barents Sea and the WSC. Drifter studies during (Poulain et al., 1996) have shown that the lower salinity Norwegian Coastal Current (NCC) and the Norwegian Atlantic Current (NAC) merge where the NCC is forced offshore near Lofoten at on. Vessel mounted ADCP and SeaSoar/CTD sections off Lofoten in May 1993 (Orvik et al., 1995) confirm that the NCC and at least the continental slope branch of the NAC appear as one current in this area. (Various hydrographic data show that there may be a baroclinic branch of the NAC further west over deep water also in this steep slope area, providing a continuous pathway from the deep ocean branch in the Svin~y section (Mork and Blindheim, 1998) to a similar offshore branch of the WSC further north. This offshore part feeds recirculation in the Fram Strait but appears to have no direct impact on the flow across the BSO.) Along the continental slope the current emerges as a swift, topographically steered current shedding eddies and offshore extending filaments. Upstream fluctuations of 3.5 to 5 day period seem to grow in this area primarily due to barotropic instabilities producing eddies and waves with km length scales (Orvik et al., 1995). Furthermore, the current is seen to partly intrude on the narrow shelf between 69 and 70 N. In this area, the continental slope gets gradually steeper. It could be that the steepening of the continental slope is a contributing factor for generation of flow on top of the shelf (Hill, 1995). Topographic irregularities not included in simple models like the one by Hill (1995) may also contribute to cross-isobath flow. The bottom topography shallower than 300 m is very complex between 69 and 71 0, so if the slope current interacts with the variable slope and generates circulation on the shelf, the dynamics of this process may be hard to observe in detail. We conclude that the flow in the area just upstream of the BSO, although guided by topography, may be quite variable on time scales of a few days. The flow contains both AW and a seasonally varying contribution from coastal water which is most obvious in summer. During winter, the fresh water contribution is smaller and more mixed in the vertical removing much of the baroclinicity often seen also in the southern part of the BSO in summer. 10

12 2 THE ROLE OF THE BARENTS SEA IN THE ARCTIC MEDITERRANEAN CLIMATE SYSTEM 2.1 Conditions in the Norwegian Sea Although not strictly upstream of the BSO, we mention here also some findings by van Aken et al. (1995) on the structure of the Arctic frontal zone bordering the NAC to the west at the approximate latitude where it passes the Bear Island and changes name to the WSC. van Aken et al. (1995) used data from 7 cruises between October 1988 and August High resolution zonal sections at I N gave geostrophic transports of the upper 1000 m of the water column between 4 0 E and 10 0 E relative to 1000 m ranging from 2.9 to 5.4 Sv with a mean of 3.8 Sv. This branch is related to the offshore branch in the Svinoy section having a mean transport of just less than 3 Sv in the data discussed by Mork and Blindheim (1998). The numbers from the two sections may differ for a range of reasons including different definitions of the waters included in the transport estimates, shifts between baroclinic and barotropic components of the transport, and different observation periods. van Aken et al. (1995) also used XBT surveys and absolute currents from vessel mounted ADCP with nondegraded Global Positioning System (GPS) navigation to study mesoscale variability in the area from 73 0 N to 76 0 N and from 2 0 E to 10 0 E. Warm core anticyclonic eddies and cold core cyclonic eddies were found along the frontal zone with scales increasing eastward with the increasing baroclinic Rossby radius of deformation, with mean eddy diameters of about 40 km west of 6 0 E and 60 km east of 6 0 E. The Arctic frontal zone as a whole showed meanders with wavelengths of more than 100 km. The scales are comparable to those observed by Orvik et al. (1995) near the shelf break further southeast. The match in scales may well be coincidental, however, since the mechanisms invoked by Orvik et al. (1995) to explain eddy generation near the shelf slope are different from those expected to be dominant in the baroclinic offshore branch. Also there are clear upper and lower limits to the scales that may be described by surveys such as those of van Aken et at. (1995) and Orvik et al. (1995) so in reality we have very little information about variability on scales below 10 km and above 100 km. We must conclude that the shelf slope and deep ocean branches of the NAC can behave quite uncorrelated unless they happen to be influenced similarly by correlated aspects of external forcing. The study by Blindheim et al. (1998) shows that the temperature and salinity in the deep ocean branch is negatively correlated with the NAO, while a positive, but weak correlation exists between NAO and temperature and salinity in the shelf slope branch. The continental slope in the Norwegian Sea west of the Bear Island Trough received strengthened attention after the wreckage of a Russian nuclear submarine approximately 200 km southwest of Bear Island in An overview of existing information about near bottom currents in the area is given by McPhee et al. (1998). Measurements by Blindheim (1994) in May-July 1993 indicated that the current from 667 m down to the bottom at m was virtually depth independent, mainly directed along the topography, and showed flow reversals at time scales of days and much energy in the 4-5 day band with velocities up to 30 cm/s and typical values above 10 cm/s. McPhee et al. (1998) measured currents at three moorings in the same area from September 1993 to October The dominant flow was along the slope with magnitudes up to 50 cm/s and periodicities of around 20 days while the cross-isobath velocities showed persistent and distinct oscillations on 5 day time scales, which are also present in the along-isobath flow but not with the 11

13 2 THE ROLE OF THE BARENTS SEA IN THE ARCTIC MEDITERRANEAN CLIMATE SYSTEM 2.2 Climatic state of the Barents Sea same statistical significance. McPhee et al. (1998) concluded that the observed oscillations could be interpreted as topographic Rossby waves moving northward with a phase speed of 0.2 cm/s. The measured currents were strongest in December to early June. 2.2 Climatic state of the Barents Sea Winter surface temperatures are representative of most of the water column and thus useful indicators of the climatic state of the Barents Sea (Midttun, 1990). The mean winter surface temperatures for presented by Midttun (1990) show a broad region with temperatures above 3 DC in the entire southwestern Barents Sea. The strongest temperature gradient occurs in the Polar Front close to Bear Island. In the central Barents Sea, the transition to temperatures below o C occurs over a distance of about 200 km, except in a zone 200 km from the Norwegian and Russian coast where temperatures above 0 DC extend further east associated with a coastal branch of warm water penetration eastwards in the Barents Sea and northwards west of Novaya Zemlya. In the Barents Sea Opening the warmest mean winter temperatures approach 5 DC between 71 and 72 D N (Midttun, 1990). The 11 year period which is the basis for this mean is reported to be slightly colder than a long term mean based on Russian data from We judge the gross spatial features quoted here to be representative of the period just before the observations described in section 3 and also a useful approximation to a longer term mean state. Midttun (1990) also calculated the mean winter sea surface temperature in the Barents Sea in each of the years 1977 to 1987 in the quadrant south and west of 75 D N, 41 E, and showed it to vary between 1.1 C in 1979 and 3.3 DC in 1983 with a mean close to the value of 1987 which was 2.3 DC. For these calculations a surface temperature of -1.8 DC was used in ice covered parts. The areal mean winter sea surface temperature is clearly negatively correlated with the maximum winter ice extent and positively correlated with the mean subsurface temperature in the Barents Sea Opening at m depth between 71 D 30 I Nand IN. The latter has an interannual variability of around 2 C and a mean seasonal variability of similar amplitude with a maximum in October and a minimum in April. The temporal resolution here is limited as no data from November or December were available for the analysis. Sea ice formation within the Barents Sea separates fresh water from the water column allowing it to be exported, in the form of ice or as near surface melt water, by mechanisms and through routes which could be quite distinct from those governing cold shelf water, recirculating Atlantic Water, or any of the intermediate water mass products. Midttun (1985) showed selected observations from the period 1971 to 1981 of high salinity, cold bottom water in the eastern part of the Barents Sea and north of Bear Island towards Storfjorden. Midttun (1985) sketched a multi-annual process whereby cold deep water from the previous winter could lead to conditions supporting stronger ice growth and more brine rejection the following year, so that the pool of bottom water would build up over several years. He also showed the existence of outflowing water north of Novaya Zemlya towards St. Anna Trough with temperatures down to below -1 C and salinities above in the summers of 1971, 1973 and 1974 for which data were available from the northeastern part of the 12

14 2 THE ROLE OF THE BA.RENTS SEA IN THE ARCTIC MEDITERRANEAN CLIMATE SYSTEM 2.3 Exchanges with the Arctic Ocean and the Kara Sea Barents Sea. Midttun (1985) further documented extensive distribution of cold water between Sentralbanken and Novaya Zemlya and from Novaya Zemlya towards Frans Josef Land in September There was only a small core of water with temperatures above 0 C centered around 100 m depth east of Sentralbanken. Further towards the northeast, the warmest water had temperatures between and 1.0 C and was found at depths between 150 and 250 m. Midttun (1985) suggested that the observed interannual variations in the thermohaline structure in the northeastern Barents Sea does not necessarily imply strong interannual variability of the dense water formation process, but may be caused by variations in the inflow from the west, and in particular that this inflow was weak in the cold year Martin and Cavalieri (1989) estimated the contributions from ice freezing in polynyas near Novaya Zemlya, Frans Josef Land and in the Kara Sea to formation of Arctic Ocean intermediate and deep water, using satellite data to map the evolution of the polynyas over the first part of winter in 1978 to Outflow of cold and saline water from the Barents Sea to the Norwegian Sea and the Arctic Ocean may gain an additional driving force from the suspension of bottom sediments augmenting the buoyancy forcing (Fohrmann et al., 1998). Modeling of such bottom plumes depends on empirical parameterizations of several small scale processes and contains considerable uncertainties, but the limiting factor for this mechanism to be of long term, climatic significance is probably the availability of resuspendable sediment (Fohrmann et al., 1998). On the basis of process modelling and field observations of sediment accumulation on the slope in the Norwegian Sea northwest of Bear Island, Fohrmann et al. (1998) suggested an order of magnitude for such transports from the Barents Sea and other shelves in the Arctic Ocean of 0.1 Sv or larger. 2.3 Exchanges with the Arctic Ocean and the Kara Sea Loeng et al. (1997) gave a review of water transport through the Barents Sea based on previous publications and current measurements in the strait between Novaya Zemlya and Frans Josef Land from October 1991 to September The measurements gave an annual mean outflow of 1.9 Sv and inflow of 0.3 Sv. The outflow showed a clear variation during the period with average in December through February of 2.7 Sv while in June through August the mean outflow was 1.2 Sv. A very recent study of the Eurasian Basin by Schauer et al. (1998) included data from the St. Anna Trough taken in summer Both from geostrophic calculations and vessel mounted ADCP, a bottom intensified outflow of 2 Sv was found, with temperatures below 0 C and salinities between 34.7 and 34.9, interpreted as originating from the Barents Sea and corresponding to the type of outflow measured during a full year by Loeng et al. (1997). Although the measurements of Schauer et al. (1998) are of short duration and only obtained during summer, the presence of the distinct outflow cores along the bottom and eastern slope of the St. Anna Trough confirm an outflow similar to that documented by Loeng et al. (1997). In addition to the variations in transport measured by Loeng et al. (1997) and interpreted as seasonal changes, interannual variations in the outflow are expected, both in the mass transport and its properties, depending on 13

15 3 ACOUSTIC DOPPLER CURRENT PROFILER OBSERVATIONS the extent of dense water formation in the Barents Sea during the preceding winter(s), expected to be flowing out driven by its negative buoyancy, and the variable contribution from direct wind driven and remotely driven throughflow (Midttun, 1985). Exchanges south of Novaya Zemlya and west of Frans Josef Land are also summarized by Loeng et al. (1997) and shown to be smaller in the mean. They would need to be included in complete mass budgets for the Barents Sea (Rudels, 1987), but in this study we neglect coastal water inflow in the southern part of the BSO and net outflow south of Novaya Zemlya as well as inflow of ice and (near surface) water from the Arctic and across Svalbardbanken, except for a brief discussion in connection with the heat budget in section 6.2. Throughflow of Atlantic derived water from the southwestern Barents Sea to the Arctic Ocean between Svalbard and Frans Josef Land has not been considered important by earlier investigators, partly because the water in the potential overflow regions in the shallowest parts of the Barents Sea is so modified that it can hardly be classified as AW (Pfirman et al. (1994), Steele et al. (1995)). However, we will return to the possibility of direct throughflow to the north in the discussions of interannual variability in section Acoustic Doppler Current Profiler Observations Shipborne Acoustic Doppler Current Velocity observations from the period are used. The data are obtained from IMR ships occupying the repeat section across the BSO in connection with climate and fishery investigations, figure 1. Data are collected while the ship is steaming as well as while doing station work and biological sampling. AU data have been collected with the same instrumentation based on 150 khz RD Instruments ADCP with bottom track, set to 10 minute averaging and 8 m vertical bins. Technical information about installation and data collection is given in Blindheim et al. (1988). Position for each profile is logged from the ship navigation system. A uniform quality check is applied to the data, rejecting individual bins based on the per cent good pings, and leaving blanks (no data) in bins for which the data quality is judged to be poor. Uniform criteria are applied for all data throughout. There is some variation in the time consumed for completing the section, and several data points are rejected in connection with station work and short detours from the main section track. The 10 minute averaging implies a maximum distance between consecutive profiles of 3-4 km, most often less depending on ship speed. Typically the upper few bins are rejected, occasionally more in conditions with high seas and strong winds, and the lower few bins are rejected, probably because of bottom reflections and reduction of signal strength in the deepest parts. The absolute horizontal velocity is logged with the ship speed being removed from the raw observations based on the bottom tracking in the usual way. Here we only focus on the velocity component transverse to the section. In order to compute transports we need to estimate velocities near the surface and bottom where original data have been rejected because of poor quality. We adopt a simple procedure with linear interpolation from the deepest accepted bin to zero velocity at the bottom depth (the depth is also recorded by the ADCP) and constant extrapolation from 14

16 3 ACOUSTIC DOPPLER CURRENT PROFILER OBSERVATIONS the shallowest accepted bin to the sea surface. We note that the latter extrapolation is likely to underestimate Ekman transports (Saunders and King, 1995), but make no attempt at correcting for missing intensification close to the surface, realizing that the Ekman depth is expected to be large in high latitude and weakly stratified waters, so that a major portion of wind driven transport should be included with the present procedure. Initial examination of the transverse velocity sections revealed several cases with cells of in and outflow across the BSO. With expected tidal currents in the area up to cm/s (Kowalik and Proshutinsky, 1995), and knowing that the typical time used by the vessel to complete any one section was 1-2 days i.e. close to the semi-diurnal and diurnal tidal time scales, tidal currents were suspected to be a major contributor to the observed structures. While tidal currents are obviously important for mixing processes in the area and interesting in their own right, a more pressing objective was to use the observations as basis for estimates of mean or filtered currents with residual tidal currents included, if significant, but the actual tidal currents at the time of sampling excluded. This was done by subtracting from each of the observed and vertically extrapolated and interpolated transverse velocity profiles, transverse tidal currents predicted from Gjeviks tidal model (Gjevik et al., 1994), and kindly supplied by Bj~rn Gjevik (pers. comm.). The tidal currents were taken from model output at the time of ADCP data collection using linear interpolation in space from the nearest grid points of the tidal model. The predicted barotropic tidal velocities are reduced towards the bottom in the same way as each of the ADCP profiles, i.e. linearly interpolated from the depth level of the deepest accepted ADCP bin to zero at the bottom. In this way we produce modified velocity sections which are used in the further analysis. It is to be noted that even if the direct tidal currents from the major tidal constituents, with correct phase and interference patterns according to the numerical tidal model, are removed by this procedure, the resulting modified velocities may still contain high frequency variability in the near-inertial frequency range. The modification process is therefore not a time filter in the ordinary sense, but it represents an attempt at removing the known components of the high frequency variability. Somewhat surprisingly, all major cell structures south of the Bear Island Trough remained after this modification. We consider the subtracted tidal currents to be realistic, and conclude that the observed structures in figure 2, see also table 1, are not caused by tidal aliasing associated with the ship motion through a few tidal periods. The major effects of the tidal correction are felt in the shallower parts of the slope of Svalbardbanken approaching the Polar Front. In this area the topography is steep. With assumed constant ship speed there are few ADCP data per bottom depth interval rendering the averaging process inherent in the basic ADCP data processing prone to errors. For this reason and because of the strong tides there, which may require an even finer resolution than Gjeviks model (Kowalik and Proshutinsky, 1995), we will not discuss data from less than 100 m depth close to Bear Island. This effectively excludes water north of the Polar Front but allows the AW domain to be studied in its entirety with some caveats for the region shallower than 200 m on the slope of Svalbardbanken. From figure 2 we note the unidirectional flow and extremely large transport in the very first data set. One might be suspicious of false velocities associated with incorrect ship velocity correction 15

17 3 ACOUSTIC DOPPLER CURRENT PROFILER OBSERVATIONS (a) October 1988 (b) September 1990 (e) September 1991 (d) October 1991 (e) March 1993 (f) October Figure 2: Transverse velo city measured by ADCP from crossings of the BSO at Limes given ill table 1 and co rrected [or tides as described in the text. Positive direction is inflow to the Barents Sea. 16

18 3 ACOUSTIC DOPPLER CURRENT PROFILER OBSERVATIONS (al Oc to ber 1988 (b) September (c ) September 1991 (d) Oct ober 1991 (e l!vlarch 1993 (f) Octo ber 1993 Figure 3: Temperature. Exact times are given III table 1 17

19 3 ACOUSTIC DOPPLER CURRENT PROFILER OBSERVATIONS '~ d (a ) Oct ob er 1988 (b) September 1990 '~d '~ d ( c) September 1991 (d) Octob er 1991 '~ d '"",d (e) l\'la rch 1993 (f) Octob er 1993 Figure 4: Salinity. Exact times are given in ta.ble 1 18

20 3 ACOUSTIC DOPPLER CURRENT PRO FILER OBSERVATIONS Section # Year.month. day. time Inflow Outflow Net Inflow Heat transport Sv TW a : b : : c : : d : : e : : f : : Mean Table 1: Time of data collection and transport through the sections or ADCP alignment errors shortly after the ADCP was installed. We have been able to go back to an even earlier data set collected during a test section approximately one month earlier with the same ship immediately after installation of the ADCP. The data from that section were not logged in the same way as later data because of incomplete data logging facilities onboard. However the data are shown in Blindheim et al. (1988) and display patterns more like some of the later sections, and no unidirectional flow. Thus we conclude that the data we show in figure 2 are all equally trustworthy. The extreme transport in the first displayed section, see figure 2, is believed to be a true snapshot of the situation at the time of sampling. Whether the transport in fall 1998 was on average higher than any of the later sampled periods is impossible to say from these ADCP data alone. It may also just be that the first crossing happened to be an extreme sample from a highly variable ensemble. The section coverage to the south and north is slightly variable between the 6 realizations, and we are missing the very northernmost part of section e) which was taken in March 1993, but the main Atlantic domain (most clearly delineated by salinity in figure 4) between 71 and 74 0 N is covered in all sections. We will not put any emphasis on the coastal water observed south of 71 o N where also highly variable topography may complicate interpretation. Common features of most parts of most of the sections are the barotropic dominance (remember we have interpolated towards enforced zero current at the bottom, so what may seem as vertical shear in the lower parts of some figures may be caused by this procedure) and the cell structures of in and outflow. The larger scale features are more obvious in some realizations than others, but it appears as a recurrent feature that the inflow on the southern side of the Bear Island Trough is predominantly concentrated in two main cores with return flow between them. An inflow core centered around 400 m depth, I N occurs in all sections except the first section which has inflow everywhere, and the fifth section which was taken in March 1993 and which has a core displaced slightly farther north. A second core further south also exists in all sections except the first one. It has a more variable position between 71 and 72 0 N at 200 to 300 m depth. The average distance between the two cores is around 150 km. These features and the outflow between them near 72 0 N invite interpretations in relation to topography. E.g. the NAC could be split by the topographic high slightly west of the section at 72 0 N, or there could be preferred length scales with semi-permanent 19

21 4 EXCHANGE PROCESSES eddy structures locked between Spitsbergenbanken and the banks close to Norway. On the slope of Svalbardbanken there is sometimes outflow and sometimes inflow with no apparent preferred flow direction (note our previous caveats about data quality in the shallower parts). The root mean square velocity is of order 20 cm/s, and the dominant length scale, in addition to the 150 km scale structures mentioned above, is so short (order 10 km) that it is barely resolved within the averaging time of the data logging. This spatial scale is shorter than that obtained on the mid-norwegian shelf from a larger data set collected with the same type of instrumentation during a campaign in March 1988 (Haugan et al., 1991), and implies that objective space time mapping of variability in the BSO would be difficult even with a dense net of sections. Temperature and salinity data obtained from CTD stations concurrently with the ADCP are shown in figures 3 and 4. We note the vertical and horizontal homogeneity in the only section obtained during winter time, and the shallowness of the stratified layer occurring in the summer/early fall occupations of the section. These are all well known features from many previously analysed hydrographic sections in the literature. The temperature and salinity sections are shown here for completeness and as reference for later discussions of this and other data sets. Computed heat transports based on 0 C reference temperature are shown in table 1. The significance of these mean values as well as the mean net volume transport in table 1 may be questioned in view of the observed variability. We will discuss the quantitative results further in section 6, but before that we turn to an overview of processes which may govern the observed exchanges, and revisit other published data in view of the information contained in the ADCP data set. 4 Exchange processes Loeng et al. (1997) discuss variability of the transport across the BSO based on running of the barotropic wind driven model of Adlandsvik and Loeng (1991) with realistic winds for the period The model will be discussed again in section 5, but it is appropriate to consider some results from the model at this point to initiate the discussion of processes of recirculation and throughflow, and to put the time period of our ADCP measurements into a longer perspective. The modeled flow across the BSO directly driven by local wind is largest in December with a mean of 0.4 Sv (Loeng et al., 1997). The standard deviation of the interannual variation in the November through March period is about 1.0 Sv, and about 0.5 Sv in May through September. Loeng et al. (1997) interpret the monthly variability of the modeled transport as a perturbation of the long term mean throughflow driven by other forcing mechanisms. The three winters of 1991, 1992 and 1993 gave the three largest monthly mean wind driven transports in the whole period from 1970 to 1994 with perturbations in one individual month above + 3 Sv in early 1991 and 1992 and above Sv in early The one-year average of monthly values peaked at Sv in the same three years, while the lowest one-year average in the whole period was reached in 1987 with Sv. The weak wind driven throughflow the year before our first ADCP data were collected, and the exceptionally strong throughflows driven by local wind during the winters of 1991 through

22 4 EXCHANGE PROCESSES should be borne in mind when interpreting the ADCP data. More generally, the variability due to local winds is a factor that needs to be considered in conjunction with all data sets from the area. Another important aspect of the model results is that the mean modeled throughflow for the whole period 1970 through 1986 was only 0.06 Sv (Adlandsvik and Loeng, 1991). This indicates that while the local wind may influence variability on daily to interannual time scales, it is not a significant driver of mean throughflow on decadal or centennial time scales. In this connection it should be mentioned that Hill and Lee (1957) investigated the relation between southerly winds and baroclinic geostrophic transports in the WSC immediately west of Bear Island. Based on 30 repeats of a section during 1949 to 1956 Hill and Lee (1957) established a positive correlation between the southerly wind component during the 10 days preceding the occupation of the section and the computed volume transport, but there was also a considerable component of the transport which was independent of the wind. The barotropic component is missing from the measurements of Hill and Lee (1957), but the results of their analysis can be taken as an indication that the WSC, in a manner similar to the Barents Sea branch of AW (Adlandsvik and Loeng, 1991), can be characterized as rapidly varying on time scales of days, with some of the variability governed by local winds. Jonsson (1991), Jonsson et al. (1992) and Morison (1991) described monthly and seasonal variations of the WSC correlated to wind power and wind stress curl. Tidal rectification is important over shallow banks, in particular Svalbardbanken (Kowalik and Proshutinsky, 1995), but can be neglected as a main driver of large scale circulation. Another local forcing beside the wind could be the densification of water due to cooling and brine rejection associated with ice freezing in the Barents Sea in winter (a process for which tides may be very important by creating openings in the ice). Open water regions are effective in water mass formation compared to ice covered regions due to the efficient heat loss from open water (Simonsen and Haugan, 1996), but it is in the seasonally ice covered part of the Barents Sea that the more extreme water masses are formed with temperature at the frezing point and elevated salinity. The density increase per unit surface heat loss is much larger when ice is freezing than when water is cooled towards the freezing point (Midttun, 1985), so the most efficient water mass formation in terms of density increase occurs when new ice is freezing. The formation of the denser fractions is important for the fate if not the rate of the outflow to the Arctic Ocean and the Norwegian Sea, in particular for its contribution to renewal of deep and bottom water in the adjacent basins (Jones et al., 1995). Following a review of conditions and processes in the Bear Island Trough and Hopen Deep in section 4.1 and evidence of throughflow in section 4.2, we discuss in section 4.3 both baroclinic forcing of large scale flow due to density gradients and processes for dense water escape from formation regions. In addition to the local processes and forcing factors, external pressure gradients are imposed on the Barents Sea. Sea surface slope across the barotropic component of the AW flow along the continental slope in the Norwegian Sea and large scale sea level differences between the Norwegian Sea and the Arctic Ocean/Kara Sea remain as major external forcing parameters for mean AW throughflow and possibly also for its variability. These factors are included in several of the modeling studies in section 5, and discussed again in section 7 in connection with long term 21

23 4 EXCHANGE PROCESSES 4.1 Circulation in the Bear Island Trough and the Hopen Deep variability, but should be kept in mind also when we now review local observational structures. 4.1 Circulation in the Bear Island Trough and the Hopen Deep Semtner (1987) in a coupled ice-ocean model of the Arctic Mediterranean obtained a cyclonic Barents Sea gyre with mean strength of about 7 Sv in the period December through March, and less than 2 Sv from May through September (his Fig. 11). About half of the gyre strength is associated with throughflow towards the Arctic Ocean and Kara Sea while the other half is local recirculation within the Barents Sea. The gyre was explained by conservation of potential vorticity and favorable forcing of the tangential component of the wind integrated around lines of constant f/h. The model was coarse resolution (110 km grid spacing) with corresponding coarse representation of the topography. Still it points to the possibility of a wintertime cyclonic circulation in the western Barents Sea. If such a circulation exists, it may interact with the Norwegian Atlantic Current and effect exchange of Atlantic Water across the Barents Sea Opening. The model simulation of Semtner (1987) had very coarse resolution. Our ADCP data, primarily from summer like most other observations, indicated recirculation but with a different distribution not resolved by the Semtner (1987) model. We now review other existing evidence about the flow pattern in the area. In a comprehensive overview of the Greenland-Iceland-Norwegian Sea, Hopkins (1991) also discussed exchange across the BSO. He interpreted the distribution of the 35 isohaline around the Bear Island Trough as indicative of a cyclonic recirculation that he named the Bj rn ya Trough Current. Such a recirculation was not apparent in dynamic anomaly maps discussed by Hopkins (1991), and no direct current measurements were shown, but Hopkins (1991) interpreted the presence of AW in the entire Bear Island Trough and Hopen Deep as circumstantial evidence for the inferred recirculation. Hopkins (1991) furthermore reviewed the evidence for outflow of dense water to the Norwegian Sea along the bottom of the Bear Island Trough and the slope of Svalbardbanken. The monumental work by Helland-Hansen and Nansen (1909) also deserves mentioning in this context. Even if they focussed on what they called the Norwegian Sea (presently known as the GIN Sea) they included the BSO, and presented a map of circulation at 100 m depth (their figure 106) which shows the main inflow through the BSO at , N and a cyclonic eddy or recirculation located over the deepest part of the Bear Island Trough between 72 0 Nand 74 0 N. The near surface drifter study by Poulain et al. (1996) showed a strong inflow to the Barents Sea close to the Norwegian coast and a weaker inflow in the Bear Island Trough. The drifters which entered the Bear Island Trough had highly circuitous paths and residence times of 1-2 months in the trough. The pattern with a distinct coastal branch is known from early estimates based on geostrophy from summer occupations of the section. We caution however against taking this branch to be representative of a large fraction of the AW inflow or of winter conditions when stratification is much weaker. Our ADCP data suggest that deep reaching cores of inflow are located further offshore and we propose that the surface intensified coastal branch is a seasonally varying phenomenon which is considerably weakened during winter. Johannessen and Foster (1978) investigated the hydrographic structure of the Polar Front close 22

24 4 EXCHANGE PROCESSES 4.1 Circulation in the Bear Island Trough and the Hopen Deep to Bear Island with 11 high resolution (2 km) XBT sections in July 1974, showing that the front was closely governed by topography and was locked at the outer part of the shelf approximately following the 100 m isobath. Ancillary data showed that the front experienced a perturbation of up to 10 km during a semidiurnal tidal cycle (Johannessen and Foster, 1978). Gawarkiewicz and Plueddemann (1995) and Parsons et al. (1996) used a 12 day time series from current meters moored at 380 m bottom depth on the northern slope of the Bear Island Trough at 22 0 E in August 1992, and vessel mounted ADCP in the shallower parts of the slope, to infer a westward mean along slope flow of Atlantic Water south of the Polar Front with a speed of about 0.1 ms- I. Gawarkiewicz and Plueddemann (1995) used the observations to support what they called a new circulation scheme for the western Barents Sea with eastward flow of Atlantic Water along the southern slope of the Bear Island Trough and splitting between Nordkappbanken and Sentralbanken. The shallower part of the inflow (above less than 260 m bottom depth) would continue southeastwards through the sill south ofsentralbanken, while the deeper part would follow the bottom topography cyclonically around the Hopen Deep and exit south of Bear Island. Based on the same current measurements and a range of additional data from August 1992, Parsons et al. (1996) presented an in-depth analysis of the summer time frontal structure, and described it to contain a tidally driven dynamic mixing regime situated above the 300 m depth contour and underneath the Polar Water upslope. The velocity structure is primarily barotropic. South of the front there is a baroclinic near surface inflowing component of 1-4 cm/s. Parsons et al. (l996) also noted in passing that transmissivity was reduced in the lower 100 m near the base of slope, indicative of suspended sediments, while generally the bottom boundary layer in the slope and front region was about 40 m thick. Harris et al. (1998) used archived hydrographic data from the southwestern Barents Sea from after 1950 in a study of Polar Front structure and water mass distribution. A total of 29 transects, each containing a minimum of three casts and crossing the Polar Front close to Bear Island, gave the possibility to synthesize a seasonal cycle of water mass structure near the Polar Front. The median number of transects per month is 2, and there were no transects from January or July in the data set. Several caveats are in order for this data set which is comprised of observations only at standard depths taken from different years in a region noted for its interannual variability (Blindheim, 1989), but the observed summer time temperature and salinity stratification and winter time virtual homogeneity is probably a generally valid attribute. A larger historical data set distributed over the entire southwestern Barents Sea was used by Harris et al. (1998) to demonstrate spreading of sea ice melt water over the Atlantic domain in summer and convective mixing of this water with AW, producing modified AW during fall and winter. Fresh water from coastal sources in mainland Norway was discarded by Harris et al. (1998) as a source for water mass formation in the western Barents Sea because of the dominant flow of these waters towards the east along the coast of Norway and Russia. Harris et al. (1998) further used the historical data set to study vertical profiles taken in September and April in an unspecified location within the Atlantic domain (with water depth 2:: 300 m) showing homogenization of salinity, and a depth averaged temperature drop of 3.7 C (from 23

25 4 EXCHANGE PROCESSES 4.1 Circulation in the Bear Island Trough and the Hopen Deep 4.7 C to 1 C) during the 7 months. This is close to the 4 C cooling that would result in a 300 m water column subjected to a 270 Wm- 2 surface heat loss (Hiikkinen and Cavalieri, 1989) during 7 months. Harris et al. (1998) noted that water parcels following the cyclonic recirculation path suggested by Gawarkiewicz and Plueddemann (1995) with a length of 1750 km, would use about 7 months if the velocity is 10 cm/s. Although this residence time was used by Harris et at. (1998) to argue that the observed seasonal cooling is consistent with cyclonic recirculation in the Bear Island Trough, we note that the approach has several shortcomings. First, the seasonal variability in temperature of the AW flowing in from the Norwegian Sea (see our March section in figure 3) would in itself induce seasonal variability in temperature at a fixed location in the trough. This influence would have to be filtered out from that ofthe surface flux along the path in the Barents Sea. Second, if the stable barotropic recirculation of Gawarkiewicz and Plueddemann (1995) dominated (their Fig. 13, see also section 5), there would be little interaction between the inflowing and outflowing branches. Then, by the end of winter one would expect a colder outflowing branch along Svalbardbanken than the inflowing branch on the southern side of the Bear Island Trough. The difference would not be as large as the 4 C proposed above, because of the upstream cooling of the inflowing AW from the Norwegian Sea, but it should be measurable, and would be an integral measure of the impact of the detour into the Barents Sea on the AW returning to the GIN Sea system. Large systematic northsouth gradients across the major parts of the BSO by the end of winter have however never been reported. There is a weak overall tendency of northwards cooling, and a dominance of much colder water near the bottom of the trough and along the deeper parts of the slope of Svalbardbanken, but this occurs in summer also (Blindheim, 1989) at a time when there would be negligible surface heat flux. The lack of consistent north-south temperature structure, except for the colder near bottom water in the north, shows, as do the ADCP data, that the exchanges across the BSO and the circulation in the Bear Island Trough must be different from and more complex than the scheme of Gawarkiewicz and Plueddemann (1995) which was largely based on short term summer observations supplemented by idealized numerical modelling. Blindheim (1989) showed current measurements from the Barents Sea Opening from September to October 1978, but had to use data from October 1978 to January 1979 over the north slope near the deepest part of the Bear Island Trough to arrive at composite transport estimates. The data cover the inflow branch in the south as well as the polar outflow on Svalbardbanken. A total inflow of 3.1 Sv and outflow of 1.2 Sv was estimated. Blindheim (1989) notes that the long distances between moorings may have smoothed out permanent features such as a countercurrent on the southern slope of the Bear Island Trough. Of the outflow estimated by Blindheim (1989), 0.8 Sv referred to water below the observed westward current minimum at about 100 m depth on the slope of Svalbardbanken. This water was named bottom water, and had temperatures between 0 and 2.5 C, salinity above 34.9 and (To above It was linked to high salinity, cold and dense water formed primarily over banks in the Barents Sea in the winter time and observed with (To ;::: 28.1 over large parts of the Barents Sea in different years (Blindheim, 1989). Blindheim (1989) pointed out the strongly barotropic character of currents in the Bear Island 24

26 4 EXCHANGE PROCESSES 4.1 Circulation in the Bear Island Trough and the Hopen Deep Trough with relatively high velocities close to the bottom, and with current reversals and instabilities or eddies possibly linked to the lateral shear between inflow and outflow. The interannual variability of temperature and salinity near the bottom of the Bear Island Trough is quite large. In the period 1968 through 1986 discussed by Blindheim (1989), temperature varied between 0 and 2.5 C and salinity varied between and The seasonal variability of the bottom water in the Bear Island Trough is much smaller than the interannual variability. It is obvious that the influence from interannually variable processes in the Barents Sea contribute to the observed variability in the Bear Island Trough, since the properties of the northward flowing AW in the NAC has a weaker variability (Gammelsr0d et at., 1992). Blindheim (1989) also showed data from the continental slope in the Norwegian Sea west of Bear Island, indicating outflow from the Bear Island Trough intruding at depths of 600 to 800 m below the main northward flow of AW in the area where the NAC changes name to the WSC. Adlandsvik and Hansen (1998) in a model study to be discussed in section 5, showed moored current meter data from the slope of Svalbardbanken. Five moorings with up to 6 recording current meters each were placed on a section from less than 50 m depth on Svalbardbanken between Bear Island and Hopen across the slope to more than 350 m depth in the Bear Island Trough from November 1987 to April Monthly mean currents for each of the four months November through February are shown in Adlandsvik and Hansen (1998). For all four months, there is inflow over depths larger than 150 m, and the southern limit of the inflowing warm core jet is captured at the deepest mooring. The largest monthly mean velocities in all four months are found in the upper levels of the mooring at 280 m depth and are between 5 and 10 cm/s. Monthly mean values near the bottom vary from 1 to 6 cm/s. Adlandsvik and Hansen (1998) do not calculate the total transport associated with the warm core jet, but from their contoured mean velocities, the jet is seen to represent a mean transport of order 0.5 Sv. When compared to the current meter data described by Blindheim (1989) which were obtained 9 years earlier and further west, there is a considerable difference between the outflowing structure described by Blindheim (1989) and the inflow described by Adlandsvik and Hansen (1998) over the same depth range of the slope of Svalbardbanken. It seems unlikely that the difference in east-west location along the slope can explain the differences. If both data sets represented a permanent flow pattern, the flows would need to be fed by water crossing the Bear Island Trough from the south somewhere between the sections. Such flow would strongly violate potential vorticity conservation when moving towards shallower depths at Svalbardbanken. Furthermore, the outflowing water in is so much colder than the inflowing water in , that the difference must be caused by a different history of water mass modification. It is therefore clear that considerable monthly and possibly seasonal and interannual variability exists of the currents and water masses south of the Polar Front along Svalbardbanken. We note that the volume budget of the Barents Sea based on Blindheim (1989) would loose an outflow of 0.8 Sv and gain an inflow of about 0.5 Sv if the fall 1978 data from the southern part of the BSO had been combined with winter data from (Adlandsvik and Hansen, 1998) rather than those originally used by Blindheim (1989) which were from the winter of We further 25

27 4 EXCHANGE PROCESSES 4.1 Circulation in the Bear Island Trough and the Hopen Deep note that 1979 was the coldest year and had the largest maximum ice extent of all years with modern data coverage of the Barents Sea, and that Midttun (1985) suggested that this cold state was linked to an anomalously weak inflow of AW in the preceding period, suggesting, as do the data of Adlandsvik and Hansen (1998), that the net inflow computed by Blindheim (1989) although a valid estimate of the situation in , is anomalously weak in a longer time perspective. Pfirman et al. (1994) used summer data from 1981 and 1982 to give a comprehensive description of the distribution of water masses in the northern Barents Sea. A hydrographic transect obtained along I N in 1981 showed high salinity cores (S 2:: 34.94) centered at about 75 m depth located above the m depth contours in the Hopen Deep (their Fig. 6B). Pfirman et al. (1994) interpreted these cores, which were found both on the slope of Sentralbanken and Svalbardbanken, as cyclonically recirculating modified Atlantic Water. They cited earlier summer current meter data in support of northward barotropic flow on the western flank of Sentralbanken, and older Russian literature in support of the recirculation along Svalbardbanken. Pfirman et al. (1994) classified the presumed recirculating water as Southern Barents Atlantic-derived Water (SBAW). This water class is characterized by a temperature and salinity maximum, and in late summer it is found with salinity from to and temperatures above 0 C at depths between 75 and 200 m. In the center of Hopen Deep it fills the whole water column. The temperature section across the Hopen Deep in summer 1981 (Pfirman et al. (1994) Fig. 6A) clearly shows that the recirculating core is in contact with colder water on the slopes while higher temperature prevails in the center of the trench. Cold Bottom Water (T :::; -0.5 C and (79 2:: 28.0) is present at about 225 m on the slope of Svalbardbanken. This water was presumed to be flowing westward. We note that a section across the Hopen Deep from Sentralbanken towards Svalbardbanken southwest of Hopen in July 1986 discussed by Quadfasel et al. (1992) (their Fig. 7) also shows two cores of high salinity Atlantic Water (this time above 35.05) similar to those discussed by Pfirman et al. (1994) and interpreted as cyclonically recirculating Atlantic-derived Water. It appears that the salinity and degree of modification of the water can vary from year to year, but the two topographically steered cores occur repeatedly. This recirculation comes in addition to the flow of cold bottom water, which due to its high density has no other escape route than through the Barents Sea Opening once it is in the trench at depths greater than 280 m. One must also expect interaction and mixing between the various modified water masses which recirculate close to each other. Quadfasel et al. (1992) compare salinities of Atlantic Water on the south and north side of the Bear Island Trough, and link the reduction in salinity from south to north to effects of mixing when recirculating Atlantic Water interacts with less saline near surface water in the Hopen Deep. The variability in properties of surface water as well as bottom water and A\V preclude firm estimates of mixing ratios based on temperature and salinity analysis. \Vhile many authors have argued for a cyclonic recirculation, as may be expected from conservation of potential vorticity and topographic steering, the evidence has been scattered and not always convincing. The direct current measurements which exist, show that AW in on the slope of Svalbardbanken tends to flow east (Adlandsvik and Hansen, 1998) while colder water in flowed west 26

28 4 EXCHANGE PROCESSES 4.2 Throughflow (Blindheim, 1989). The ADCP data from show transport in both directions. Westward currents dominated in a period in August 1992 (Gawarkiewicz and Plueddemann (1995), Parsons et al. (1996)). 4.2 Throughflow We start with estimates of net flow through the BSO and move progressively towards the northeast. The ADCP data give information about the structure of the flow, but the variability makes average flow estimates based on only 6 realizations rather useless as predictions of any mean state. Instead we will use the information from the ADCP data to critically assess previous longer term data and their interpretations. Loeng et al. (1997) review several Russian estimates of geostrophically estimated AW inflow to the Barents Sea across the BSO. The five different papers mentioned by Loeng et al. (1997) all give a clear seasonality with maximum from September to January and minimum from April to June. The highest geostrophic inflow is 2.1 Sv in December, at which time the mean of the five estimates is 1.9 Sv. The lowest inflow estimate is 1.2 Sv in June, at which time the mean is 1.4 Sv. It is noteworthy and perhaps surprising that these estimates which are based on data from different time periods all give similar seasonal features (It is also striking that Timofeyev (1961) obtained a very similar seasonal cycle of the baroclinic geostrophic transport of water warmer than 0 c and with salinity greater than 35 in sections across the WSC at 78 0 N based on data from 1933 to 1960 with au transports about twice as large as those discussed here for the BSO). The largest deviation between the studies for any given month is only 0.4 Sv. However, it is clear that the baroclinic geostrophic flow is only one component of the inflow. There is a need to account for the barotropic component of inflow, which may well be larger than the baroclinic. In Blindheims current measurements from , already discussed in section 4.1 (Blindheim, 1989), the net transport through the BSO was 1.9 Sv. Since we are focussing on AW, we should neglect the Polar Water, and subtract only the deep and bottom water outflow from the inflow, giving a net inflow of 2.3 Sv. We note that both 1978 and 1979 were cold years in the Barents Sea with 1979 giving an unusally extensive sea ice cover (Midttun, 1989), and that Midttun (1985) has proposed that the existence of large amounts of cold bottom water in the Barents Sea is indicative of a relatively weak AW throughflow. If we used the current meter data from the slope of Svalbardbanken (Adlandsvik and Hansen, 1998) together with the data of Blindheim (1989) from the region southwards of the deepest point in the Bear Island Trough, we would get an increase of 1.3 Sv and a total net throughflow of AW of 3.6 Sv. This combination of data from different years is of course quite questionable since we cannot exclude the possibility that the in and outflow cells shifted position and e.g. that the warm core jet from was compensated by return flow in an area where the data from show inflow. However, even the original analysis of Blindheim (1989) contained combination of data from different time periods. One of the moorings which was in place for the period June 1978 to March 1979, indicated that the inflow was stronger during the following winter period than during the main measurement period September to October 1978 (Blindheim, 1989). With the paucity of time series observations available, and 27

29 4 EXCHANGE PROCESSES 4.2 Throughflow knowing the costs and difficulties involved in obtaining representative data, we feel justified in trying to make the most out of the data which do exist. Caution has to be exercised since all the data are from a short period of time, and the spatial coverage is poor in relation to the variability on short spatial scales demonstrated in section 3. There is therefore a considerable uncertainty in these estimates. However, from the important if not dominating barotropic contribution of all measurements, it appears quite clear that the old geostrophic estimates centered around 1.5 Sv are not representative of annual mean througflow. Loeng et al. (1994) discuss current meter data obtained over bottom depths of250 m and 278 m between Sentralbanken and Storbanken from September 1992 to October 1993, and hydrographic data from the whole Barents Sea during the same years. AW with salinity above 35 and temperature above 2 C is present between the two banks in September of both years. However, hydrographic surveys during fall show that water temperatures both at 100 m depth and near the bottom drop rapidly to values below 0 C further north and east in the Barents Sea, indicating that this passage is close to the northeast extreme of the extent of warm AW in the Barents Sea in summer. On top of the banks and generally in the Barents Sea there was a much wider distribution of cold bottom water in fall 1993 than in fall The increased volume of cold bottom water was accompanied by an increase in the strength of near surface stratification in fall 1993 explained by more near surface meltwater after a winter with much ice freezing and brine rejection. A notable feature of the temperature and salinity distribution over the banks, in particular in 1992 (Loeng et al., 1994), is that the bottom water characteristica are much more evident in temperature than in salinity. The salinity is often lower in the cold bottom water than in the intermediate water above. This must indicate that the water column salinity during freezing ha.., been lower than that of the AW inflow (Harris et al., 1998) which subsequently intrudes above the bottom water on top of and along the slopes of banks. Thus the water column in summer/fall below the surface layer, may in fact be weakly unstably stratified in salt but is kept stable by the temperature difference between AW and bottom water. The current measurements of Loeng et al. (1994) showed that tides dominate, but monthly mean velocities up to 12 cm/s are observed with typical values from 3 to 5 cm/s. At almost all depths there were periods when the residual current changed direction in intervals of 3-4 days, and periods when the residual current direction was more stable, but the speed varied with a similar frequency. The dominating direction is towards the northeast and there is a clear seasonality in the current strength with a maximum in winter. Since there are no hydrographic data from the winter season, one can only speculate about the thermal modification of AW moving northeast between the banks in winter when the current velocity is stronger than in summer. The ice cover and salinity stratification at the base of the mixed layer limits upward heat loss, so possibly AW with temperatures above 2 C can penetrate further in winter than in summer. In this connection it is worth mentioning that Loeng (1991) concluded from different sections of the Barents Sea during many years that when cold water is displaced by warmer water, most of the associated local temperature increase takes place in late autumn or early winter. This can be understood by the stronger throughflow starting in late fall as seen in these current measurements and those from the 28

30 4 EXCHANGE PROCESSES 4.2 Throughflow BSO (Blindheim (1989), Adlandsvik and Hansen (1998)). Steele et al. (1995) showed hydrographic data from a drift section entering the Barents Sea in the north in early December 1988 and exiting via the Hopen Deep and Bear Island Trough in early January No brine enriched water was encountered, but water with properties matching those of cold halocline water of the Arctic Ocean was found primarily near the marginal ice zone in the Barents Sea, and explained to be formed by variable contributions from direct surface cooling of AW and from melting of sea ice. The cold halocline water formed in this way directly from AW in the Barents Sea would flow north and east (the latter supported by geostrophic currents from the section) towards the Arctic Ocean, rivalling the previously proposed, but not documented or localized sources of Arctic halocline water associated with ice freezing in shelf polynyas. The study of Steele et al. (1995) suggested a major role of the Barents Sea as source for the halocline water of the Eurasian basin, but documented no transport rates of such water. The data of Steele et al. (1995) do not exclude that freezing later in winter may produce more saline water which could renew deeper water of the Arctic Ocean rather than or in addition to halocline water. When discussing and discarding connections between the southern Barents Atlantic derived water and northern Barents Atlantic derived water of Pfirman et al. (1994), Harris et al. (1998) state (their page 2914) that they are not aware of a dynamical mechanism whereby AW can flow through the Polar Front region and over the sill between Svalbardbanken and Storbanken. We are not aware of current measurements at that sill, but the measurements of Loeng et al. (1994) over the sill between Sentralbanken and Storbanken do show flow of AW. The latter sill depth is 280 m while the former is 200 m, which would admittedly provide a stronger obstacle to flow. The deeper of the two sill depths is comparable to that south of Sentralbanken, so it is evident from the current measurements (Loeng et ai., 1994) that not all water above isobaths shallower than 280 m in the Bear Island Trough - Hopen Deep deflects eastwards south of Sentralbanken to feed the southern branch of AW throughflow. This shows that topographic control as invoked by Gawarkiewicz and Plueddemann (1995) can not explain the entire flow pattern in the area. If there is a large scale pressure forcing which creates a warm core jet along the slope of Svalbardbanken, it is expected that this same pressure forcing also generates northward or northeastward flow across the different sills. Strongly modified AW was shown by Steele et al. (1995) to have eastward baroclinic velocity components of up to 4 cm/s near the sill between Svalbardbanken and Storbanken. Northward baroclinic components or barotropic velocities are not known in this area. Based on ice maps and surface and subsurface temperature measurements, it is seen that the branch of AW which flows eastwards south of Sentralbanken normally is able to maintain high temperatures further east than the branch flowing through the Hopen Deep and north of Sentralbanken. In summer and fall, this is to be expected because of the higher temperatures in the parts with a larger contribution from coastal water with salinity stratification favoring seasonal surface heating. In contrast, the Hopen Deep branch may be fed primarily from the deeper levels of the AW in the Norwegian Sea with a weaker seasonal temperature cycle. In winter, the coastal branch should be the coldest upstream, but the effects of this may not be detected in hydrographic data which are mainly from the summer period, nor on the ice edge due to the delay in moving 29

31 4 EXCHANGE PROCESSES 4.3 Effects of buoyancy forcing from the BSO to the eastern Barents Sea. The average temperature at m depth in repeat sections crossing the southern branch of AW inflow east to 37 E are well correlated on interannual time scales with average temperature from the same depth interval in the southern part of the BSO (Midttun, 1989). A time lag is expected, but must be less than a year, and is difficult to determine from the yearly occupations of the sections. No direct current measurements are known from the southern branch. Schauer et al. (1997) recently documented the impact of the Barents Sea outflow through the St. Anna Trough on waters of the Eurasian Basin, basically confirming the scheme of Rudels et al. (1994). Based on data from summer 1993 in the Nansen Basin west and east of the trough, and supplementary information from the eastern Barents Sea, Schauer et al. (1997) deduced integral effects of the outflow on shelf slope mixing processes and formation of intermediate water in the N ansen Basin. The Barents Sea branch contained colder and less saline components than the Fram Strait branch. It mainly influenced the upper 1300 m of the Nansen Basin and constituted about 50 % of the upper ( m) intermediate layer of the boundary flow along the slope of the Arctic Ocean east of the St. Anna Trough and 80 % of the lower ( m) intermediate layers of this flow. This shows the importance of the Barents Sea throughflow for the structure and renewal of intermediate waters of the Arctic Ocean. The temperature, salinity and density range of the outflow from the Barents Sea to the Arctic Ocean is much wider than that of the inflowing AW in the BSO (Rudels et al., 1994), and the hydrographic properties of the outflow are expected to vary significantly from year to year. 4.3 Effects of buoyancy forcing We consider two ways in which local formation of dense water could affect circulation. The first is by creation of large scale baroclinic pressure gradients across some frontal zone between the dense and less dense waters creating geostrophically balanced flow along the front. The second is the possibility of cross frontal transport driven by the negative buoyancy of the dense water in combination with variable bottom topography. Midttun (1985) presented observations of dense water formed by ice freezing and salt rejection in the Barents Sea, with emphasis on the eastern Barents Sea and Storfjorden where extreme salinities well above the salinity of Atlantic Water have been found, also by other investigators including Anderson et al. (1988), Quadfasel et al. (1988) and Schauer (1995). Based on may years of investigation, Midttun (1985) stated that cold, saline bottom water is generally widespread over the Barents Sea. Midttun (1989) in particular emphasized Sentralbanken, where there is large interannual variability of sea ice cover, as a site for bottom water formation, but also mentioned Svalbardbanken as a source area for deep water in the Bear Island Trough. Transport away from the banks may occur rather slowly as indicated by the survival of cold bottom water many places in the Barents Sea, sometimes over several summers (Midttun, 1985) even if surrounded by warm AW. Quadfasel et aj. (1993) showed a section across the eastern part of Sentralbanken taken in August 1990 in which water colder than -1.5 C was present on 30

32 4 EXCHANGE PROCESSES 4.3 Effects of buoyancy forcing the southern slope of the bank close to the top of the bank. Water colder than -1.0 C occupied more than 120 km of the section at and near the bank, and water colder than -0.5 C extended about 200 km. On the northern slope, water above 0 C was present close to the bank, but to the south the warmer water (up to more than 3 C) was found much further away centered over the deepest depression. A maximum (J'(J of up to may be experienced in the main Barents Sea (Quadfasel et al., 1992) while the typical (J'(J of inflowing AW in the BSO is between 27.5 and These characteristica do not vary much over the seasons. If comparatively less dense AW meets denser water along a frontal zone with typical density difference 6p = 0.3kgm- 3, the associated geostrophically adjusted transport fr; H2 6p over a vertical depth scale H 200m becomes of order 1 Sv (g is the gravitational acceleration 10ms- 1, p is density of order 10 3 kgm- 3, and f is the Coriolis parameter of order 1O- 4 s- 1 ). The direction of the flow would be for near surface AW to move approximately eastward along isopycnals in the Barents Sea, which would contribute to throughflow, or for dense near bottom water to move westward contributing to recirculation if fed by inflow south of the idealized front. ~ote that the east-west idealization is a crude one. In reality, as just described, the front is strongly affected by bottom topography. A more quantitative estimate could be made by using detailed depths and locally observed density gradients. However it will remain only a diagnostic deduction from the observed density field, and does not fully explain the physical mechanism for the inflow or the extent of densification and buildup of density gradients. Dense water is thought to be preferrentially formed in shallow water when surface convection reaches the bottom and ice formation starts. The temperature of the dense water is tied to the freezing point, but the salinity could become higher or lower than that of AW depending on the conditions in fall and the total ice growth. Backhaus (1996) demonstrated in a model study the strong dependence of the rate of dense water formation in the ice covered parts of the Barents Sea on episodic wind events which could create openings of the ice cover in particular in the lee of islands, but the model results also showed strong sensitivity to the AW and freshwater content of the water column. The large winter air-sea heat fluxes which occur occasionally in leads and polynyas, are experienced throughout the winter in the Atlantic domain adjacent to the ice where also stirring by wind stress and Ekman pumping by stress divergence (~ed, 1983) affects the stability of the water column. So a density increase does occur also in the AW domain, and with its high salinity, pure AW would be a prime candidate for creation of dense outflow water if it could be cooled sufficiently. On the other hand AW does not easily move onto shallow banks due to topographic steering. For the outflow of strongly modified water created primarily over banks in winter, we need a mechanism for the water to escape (Kill worth, 1983). Cold, dense water on Sentralbanken may have long residence time and weak mixing with the anticyclonic circulation of Atlantic Water around the bank due to the well known Taylor column effect of banks. This mechanism has been suggested to be important also for localization of dense water formation near Maud Rise in the Weddell Sea (Alverson and Owens, 1996). In the first instance, elevated density of the cold water assisted by the thermobaric effect (increa..'le of compressibility with reduction in temperature) will 31

33 4 EXCHANGE PROCESSES 4.3 Effects of buoyancy forcing force a downflow along the slope of the bank, deflected to the right by earth rotation depending on the frictional influence. However, the dynamics are not so simple. While the formation of dense waters on shelves has been studied in several papers (e.g. Kampf and Backhaus (1998», the problem of collapse and subsequent downslope flow of dense water from banks appears to have received less attention. In an overview of processes involved in shelf convection, Backhaus et al. (1997) included a simple qualitative two layer model experiment of the rotational gravitational collapse of dense water on a uniform slope of 0.02 o. A representative water column of 100 m height and 45 km diameter, and some unspecified density contrast, was found to move with a speed of order 1 km per day, supporting the observations of long lived anomalies perhaps over a whole summer in the Barents Sea. However, we need a deeper understanding of the processes involved. Model studies of bottom arrested gravity plumes using specified upstream conditions and focussing on the active bottom layer alone (e.g. Jungclaus ei al. (1995), Alendal et al. (1994)) provide limited guidance on the coupled problem which requires upper layer convergence to fully or partially compensate the lower layer divergence over the bank. It is likely that both radiation by near-inertial waves and advection associated with baroclinic instability contribute to draining potential energy from a localized patch of dense water. This has been modeled with application to deep convection without the topographic interaction by Hermann and Owens (1990) and in several later model studies by other authors showing that a geostrophically adjusted state with a cyclonic rim current around the dense patch becomes baroclinically unstable resulting in mesoscale eddies. Adding interaction with topography would complicate the problem of the collapse of a cylinder (Dewar and Killworth, 1990), since the scale of the bank in relation to the baroclinic Rossby radius of the stratified water around the bank may be important. Non-axisymmetric perturbations are required since it is easy to see that restriction to axial symmetry yields only trivial solutions in a configuration with vertically homogeneous water on a cylindrical bank and continued densification over the bank (see however Gill et ai. (1979) for a two-layer axisymmetric case where densification is parameterized by localized mass transport from the upper to the lower layer). This implies that the joint effect of baroclinicity and relief (JEBAR) (Huthnance (1984), Sl~rdal and Weber (1996)) becomes another source term for the vorticity in addition to vortex stretching. Wheless and Klinck (1995) investigated a related problem with supply of dense water to a sloping shelf where the water column is everywhere assumed well-mixed, and showed in that case that the JEBAR is capable of transporting fluid across isobaths. The well known preferred anticyclonic circulation around banks, assumed critical in the localization of the dense patch in the first place (Alverson and Owens, 1996), may also affect the baroclinic instability process. Gjevik et al. (1994) show in both barotropic and baroclinic model cases, that longs he If currents near banks can generate anticyclonic flow not only around banks but also on top of the banks, i.e. precisely where the cyclonic rim currents generated by elevated density are expected. The most relevant and illuminating modelling studies for the present problem are probably those of Gawarkiewicz and Chapman (1995) and Chapman and Gawarkiewicz (1997). They used a semispectral three-dimensional numerical model with mesoscale resolution in studies where an area close to the coast was subjected to a negative surface buoyancy flux. With or without a bottom 32

34 5 PUBLISHED MODEL STUDIES slope, the modeled density anomaly in the patch increases linearly with time until a certain level, depending on the slope (Gawarkiewicz and Chapman, 1995), is reached. At that time instabilities develop into mesoscale eddies whose scale is determined by the baroclinic Rossby radius, and it is these eddies which mediate the offshore transport of dense water. Coates and Ivey (1998) used a laboratory model of a similar situation as that in the model of Gawarkiewicz and Chapman (1995) and confirmed the stabilization of the density anomaly in case of a coast ally bounded densification region. Chapman and Gawarkiewicz (1997) followed with further numerical modeling studies and showed that a scaling proposed by Visbeck et al. (1996) for deep ocean convection could be modified to the cases studied in the numerical model. Chapman and Gawarkiewicz (1997) came up with expressions for the equilibrium density anomaly as well as the time to reach this equilibrium when the forcing decay region is wider than the baroclinic Rossby radius. Both the density anomaly and the time depend on the efficiency of eddy exchange, and the numerical experiments indicate that this efficiency is several times larger in shallow convection than in open ocean deep convection. We note that the density contrast between the bank water and the surrounding A W may be quite small simply because of the larger contribution from surface meltwater residing over the banks in fall. This meltwater is mixed down in fall, reducing salinity in the whole water column before freezings starts. If there has been efficient removal of melt water during summer, e.g. by wind, then more saline water may occur as a result of ice freezing. This points to a possible forcing by the ice and freshwater circulation, expressed by the salinity in fall, on the degree of dense water formation during winter. Alternatively, the upper limit to density may be controlled by the mesoscale dynamical mechanisms just mentioned and as such be less dependent on the initial freshwater content. We conclude that the dynamics of establishing an equilibrium lateral density contrast in the Barents Sea is not properly understood, that it may depend on the initial freshwater content of the water column on the bank, the strength of AW flow (its degree of deflection around banks and its advective resupply of heat and buoyancy) in addition to the size and slope angle of the bank and the surface forcing. It appears as an interesting future modelling task to investigate the adjustment process and generation of mesoscale variability associated with dense water localized over a bank, and its possible dependence on background flow such as expected from AW in the Barents Sea. At present however, it seems that there is little that can be concluded on pure theoretical grounds beyond the discussion of mechanisms given here. 5 Published model studies The available observations are incomplete and require interpretation. Numerical models can guide such interpretation by facilitating quantitative analysis as well as qualitative identification of driving forces which warrant further consideration. Numerical models can also be used to test ideas and mechanisms suggested from theoretical studies. We mention here only model studies which have produced results of direct relevance to Barents Sea recirculation and throughflow of AW, recognizing that several published large scale ice and ocean models of the Nordic Seas and the 33

35 5 PUBLISHED MODEL STUDIES Primary Reference Throughflow Recirculation Northern slope Comment Hibler and Bryan, 1987 Unknown Unknown Outflow Coarse res. Semtner, Sv Sv Outflow Seasonal Adlandsvik...) Sv Not shown Not shown Wind only Harms, Sv None Inflow NAC and wind St~le-H..., Sv specified Not quantified Outflow Driven by BC Legutke, 1991 Closed Not quantified Outflow Cut NE corner Hiikkinen..., 1992 Broad Apparently none Inflow Sigma coord. McClimans..., Sv Unknown Inflow Driven by BC Gawarkiewicz..., Sv 0.9 Sv Outflow Idealized Engedahl, Sv Unknown Inflow Diagnostic Aukrust..., Sv None (?) Inflow (?) WSC 0(1 Sv) Holland..., Sv ~one Inflow Seasonal Weatherly..., 1996 Not quantified None (?) Inflow Too much ice Proshutinsky..., (1 Sv) Unknown Unknown Wind only Harms, Sv Not shown Not shown Sept.-March Gerdes..., Sv Apparently not Inflow High vert. res. Zhang..., Sv Variable Variable ,89-96 Engedahl..., Sv ~ 1 Sv Outflow Diagnostic Adlandsvik..., Sv Not quantified Warm core jet 4 km Sv.bank Li..., 1998 Not shown Apparently not Warm core jet 3 km Sv.bank Table 2: Overview of model studies Arctic Ocean, while adequate for the purposes for which they were designed, have not provided sufficient detail of the currents in the Barents Sea to be included here. Also pure tidal models, some process oriented local model studies and detailed studies of the GIN Sea or Arctic Ocean, are left out of the present overview. Some insights from a wider range of models have already been mentioned and other aspects are taken up in sections 6 and 7. The review here is restricted to published model studies available in the open scientific literature. There are large differences in model approach, numerical techniques, spin-up, forcing data, diagnostic or prognostic use of data etc. among the models. We can not give a full account of these aspects, but mainly mention relevant results and refer to the references for details. Table 2 summarizes the results of each model concerning throughflow, recirculation and the structure of the flow on the northern side of the Bear Island Trough. The models are treated in approximate chronological order, which to a certain extent but not completely coincides with increasing complexity and improved numerical resolution. Hibler and Bryan (1987) were among the first who coupled a 3D ocean model to a dynamicthermodynamic ice model of the Arctic Mediterranean. They showed that the ocean model, which was diagnostic in their case, was essential for providing the heat flux required to avoid unrealistic large ice cover in the western Barents Sea and the adjacent parts of the Norwegian Sea. With a grid spacing of they were unable to simulate shelf currents in any detail, but they noted that on the shelves in contrast to the deep ocean, fluctuating currents were dominant while mean currents were weak. With only three velocity grid points across the BSO, they showed (their Fig. 11) mean inflow in the southernmost and outflow in the northernmost part at the second level of the model centered at just over 50 m depth. 34

36 5 PUBLISHED MODEL STUDIES The study of Semtner (1987) has already been mentioned in section 2. His model had a slightly finer resolution than that of Hibler and Bryan (1987) allowing the simulation and analysis of a Barents Sea throughflow and recirculation and identification of its strong seasonality with maximum strength in winter. The modeled cyclonic gyre in the Barents Sea except the throughflow is classified as recirculation here although parts of it may actually be a closed circulation east of the continuation of the NAC as the WSC. Adlandsvik and Loeng (1991) used a barotropic model to investigate wind driven transports using realistic wind forcing for the period and no other external forcing. The mean throughflow for the whole period was 0.06 Sv (Adlandsvik and Loeng, 1991). As mentioned in section 2 this indicates that while the local wind may influence variability on daily to interannual time scales, it is not a significant driver of mean throughflow on decadal or centennial time scales. Model runs for the whole period up to 1994 are reported in Loeng et ai. (1997). Harms (1992) used a multi-level barotropic model forced by a sea surface slope at the southern boundary at 67 0 N with an inflow of 3.6 Sv chosen to represent the barotropic component of the N AC. Of this inflow to the model domain, 0.15 Sv flows into the Barents Sea. The inflow is concentrated along the northern side of the Bear Island Trough from a southward going flow west of Bear Island. The small transport was taken to indicate that wind and baroclinicity would be required to obtain a substantial flow into the Barents Sea. However also the manner in which the upstream condition was specified as a linear sea surface slope independent of topography may have some importance. The inflow in the Bear Island Trough remained when M2 tides were included producing a residual current in the opposite direction on the shallow bank north of a current front approximately over the 200 m isobath, which was interpreted as coinciding approximately with the observed Polar Front. Forcing with monthly mean winds for the period gave a mean annual cycle of transport through the Barents Sea Opening with 0.5 Sv in the winter, 0.1 Sv in summer and an annual average of 0.3 Sv, which has been counted as the modeled throughflow in table 2. The largest fraction flows out between Novaya Zemlya and Frans Josef Land but there may also be some outflow south of Novaya Zemlya and west of Frans Josef Land. The model results for the pure wind driven case are in good agreement with those of Adlandsvik and Loeng (1991) using a similar model. The mean throughflow in the model of Harms (1992) is low, and does not alter the conclusion from the Adlandsvik and Loeng (1991) model study that long term mean throughflow must be governed by other forcings than the local wind, which however can generate large interannual and shorter term variability. It is somewhat surprising that the NAC inflow produces so small throughflow. St0le-Hansen and Slagstad (1991) presented a regional baroclinic model with 20 km horizontal resolution coupled to a thermodynamic free drift ice model. It was forced by an inflow with uniform speed of 0.1 ms- 1 in the NAC and 0.25 ms- 1 in the Norwegian Coastal Current south of the BSO and a weak inflow of surface water between Frans Josef Land and Svalbard. The surface elevation between Frans Josef Land and Novaya Zemlya was adjusted to give an outflow of about 2.0 Sv, but this was done in such a manner that waves and disturbances inside the model domain should be allowed to pass through the boundary undisturbed. There was an open boundary also to the west 35

37 5 PUBLISHED MODEL STUDIES of Spitsbergen. The model was initialized with hydrographic data obtained by IMR in fall Steady state currents balancing the initial density field and the boundary conditions with no surface forcing were computed, showing outflow along the slope of Svalbardbanken and two branches of inflow across the BSO; one coastal branch and one branch situated over the southern slope of the Bear Island Trough. A fraction of the cyclonic recirculation in the Bear Island Trough and the Hopen Deep formed a closed circulation in the trench without escaping to the WSC at Bear Island. The model was then forced with simplified surface cooling, producing apparently realistic ice cover and temperature, salinity and density profiles in different parts of the Barents Sea during winter. There was little change in the structure of the currents associated with the seasonally varying density field. From 1 March, realistic winds and more complete representations of the surface heat budget was included. Ice melting and development of spring and summer stratification was then obtained. The main discrepancies with observations were explained by missing tidal mixing over shallow banks. The same basic model on a larger domain, but still with similar specified in an outflows and similar throughflow pattern, has been used by Slagstad and Wassmann (1996) for simulations of a cold year (1981) and a warm year (1984) in studies with emphasis on vertical stability and forcing of biological processes. Legutke (1991a) presented an ocean general circulation model of the Greenland and Norwegian Seas with 20 km resolution and 12 levels, of which 6 to 9 are active over most of the Barents Sea. Additional results from the same model are also presented in Legutke (1987) and Legutke (1991b). The model had an inflow of 5.3 Sv through the Faroe-Shetland Trough, in and outflows of 7 Sv through the Fram Strait and 0.6 and 5.9 Sv through the Denmark Strait, but was closed at the eastern and northern boundaries of the Barents Sea. Climatological initial data was used. Initial experiments with seasonal as well as annual mean wind stress forcing gave an inflow across the BSO, split into one coastal and one offshore branch with recirculation in the Bear Island Trough (Legutke (1987), Legutke (1991a». Simulations with synoptic winds from January 1976 to June 1978, which was a period of anomalously low wind stress curi, were then performed. The time mean transport of Atlantic Water off the mid-norwegian shelf was 10.9 Sv with a realistic distribution in different branches over topography in comparison with measurements (Mork and Blindheim, 1998). The closed boundary in the Barents Sea precludes direct application of the model results on the main topics of the present paper, but the in-depth analysis of contintental shelf waves in the Norwegian Sea (Legutke, 1991a) is relevant to the issue of the effects of wind variability on transport in different branches of the NAC. (Stevens (1991) is an example of a related study discussing several aspects of boundary conditions and forcing of the circulation in the Norwegian and Greenland Seas. The model of Stevens (1991) had the eastern boundary at 30 o E in the Barents Sea and gave a cyclonic recirculation in the Bear Island Trough. We have not included it in the table, because of the limited realism in the Barents Sea with this boundary.) In Legutke (1991a) the Norwegian Sea shelf slope branch of the NAC has higher mean transport and variability in the winter months than in summer, and its transport is well correlated with the contour integral of the wind stress along the 1000 m isobath enclosing the whole model domain on all resolved time scales. The second branch which is centered over the 1000 m bottom depth 36