The Arctic Circumpolar Boundary Current

Deep-Sea Research II 46 (1999) 1023}1062

The Arctic Circumpolar Boundary Current Bert Rudels *, Hans J. Friedrich
, Detlef Quadfasel
 Finnish Institute of Marine Research, Lyypekinkuja 3A, P.O. Box 33, FIN-00931 Helsinki, Finland
Institut fu( r Meereskunde der Universita( t Hamburg, Troplowitzstra}e 7, D-22529 Hamburg, Germany Received 3 April 1998; received in revised form 15 September 1998; accepted 10 October 1998

Abstract The main water transformations in the Arctic Mediterranean take place in the boundary current of Atlantic Water, which crosses the Greenland}Scotland Ridge from the North Atlantic into the eastern Norwegian Sea. It enters and #ows around the Arctic Ocean before it exits the Arctic Mediterranean as the East Greenland Current, primarily through Denmark Strait. On route, it experiences numerous branchings and mergings. By examining how the properties of this `circumpolara boundary current evolve, it is possible to identify and describe the processes causing the water mass transformations in the Arctic Mediterranean. It is also possible to follow the Arctic Ocean deep waters as they spread into the Nordic Seas and eventually provide 40% of the over#ow water supplying the North Atlantic Deep Water.  1999 Elsevier Science Ltd. All rights reserved.

1. Introduction The Nordic Seas } the Greenland Sea, the Iceland Sea and the Norwegian Sea } are part of the Arctic Mediterranean Sea (Sverdrup et al., 1942), which also comprises the Arctic Ocean with its two main deep basins } the Eurasian and the Canadian Basin } and its extensive shelf seas. The circulation, the water mass transformations and the ventilation of the Nordic Seas are conditioned by the circulation in the entire Arctic Mediterranean. Apart from its restricted communication with the Paci"c through the Bering Strait, the Arctic Mediterranean forms a high latitude cul-de-sac of the North Atlantic. The communication with the North Atlantic occurs mainly over the

* Corresponding author. Fax: #358-9-331-025. E-mail address: rudels@"mr." (B. Rudels)  Present address: Niels Bohr Institute for Astronomy, Physics and Geophysics, University of Copenhagen, Juliane MariesVej 30, DK-2100 Copenhagen, Denmark. 0967-0645/99/$ - see front matter  1999 Elsevier Science Ltd. All rights reserved. PII: S 0 9 6 7 - 0 6 4 5 ( 9 9 ) 0 0 0 1 5 - 6

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Greenland}Scotland Ridge, but a second out#ow passage, through the Canadian Arctic Archipelago and the Davis Strait, exists west of Greenland. The circulation is dominated by the movement of warm Atlantic Water entering across the eastern part of the Greenland}Scotland Ridge into the Norwegian Sea (Helland-Hansen and Nansen, 1909). It #ows along the Norwegian coast as the Norwegian Atlantic Current. When it reaches the latitude of the Bear Island Channel, its "rst major bifurcation occurs. A substantial fraction #ows eastward and enters the Barents Sea (Helland-Hansen and Nansen, 1909; Nansen, 1906), while the main part continues northward as the West Spitsbergen Current. Several branches are de#ected westward from the current: north of the Greenland Sea basin, north of the Boreas basin and in Fram Strait (Gladfelder, 1964). Only a smaller part of the West Spitsbergen Current eventually enters the Arctic Ocean and #ows eastward along the Eurasian continental slope. North of the Kara Sea the boundary current meets the branch that turned east and entered the Barents Sea north of Norway. This branch reaches the Arctic Ocean by crossing the Barents Sea and the northern part of the Kara Sea (Rudels et al., 1994). The combined boundary current continues eastward a short distance before it again splits. Branches leave the continental slope along bathymetric features, particularly along the Nansen-Gakkel Ridge (Anderson et al., 1989), the Lomonosov Ridge (Aagaard, 1989) and the Mendeleyev Ridge (Aagaard et al., 1996; Carmack et al., 1997). However, a part of the boundary current follows the continental slope around the entire Arctic Ocean. As this part recrosses the Lomonosov Ridge into the Eurasian Basin it meets and mixes with the other branches as they converge east of the Morris Jesup Plateau. The waters exit the Arctic Ocean through Fram Strait, where they combine with the recirculating waters of the West Spitsbergen Current to continue southward along the Greenland continental slope as the East Greenland Current. The boundary current again diverges at bathymetric features, in this case the Greenland Fracture Zone and the Jan Mayen Fracture Zone, and branches from the boundary current enter the interior of the Boreas Basin and the Greenland Sea Basin (Helland-Hansen and Nansen, 1909). Exchanges in both directions occur, and the East Greenland Current is resupplied with water masses formed in the subpolar seas. The main part of the boundary current exits the Arctic Mediterranean through the 600 m deep Denmark Strait, but its denser fractions are de#ected eastward along the Jan Mayen Fracture Zone and along the Iceland shelf slope and eventually enter the Norwegian Sea. The upper part of these waters then returns to the North Atlantic through the 850 m deep Faeroe-Shetland Channel. The Atlantic Water experiences large transformations due to its exposure to the severe high-latitude climate and due to the net freshwater input to the Arctic Mediterranean. The varying bathymetry allows di!erent processes to dominate and the separate circulation branches attain di!erent characteristics. This leads to strong frontal zones and isopycnal mixing as the branches again converge. It is our purpose to follow the `Arctic Circumpolar Boundary Currenta as it circulates around the Arctic Mediterranean. It is an e!ort to examine how the processes in the di!erent regions determine the water mass transformations, how the di!erent regions are connected, and how important they are for the formation of over#ow water and for the supply of North Atlantic Deep Water. The hydrographic stations and sections

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Fig. 1. A guide map to positions of CTD-pro"les used in this text. Bold lines at 803N and at &203E refer to the sections of Figs. 2 and 3. Shaded dots give the CTD-positions of pro"les used for Figs. 9 and 10. Full dots give CTD positions for of those pro"les used for the Figure with the number displayed nearest a cluster of such dots. Open circles refer to the CTD-positions of Fig. 13, stars to those of Fig. 7. Table 2, includes additional information on CTD-station time and position.

discussed in the text are shown in Fig. 1 and the stations positions are given in an appendix (Table 2). 2. The Norwegian Sea and the Barents Sea In the Norwegian Sea the heat #ux to the atmosphere is large. The Atlantic Water is cooled, and in winter convection reaches down to 400}600 m (Clarke et al., 1990). Because of net precipitation the Atlantic Water is not only cooled but also becomes less saline. The e!ects of cooling dominate and the Atlantic Water becomes gradually denser as it #ows northward. Because this density increase is so signi"cant, Bunker and Worthington (1976) postulated that the northern Norwegian Sea was the likely formation area of the North Atlantic Deep Water, an idea has recently been taken up by Mauritzen (1996a,b). A small part of the Atlantic Water is detached toward Jan Mayen but the course of the Norwegian Atlantic Current mainly runs parallel to the Norwegian continental

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slope. It continues north of its main bifurcation at the Bear Island Channel as the West Spitsbergen Current. The current then consists of several bands, and as it approaches Fram Strait parts of the current recirculate to the west, north of the Greenland Sea (Gladfelder, 1964), and inside as well as north of Fram Strait (Bourke et al., 1988). Only a smaller fraction of the West Spitsbergen Current then passes into the Arctic Ocean (Bourke et al., 1988). Most of it enters close to Svalbard but some may #ow around the Yermak Plateau (Perkin and Lewis, 1984) as well as through a recently discovered gap in the central Yermak Plateau (Gascard et al., 1995). The recirculation in the strait leads to a westward spreading of the Atlantic core, and warm Atlantic Water dominates the upper part of the strait (Fig. 2). As Atlantic Water enters the Arctic Ocean its high surface temperature and its large reservoir of sensible heat lead to the melting of sea ice and to the maintenance of an ice free area north of Svalbard (Whalers' Bay), promoting a large heat loss to the atmosphere (Aagaard et al., 1987; Untersteiner, 1988). The melt water becomes mixed into the cooled upper part of the Atlantic Water, creating a less saline surface layer. This melt water input has also been observed from Oxygen measurements (OG stlund and Hut. 1984; Schlosser et al., 1994a,b), which showed that the less saline surface layer in the Nansen Basin largely results from ice melt. In winter the surface layer is homogenised by haline convection down to the thermocline and no cold, isolating halocline is present above the Atlantic Water (Rudels et al., 1996). The Atlantic Water as well as the surface layer #ow eastward in the boundary current along the Eurasian continental slope. The Atlantic Water does not subduct below the Polar Surface Water, as it enters the Arctic Ocean. Its upper part becomes transformed into a deep ('100 m), less saline surface layer which, together with the permanent ice cover, partly insulates the Atlantic Layer from the atmosphere. North of Norway, the Atlantic Water enters the Barents Sea together with the Norwegian Coastal Current (Fig. 3). This in#ow splits. One part continues eastward south of the Central Bank, the other #ows west of the Central Bank into the Hopen Deep. This branch again splits in the northern part of the Hopen Deep. One part returns to the Norwegian Sea along the Svalbard Bank, a smaller fraction crosses the sill between EdgeoK ya and the Grand Bank into the northern Barents Sea, and the rest #ows east north of the Central Bank (Loeng, 1991). The returning Atlantic Water has become denser by cooling and less saline, partly by net precipitation, partly by exchanges of Arctic water across the polar front (Loeng, 1991; Harris et al., 1998). The heat loss in the Barents Sea is as large as in the Norwegian Sea (Vowinckel and Orvig, 1970), and since the depth of the Barents Sea (200}300 m) is less than the level reached by winter convection in the Norwegian Sea the cooling here is stronger. The exchanges across the Polar front also bring Arctic water on top of the Central Bank, and the densest water found at the bottom of the Hopen Deep is likely to have formed by haline convection over the Central Bank and then drained down the sides of the bank into the deeper depressions (Midttun, 1985; Loeng, 1991; Quadfasel et al., 1992). This less saline, but cold and dense bottom water provides the densest return #ow down the Bear Island Channel into the Norwegian Sea (Fig. 3). In the other branch, the water of the coastal current, #owing over the shallower, shoreward area, has become homogenised by convection and cooled to freezing

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Fig. 2. Zonal sections of # and S along 803N in Fram Strait. (a) RV Polarstern in summer 1985, (b) RV Polarstern in summer 1987. Greenland is to the left, Svalbard to the right (from Chrubbasic, 1993).

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Fig. 2. Continued.

temperature, when it reaches the eastern part of the Barents Sea. Subsequent heat loss leads to ice formation and brine rejection, and cold, dense bottom water is formed over the shallow banks (Nansen, 1906; Midttun, 1985). It sinks down the sides of the banks into the depressions, where it mixes with and cools the Atlantic Water. As the winter progresses the density of the water formed on the banks increases and by the

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Fig. 3. Meridional sections of #, S and r at the western Barents Sea shelf break. Norway is to the left,  Svalbard to the right.

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end of the winter its density is high enough for it to sink into and "ll the deepest depressions. This has been observed throughout the years (Nansen, 1906; Midttun, 1985; P"rman et al., 1994) as well as explored analytically (Defant, 1961, p. 138; Midttun, 1985) and numerically by recent model experiments (Backhaus et al., 1997; Harms, 1997). The work by Backhaus et al. (1997) stresses the importance of lee polynyas and the removal of ice for reaching really high densities. The ice melts in summer by incoming solar radiation, and by heat from below if it drifts over the warmer Atlantic Water. This leads to the formation of a colder and less saline surface layer, referred to as the Arctic Water in the Barents Sea (Loeng, 1991), where it constitutes the surface layer north of the Polar front. It is homogenised in winter by haline convection down to depths larger than 100 m. Although its salinity is higher than that of the surface water in the Arctic Ocean, the convection in the northern Barents Sea does not penetrate into the underlying Atlantic Water. Yet much of the Arctic Water supplies the less dense, cold out#ow in the Bear Island Channel to the Norwegian Sea (Loeng, 1991; see also Fig. 3). Nansen (1906) believed that the dense bottom water formed in the eastern Barents Sea would follow the main #ow between Franz Josef Land and Novaya Zemlya eastward into the Kara Sea. This assumption is corroborated by hydrographic sections (Midttun, 1985) and by year-long current observations (Loeng et al., 1993) in the passage between the two seas, as well as by model studies (Harms, 1997). Part of this water, presumably the less dense fraction, will reside in the Kara Sea as a receptor for the run-o! from the large rivers entering the Kara Sea. The main and denser part, however, continues down the St. Anna Trough as a cold, dense wedge on the eastern side of the trough beside the core of warmer Atlantic Water from Fram Strait, which makes a loop into the St. Anna Trough (Hanzlick and Aagaard, 1980). The assessment of the relative importance of the two pathways for Atlantic Water from the Norwegian Sea to the Arctic Ocean has varied throughout the century. Helland-Hansen and Nansen (1909) believed that the main in#ow occurred over the Barents Sea and that the West Spitsbergen branch was so reduced by the westward recirculation that only a small fraction of it entered the Arctic Ocean. The West Spitsbergen Current in Fram Strait was thus considered much weaker than the East Greenland Current leaving the Arctic Ocean. Coachman and Aagaard (1974) assumed the in#ow over the Barents Sea to be small, and in their mass budget for the Arctic Ocean Aagaard and Greisman (1975) postulated that the transports of the East Greenland Current and the West Spitsbergen Current were in balance and almost an order of magnitude larger than the exchanges through the other passages. Now the views again have changed. In their recent freshwater budget for the Arctic Ocean Aagaard and Carmack (1989) assume that the in#ow over the Barents Sea is perhaps twice that through Fram Strait (2 Sv vs. 1 Sv).

3. The Eurasian Basin The boundary current #owing along the Eurasian continental slope interacts with the shelf waters. Atlantic Water enters the Barents Sea in the Victoria Channel

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(Mosby, 1938), and P"rman et al. (1994) argued that this in#ow from the north supplies most of the cold, less saline Atlantic Water found in the deeper parts of the northern Barents Sea. Atlantic Water from Fram Strait also enters the St. Anna Trough (Coachman and Barnes, 1963; Hanzlick and Aagaard, 1980). Most of this water just loops into the trough, but a fraction may continue and circulate around Franz Josef Land to enter the Barents Sea from the east (Hanzlick and Aagaard, 1980). Dense and cold shelf water from the Barents Sea, draining down the slope and penetrating into the Atlantic Layer of the boundary current, also has been observed (Rudels, 1986; Schauer et al., 1997). The largest changes in the boundary current occur when the two in#ow branches, the Fram Strait branch and the Barents Sea branch, meet in and east of the St. Anna Trough. Because of their di!erent mixing histories, the characteristics of these two branches have become quite distinct. Below the upper layer the Fram Strait branch has been largely protected from interactions with the ice and with the atmosphere, and it has retained much of its initially high temperature. Some cooling has taken place though, and Swift et al. (1997) estimate a heat loss of 80 W m\ from the warm core along the Barents Sea slope. Since the interaction with the sea surface normally involves entrainment into the upper layer rather than a cooling of the water below, this observed cooling could be caused by double-di!usive heat #uxes through diffusive interfaces, which do not involve any downward mixing of mass. Another possible explanation would be isopycnal mixing with water of the interior basin. The salinity maximum is still present in the Atlantic Layer but it is found deeper here than the temperature maximum; and the Fram Strait branch is strati"ed primarily in temperature (Schauer et al., 1997; see also Fig. 4). This is in contrast with the much colder and less saline Barents Sea branch. It is almost homogenous in temperature (#&!0.5) but strati"ed in salinity (see Fig. 4). The density range of the initially 300 m thick out#ow from the St. Anna Trough corresponds to a density range found over 800 m in the Fram Strait branch. As the Barents Sea water enters the Nansen Basin the denser part slides down the slope while the less saline upper part becomes con"ned to the shallower part of the slope. This gives the strange impression of almost similar potential temperature pro"les and largely di!erent salinity pro"les across the slope, as is seen on stations taken north of Severnaya Zemlya (Fig. 4). The two branches here #ow side by side, separated by a sharp frontal zone. Large and mostly density-compensating inversions are seen in the potential temperature and salinity pro"les indicating strong isopycnal mixing (Fig. 4). The inversions are connected with sharp gradients separating almost homogeneous layers. The almost densitycompensating property steps between the layers allow double-di!usive convection to operate, which could drive the water in the layers across the front into the opposite water mass and expand the frontal zone into the two branches. Layer structures are observed also in the interior of the basin (Fig. 5). Velocities in similar layers, observed in isolated lenses of Mediterranean Water in the North Atlantic, are found to be small, about 0.001 m s\ (Ruddick and Hebert, 1988). If such velocity also applies for the layers found in the Arctic Ocean, it would be too small to account for the observed wide spreading of the layer structures. Quadfasel et al. (1993) and Rudels et al. (1994) therefore assumed that the boundary current partly becomes

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Fig. 4. CTD-pro"les obtained in 1995 by RV Polarstern north of Severnaya Zemlya. (a) #}S curves, (b) #}S curves (blow-up of the deep data), (c) potential temperature pro"les (#), (d) salinity pro"les (S) (cf. Fig. 1).

detached from the continental slope at the Nansen-Gakkel Ridge. Such a separation of the boundary current between the Barents Sea slope and the Laptev Sea slope is indeed suggested by the observations by Schauer et al. (1997). The fraction that returns along the ridge is dominated by water of the Fram Strait branch, but as it turns this branch has already developed a layering structure. Because of the broadening of the boundary current during this separation, the extent of the layering would increase more rapidly than by motions induced by double-di!usive convection. As another bifurcation of the boundary current occurs at the Lomonosov Ridge (Aagaard, 1989), the layers could extend over a large part of the Amundsen Basin and be advected with the main #ow along the basin from the Laptev Sea slope toward Greenland. That a return #ow occurs along the Nansen-Gakkel Ridge was "rst postulated by Anderson et al. (1989) based upon observations from the Polarstern 1987 Arctic

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Fig. 5. CTD-pro"les obtained in 1991 by IB Oden in the interior of the Eurasian and Makarov Basins. (a) #}S curves, (b) #}S curves (blow-up of the deep data), (c) potential temperature pro"les (#), (d) salinity pro"les (S). (cf. Fig. 1).

Expedition. Its strength and also its existence have been doubted by Swift et al. (1997) and by Steele and Boyd (1998) because no distinct, warmer core of Atlantic Water is seen on the southern side of the ridge on hydrographic sections across the basin. This of course, may be due to the small distance between the two #ows, which does not allow for a stagnant intermediate water column to develop. Instead Swift et al. (1997) assume that the interior of the Nansen Basin is heated via double-di!usive convection by the spreading of layer structures in the warm core of the Fram Strait in#ow. The layering would then start to form as the Fram Strait branch enters the Arctic Ocean (Perkin and Lewis, 1984). Using the high value of 0.005 m s\ for the velocity in the layers estimated by Carmack et al. (1997) (see below) the horizontal heat transfer into the Nansen Basin and the spreading of the layers can be explained. It is still open which interpretation is the correct one. There is no doubt that a bifurcation of the boundary current occurs at the Lomonosov Ridge. One part crosses the ridge and enters the Canadian Basin, one

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part #ows along the ridge towards Greenland (Aagaard, 1989). These two #ows will mostly comprise the Barents Sea branch water, but the high temperatures, observed by Carmack et al. (1995) north of the East Siberian Sea and by Schauer et al. (1997) north of the Laptev Sea, show that a trace of the Fram Strait source is also present. The eastern Laptev Sea is also the area where most of the low-salinity shelf water crosses the shelf break and enters the interior of the Arctic Ocean basins. The lowering of the salinity of the mixed layer, resulting from this in#ow, reduces the depth of winter convection. The lower part of the previously ventilated mixed layer becomes isolated from the surface and forms a halocline between the Polar Mixed Layer and the Atlantic Layer (Rudels et al., 1996). The halocline continues both with the boundary current along the slope and along the Lomonosov Ridge into the Amundsen Basin.

4. The Canadian Basin The #ow of the boundary current into the Canadian Basin has been starkly revealed by the recently observed presence of anomalously warm Atlantic Water in the Arctic Ocean (Carmack et al., 1995). The propagation of this warm pulse along the continental slope as well as along the Lomonosov and Mendeleyev Ridges into the interior of the Arctic basins (Aagaard et al., 1996; Carmack et al., 1997; Swift et al., 1997) shows a circulation pattern consisting of a basic boundary current along the slope and detached gyres connected with the bathymetric features separating the di!erent basins. A sharp front was seen between the boundary current and the Makarov Basin water column and interleaving was observed across the front (Carmack et al., 1995). Extensive layer structures were also found on the AOS94 expedition (Aagaard et al., 1996; Carmack et al., 1997), and Carmack et al. (1997) followed the individual layers across entire basins. They suggested that the spreading of the layers was due to the energy release by double-di!usive convection, and that the lateral mixing caused by the intrusions could account for the heating observed in the interior of the Makarov Basin, away from the Mendeleyev Ridge. This implies much larger velocities in the layers, as well as a much larger e!ective horizontal di!usion coe$cient than those derived from the observations of meddies in the North Atlantic (Hebert et al., 1990). This large di!erence, 3;10 ms\ vs. 5 ms\, was pointed out by Carmack et al. (1997). The AOS94 observations in the Makarov Basin indicated a di!erence also in the deeper layers between the water masses close to the Mendeleyev Ridge and those further to the north (Swift et al., 1997). The layer between 500 and 1500 m was colder and fresher close to the Mendeleyev Ridge, and Swift et al. (1997) identi"ed these characteristics as those of the Barents Sea branch and concluded that the deep interior of the Makarov Basin was ventilated by a part of the boundary current that separates from the continental slope at the Mendeleyev Ridge. To account for the di!erent characteristics of the interior of the basin they proposed that the boundary current separates from the continental slope also on the American side and that it there carries warmer, more saline water from the Canada Basin into the Makarov Basin along the

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Lomonosov Ridge. The mixing of waters from these two sources would account for the properties observed in the interior of the Makarov Basin. However, the di!erence by the two inputs must be due to processes that transform the boundary current along the loop around the Canada Basin from the Mendeleyev Ridge to the Lomonosov Ridge. That the Canadian Basin deep water is warmer than that of the Eurasian Basin has been known for a long time, and this has normally been attributed to the presence of the Lomonosov Ridge (Worthington, 1953). Aagaard (1981) noticed that the deeper layers were also more saline, which indicated that water mass transformations must occur in the Canadian Basin. That the deep water of the Canadian Basin is not a remnant from a previous colder era but is ventilated, albeit slowly, also has been demonstrated by Schlosser et al. (1994a,b). Some deep water transformations occur, and these transformations must be such that the Atlantic Water becomes colder and fresher and the intermediate and deep water warmer and more saline. The creation of dense water by brine rejection on the shallow shelves, as was discussed by Nansen (1906) and Midttun (1985) in the context of the Barents Sea, is one possible mechanism and Aagaard et al. (1985) showed high-salinity shelf water (S'36) on the Chukchi Sea shelf, which, once it crosses the shelf break, should be dense enough to sink into the deeper layers of the Canadian Basin. Plumes sinking into the halocline have been described by Melling and Lewis (1982), and Aagaard et al. (1985) suggested that the transformation of the deep water could be accounted for by descending, shaving plumes. However, such plumes would not explain the higher temperatures found in the intermediate and deep water of the Canadian Basin. Rudels et al. (1994) applied the idea of an ensemble of entraining plumes, each sinking to its own matching density level, to explain the di!erences between the Eurasian Basin and the Canadian Basin. Utilising the fact that the boundary current splits at the Lomonosov Ridge with one part entering the Canadian Basin and one part #owing along the Lomonosov Ridge (Aagaard, 1989), they compared the water columns from two stations separated by the Lomonosov Ridge, assuming that the one in the Amundsen Basin could represent the in#owing boundary current and the one on the Makarov Basin side could represent the Canadian Basin water column. They then estimated the characteristics of the entraining plumes sinking down the slope, which could explain the observed changes. The entrainment rate was assumed constant, adding twice the initial volume for every 300 m of descent. No along-slope changes were considered. However, the recent "ndings from the AOS94 expedition } showing that the Makarov Basin characteristics are composed of in#ow from the Eurasian Basin, bringing part of a `younga boundary current, and of in#ow from the American side, bringing an `olda boundary current into the basin } indicate that this assumption may not be that bad. It was found that comparably little input from the shelves was needed to explain the changes in the water column, but the large entrainment necessary to explain the transformations required really high salinities (S'36}37) on the shelves to create plumes that would sink into the deepest layer. One result was that the Canadian Basin, due to slope convection, becomes warmer than the Eurasian Basin at about 1000 m (Fig. 5) while the sill depth for the boundary current is around 1700 m. The temperature di!erence between the intermediate depth

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layers of the two basins is then not due to the presence of the Lomonosov Ridge but due to the e!ect of slope convection (Meincke et al., 1997). If slope convection from the shelves were the only process transforming the deep water, we would expect a constant temperature and increasing salinity in the deeper layers, as is seen close to the bottom in the Eurasian Basin (Fig. 5). In the Canadian Basin there is a constant salinity and a decreasing temperature in the deep water above the isothermal and isohaline bottom water (Fig. 5). Jones et al. (1995) explained this feature by assuming a deep spill-over between the Amundsen and the Makarov basins through gaps in the Lomonosov Ridge away from the continental slope and the boundary current. The colder Amundsen Basin water is more compressible than the Makarov Basin water, and a small displacement downward would, because of the thermobaric e!ect, allow the water from the Amundsen Basin to sink through the weakly strati"ed Makarov Basin water column, perhaps reaching the bottom. Applying the same entrainment rate as for the slope plumes to the water sinking down the Lomonosov Ridge, Jones et al. (1995) found that about 80% of the deep water in the Canadian Basin is renewed by such spill-over. The salinity of the Paci"c in#ow is too low to directly a!ect the boundary current. Nevertheless, it restricts the production of dense water on the shelves and, consequently, inhibits the ventilation of the deeper layers. It also makes the Polar mixed layer less saline than in the Eurasian Basin. Water of the Polar Mixed Layer which enters the Canadian Basin across the Lomonosov Ridge then becomes covered by less saline Paci"c Water and shows up as a part of the halocline. Most of the Paci"c Water leaves the Arctic Ocean through the Canadian Arctic Archipelago (McDonald, 1996), but some enters the Eurasian Basin and leaves the Arctic Ocean through the Smith Sound and through Fram Strait (McDonald, 1996; McLaughlin et al., 1996; Jones et al., 1998). As the boundary current recrosses the Lomonosov Ridge north of Greenland (Newton and Sotirin, 1997), the densest water remains in the Canadian Basin. The part of the boundary current that has circulated around the Canada Basin and #ows directly into the Eurasian Basin (the true circumpolar part) can be distinguished from the return #ow via the Makarov Basin by slightly lower maximum temperature and by smoother pro"les (Fig. 6, station 43). Its silicate content is also higher (Rudels et al., 1994; Jones et al., 1995). At the Morris Jesup Plateau the boundary current meets the di!erent #ows recirculating in the Eurasian Basin. These are warmer and more saline in the Atlantic Water, and the salinity minimum, caused by the injection of the Barents Sea branch water, is seen at intermediate depth on all stations except 49, which is above the Nansen-Gakkel Ridge (Fig. 6). The interleaving, seen in the Atlantic Layer, is similar to that found further into the Eurasian Basin (Fig. 5), and Rudels et al. (1994) observing the layers close to Fram Strait assumed that they are remnants of the isopycnal interactions between the Barents Sea and the Fram Strait branches which are advected along the basin toward Fram Strait. This interpretation di!ers from the one proposed by Perkin and Lewis (1984) and adopted by Carmack et al. (1997), who assume that the layer structures are formed by interaction between the Fram Strait in#ow and the Arctic Ocean water column and spreading into the basin. The high-salinity Canadian Basin Deep Water, which crosses the Lomonosov

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Fig. 6. CTD-pro"les obtained in 1991 by IB Oden on a section north of the Fram Strait. (a) #}S curves, (b) #}S curves (blow-up of the deep data), (c) potential temperature pro"les (#), (d) salinity pro"les (S). (cf. Fig. 1).

Ridge, shows up as an intermediate salinity maximum at about 1800 m. Most of this water continues to Fram Strait and the Greenland Sea, but a part #ows into the Amundsen Basin, providing an intermediate salinity maximum, distinguishing the deep parts of the Amundsen and the Nansen Basin water columns (Jones et al., 1995).

5. Water mass transformations and water masses The thermohaline processes in the Arctic Ocean act as separators on the entering Atlantic Water. The freshwater added by river run-o! and net precipitation, and the freezing and melting cycle create a less dense, low salinity upper water mass comprising the Polar Mixed Layer and the halocline. Ice formation and brine rejection lead to an accumulation of saline water on the shallow shelves. Its subsequent sinking down

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the continental slope as entraining boundary plumes transforms the deep water column. The upper part becomes cooler and less saline, while the deeper layers become warmer and more saline. In the deepest layer the temperature becomes constant while the salinity increases toward the bottom. This happens because the plumes have the same initial temperature and entrain the same amount of intermediate water. If the plumes reach down to layers so dense that they are only ventilated by slope convection, these layers will have the same temperature as the plumes and no further temperature change will occur. The most saline plumes reach the deepest levels and the bottom layer becomes strati"ed in salinity. The departure from this pattern in the Canadian Basin was discussed in the last section. The Atlantic Water thus also becomes transformed into colder and denser, but normally less saline water masses. We assume that the plumes ventilating the deep Arctic Ocean are short lived and of small horizontal scale. Their initial thickness, as they cross the shelf break, is expected to be less than half of the depth of shelf. That would mean about 10}15 m for most of the Arctic shelves. We assume that the entrainment is caused by turbulence generated by bottom stress, which also acts to break the balance between the gravitational and the Coriolis accelerations and allows the plumes to cross isobaths. The entrainment increases the volume of the plumes and we assume that this, essentially, leads to an increase in their width but not in their thickness. If they remain thin the dominance of bottom stress will be maintained. The small scales of the plumes make them di$cult to observe and the only reasonably well documented case is the Storfjorden out#ow (Quadfasel et al., 1988; Schauer, 1995) (Fig. 7). In this case, the boundary plume, which could be traced to the high salinity (S'35.5) bottom water of Storfjorden, is seen as a warmer, more saline bottom layer on the continental slope. The plume is also seen as a cold, saline and dense layer in Fig. 3, as it crosses the shelf break south of S+rkapp. In recent years, the Storfjorden out#ow has been the object of intense modelling activity, and the models reproduce well the path of the plume and the change of its characteristics (Jungclaus et al., 1995; Backhaus et al., 1997). A striking observation of a thick, cold and saline bottom plume was made in Fram Strait west of Svalbard in 1988 from RV Valdivia (Fig. 8). The low temperature indicates little entrainment of Atlantic Water, while the salinity, which is low compared to that of the Storfjorden bottom water, implies a large entrainment of low-salinity water on the shelf before the plume passes through the Atlantic Water. The plume was also thicker than those found by Quadfasel et al. (1988) (Fig. 7), which suggests that its source also could have been an out#ow from the Bear Island Channel further to the south. Cold, dense water has been observed to form over the Central Bank and drain into the Hopen Deep, where bottom water characteristics much similar to those of the plume have been observed (Quadfasel et al., 1992). If such bottom water returns along the Svalbard Bank and sinks down the Bear Island Channel into the Norwegian Sea, as the water seen on Fig. 3, it would avoid the Atlantic Water in the Bear Island Channel and a large part of the Atlantic Water in the Norwegian Sea. The entrainment of warmer water would then be reduced. A thicker plume also would be less dominated by bottom stress, and the turbulent activity, which homogenises the plume and drives the entrainment, would be smaller.

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Fig. 7. CTD pro"les obtained in 1986 by RV Valdivia on the continental slope west of Svalbard. (a) #}S curves, (b) #}S curves (blow-up of the deep data), (c) potential temperature pro"les (#), (d) salinity pro"les (S). (cf. Fig. 1).

The characteristics then change more slowly and the plume would become strati"ed. There are indications of this in the pro"les shown in Fig. 8. A comparison between Figs. 4, 7 and 8 shows that the plumes on the Svalbard slope are di!erent from the St. Anna out#ow. The water from Storfjorden sinks down the slope, driven by gravity, and its characteristics are changed by diapycnal mixing (entrainment). Fig. 4 shows that the St. Anna out#ow is, in reality, a strong through#ow of a strati"ed water column. As it reaches the Arctic Ocean, the less dense fraction occupies the shallow part of the slope. The denser part slides down to deeper levels displacing the Fram Strait branch. The interaction between the in#ow and the Fram Strait branch is lateral and isopycnal, and it is more likely induced by instabilities of the #ow than by bottom friction. It is the di!erence between a buoyant plume and a jet. The out#ow is con"ned by rotation to the slope and is strong enough to depress the isopycnals at the slope below the levels they have in the basin interior (Schauer et al., 1997).

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Fig. 8. CTD-pro"les obtained in 1988 by RV Valdivia on the continental slope west of Svalbard. (a) #}S curves, (b) #}S curves (blow-up of the deep data), (c) potential temperature pro"les (#), (d) salinity pro"les (S). (cf. Fig. 1).

However, on the deeper slope stations (Fig. 4) an increase in salinity, relative to the same level in the interior, is seen close to the bottom. This thin, high-salinity layer may be a trace of dense boundary plumes sinking down the slope below the main in#ow of the Barents Sea branch. The source of such saline, dense water could either be the deepest layer of the St. Anna in#ow or the shelf areas in the eastern Kara Sea or around Severnaya Zemlya. Plumes from both these sources would avoid entrainment of warm Atlantic Water. The Eurasian Basin water column is thus dominated by the St. Anna in#ow at its intermediate levels, and the e!ects of slope convection become prominent only below 2000 m. This is in strong contrast to the Canadian Basin where the e!ects of slope convection are especially strong at intermediate depth. The thermohaline processes in the Arctic Ocean have transformed the entering Atlantic Water into several water masses as dense or denser than the over#ow water supplying the North Atlantic Deep Water: The modi"ed Atlantic Water of the

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Atlantic Layer of the Arctic Ocean, historically de"ned as water warmer than 03C, the Canadian Basin Deep Water, in the Eurasian Basin identi"ed as an intermediate salinity maximum, and the Eurasian Basin Deep and Bottom Water. Between the Atlantic Layer and the Canadian Basin Deep Water the potential temperature and salinity pro"les are characterised by decreasing temperature and increasing salinity with depth (Figs. 5 and 6) and the water column is stable in both properties. The water in this depth range is denoted upper Polar Deep Water. The slope of its #}S curve indicates that it is a product of shelf-slope convection and it is seen in its purest form in the Canadian Basin (Fig. 5). In the Eurasian Basin the in#ow of the Barents Sea branch creates a salinity minimum, and occasionally a temperature minimum, below the Atlantic Layer. At least in its deeper range, it will have a similar slope as the upper Polar Deep Water of the Canadian Basin. Although it is formed by a di!erent process, the water between the salinity minimum and the salinity maximum in the Eurasian Basin also will be considered as upper Polar Deep Water. As the recombined boundary current approaches Fram Strait the di!erence between the contributions from the two basins of the Arctic Ocean in this depth and density range will be small compared with the di!erences between the waters from the Arctic Ocean in the north and from the Nordic Seas to the south. Because of the meeting and mixing of several water masses of di!erent origin that take place in Fram Strait, it might be appropriate, at this stage, to give a more systematic classi"cation of the di!erent waters even if the second deep water formation area, the Greenland Sea, has not yet been discussed. On purpose, our classi"cation is based on only 3 CTD-data sets: 1st } CTD-pro"les obtained from IB Oden in the Arctic Ocean in August/September 1991 (Fig. 9, top panels); 2nd } CTD-pro"les obtained from RV Valdivia in the Greenland Sea and in the Norwegian Sea in May 1993 (Fig. 9, bottom panels); 3rd } two pairs of CTD sections across the Fram Strait obtained by RV Lance, RV Polarstern and RV Valdivia during MIZEX 1984 (Fig. 10). Furthermore, we concentrate on Fram Strait and de"ne in six layers, separated primarily by isopycnals, two water masses, one of southern and one of northern origin. The scheme, consistent with the observed pro"les, is given in Fig. 11 and in Table 1. The water masses have been separated according to (#, S)-ranges and to (# vs. S)-slopes. The lower limit of the Atlantic Layer is taken to be 03C, which in the Eurasian Basin closely corresponds to an intermediate salinity minimum due to the St. Anna Trough in#ow. The upper Polar Deep Water is found below the Atlantic layer down to the p "30.444 isopycnal, which separates it from the Canadian   Basin Deep Water below. Canadian Basin Deep Water and Eurasian Basin Deep Water are separated by the p "35.142 isopycnal, which approximately corres  ponds to the depth of the Lomonosov Ridge (Aagaard et al., 1985). Waters with higher densities do not cross the ridge and are de"ned as Canadian Basin Deep Water. In the Eurasian Basin the water below the isopycnal p "39.738, located close to the sill   depth of Fram Strait, is de"ned as Eurasian Basin Bottom Water. In some of the layers water masses from both regions are equally prominent, in other layers one source may dominate strongly. We do not expect that the abbreviations used as short hand in this classi"cation aimed at Fram Strait shall `catch ona by workers discussing the wider realm of the Arctic Mediterranean. However, we

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Fig. 9. #}S curves of CTD-pro"les to the north (top panels) and to the south (bottom panels) of the Fram Strait obtained by IB Oden in 1991 and by RV Valdivia in 1993. The right side panels give blow-ups of the deep water range.

believe that this classi"cation is worthy of some contemplation. We try to avoid acronyms in the text and shall here mainly use Modi"ed Atlantic Water and upper Polar Deep Water, to distinguish the intermediate water from the north from, the Atlantic Water and the Arctic Intermediate Water from the Nordic Seas. The deep water encountered in the Greenland Sea and Norwegian Sea will be called Greenland Sea Deep Water and Norwegian Sea Deep Water, respectively. The Modi"ed Atlantic Water is what Mauritzen (1996a,b) has denoted Arctic Atlantic Water.

6. Fram Strait In Fram Strait the boundary current encounters, and mixes with, the recirculating branches of the West Spitsbergen Current. On the sections shown in Fig. 2, which are

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Fig. 10. #}S curves of CTD-pro"les on a northern (&80.33N, top panels) and on a southern (&78.93N, bottom panels) across the Fram Strait obtained during MIZEX'84. The right side panels give blow-ups of the deep water range.

taken north of the sill, the saline and warmer Arctic Ocean deep waters dominate on the deepest part of the section. Only close to Svalbard does the rising of the isotherms towards the slope indicate the presence of deep water from the Nordic Seas. In the central and western part of the section the salinity has a maximum at the bottom, and to the west a second intermediate salinity maximum is seen. These maxima correspond to the Eurasian Basin and to the Canadian Basin deep waters, respectively. The Nordic Seas Deep Water has been followed to the north side of the Yermak Plateau (Jones et al., 1995), but is not seen in the temperature and salinity pro"les further to the east. The salinity and temperatures vary between the two years, 1987 showing colder and less saline Arctic Ocean deep water. Such variations may be due to the position of the front between the southern and northern water masses, which could be forced to move north and south in the strait in response to external forcing.

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Fig. 11. Schematic #}S diagrams showing the boundaries between the di!erent water masses in Fram Strait (cf. Table 1). Table 1 A water mass classi"cation for Fram Strait Polar surface water; p (27.70  Polar surface water, warm (PSWw); # *0  Atlantic water; p *27.70; p (30.444   

Polar surface water, cold (PSWc); # (0 

Atlantic water, proper (Awp) Atlantic water, modi"ed (AWm) (a) # *2 (a) p (27.97; # (2; S(S     + (b) p *27.97; 0)# (2; `"ngera (b) p *27.97; 0)# (2; not `"ngera     Intermediate water; p *27.97; p (30.444; # (0     Upper polar deep water (UPDW) `robusta

Arctic intermediate water (AIW) `"ngera or `di!usivea Upper deep water; p

*30.444; p (35.142    

Nordic Sea Deep Water, type a (NDWu); S(S  Lower deep water; p *35.142;   Nordic Sea Deep Water, type b (NDWl); S(S


Eurasian Basin Deep Water, type a (EDWu); S )S(S   ! p (39.738   Eurasian Basin Deep Water, type b (EDWl); S
)S(S  !

Canadian Basin Deep Water (CDW) S )S ! Canadian Basin Bottom Water (CBW); S )S !

Bottom water; p *39.738   Nordic Sea Bottom Water (NBW)

Eurasian Basin Bottom Water (EBW)

S(S


S
)S

S "34.887!0.039#  
S "34.918!0.0014#  S "34.9!0.06# !  S "34.676#0.232# + 

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To the east the temperature range between !0.5 and 13C occupies a very thin depth interval compared to the wide spreading of these isolines seen on the Greenland slope. To the west this temperature range comprises the Modi"ed Atlantic Water and the upper Polar Deep Water and shows the results of cooling the Atlantic Water in the Arctic Ocean as well as the water added by the Barents Sea branch. There is a salinity minimum present across the section at about 800 m. To the east it is Arctic Intermediate Water from the south that #ows in the West Spitsbergen Current. It is cold and perhaps slightly more saline than the minimum seen in the west. The western minimum could be a remnant of the salinity minimum present in the Eurasian Basin (Figs. 5 and 6), but most likely it results from the recirculation of the West Spitsbergen Current. The Atlantic Water of the West Spitsbergen Current dominates in the upper 500 m of the cross section and con"nes the Modi"ed Atlantic Water to a small area close to the Greenland slope. The recirculation creates a salinity minimum between the Atlantic Water and the upper Polar Deep Water, and this minimum is probably what is observed at the Greenland slope. We may note that not only the deep water characteristics are di!erent between the years. The Arctic Intermediate Water was more prominent and the Atlantic Water thicker in 1987, suggesting a stronger West Spitsbergen Current that year, which would agree with less Arctic Ocean deep waters. However, the salinity of the Atlantic Water was higher in 1985. The low-salinity, cold Polar Surface Water is found as a thick wedge to the west. The less saline surface water resulting from ice melting on the Atlantic Water of the West Spitsbergen Current only forms a thin layer in the central and eastern part of the section. The interactions between the boundary current at the Greenland slope and the waters from the south are mainly isopycnal. Below the Polar Surface Water the potential temperature and salinity pro"les show that the recirculating Atlantic Water intrudes into and almost completely absorbs the upper part of the boundary current, the Modi"ed Atlantic Water and the upper Polar Deep Water (Fig. 12). The denser part of the upper Polar Deep Water, still recognised by its #}S slope, now shows up as the lower part of a salinity minimum in the pro"les and #}S curves. Inversions present in the upper part of the boundary current are wiped out and replaced by inversions and interleaving structures created in Fram Strait between the recirculating water and the boundary current. Interleaving is also seen in the intermediate and deep waters of the boundary current close to the slope. These intrusions are cold and less saline and indicate the presence of colder waters from the second main deep water formation area in the Arctic Mediterranean } the Greenland Sea. That these waters are found so far to the west suggests that not all the deep waters from the Nordic Seas #ow north along the Svalbard slope. Some penetrate northward more directly from the Greenland Sea. Arctic Intermediate Water interacts with the upper Polar Deep Water and the Canadian Basin Deep Water, while Greenland Sea Deep Water (NDWu and NDWl) is seen close to the bottom. The mixing is primarily isopycnal, and the sharp gradients suggest the existence of a narrow frontal zone somewhat like that found east of the

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Fig. 12. CTD-pro"les obtained in 1984 by RV Lance in Fram Strait on the continental slope east of Greenland. (a) #}S curves, (b) #}S curves (blow-up of the deep data), (c) potential temperature pro"les (#), (d) salinity pro"les (S). (cf. Fig. 1).

St. Anna Trough (Fig. 4). The mixing weakens the property contrasts between the water masses of the two formation areas and creates waters with characteristics of the Nordic Sea Deep Waters (NDW) (Fig. 11).

7. The Greenland Sea The boundary current exits Fram Strait as the East Greenland Current, transporting the recirculating Atlantic Water as well as the Arctic Ocean water masses and the main part ('90%) of the ice exported from the Arctic Ocean. The less dense Polar Surface Water follows the Greenland continental slope toward the Iceland Sea and Denmark Strait. One part, however, is de#ected eastward at the Jan Mayen

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Fracture Zone, bringing cold polar water and ice in the Jan Mayen Current towards Jan Mayen and the eastern rim of the Greenland Sea (Helland-Hansen and Nansen, 1909). The recirculating Atlantic Water often cools rapidly (to (23C) as it progresses along the Greenland slope into the Greenland Sea (Koltermann, 1991; Strass et al., 1993; Meincke et al., 1997). The horizontal maps shown by Koltermann (1991) also indicate that part of the recirculating water could enter the Boreas Basin at the Greenland Fracture Zone. However, since Atlantic Water with temperatures above 23C has been observed in the Greenland Sea as well as in the Iceland Sea (Figs. 15 and 16) this could also be a seasonal e!ect. The observations reported by Koltermann (1991) were made in May, as were those by Strass et al. (1993), while the Meincke et al. (1997) observations were made in March. By contrast the Lance stations of Fig. 15 were taken in September and the Valdivia stations of Fig. 16 in October. HaK kkinen (1987,1990) suggested that the deep and bottom water in the Greenland Sea was formed by upwelling, cooling and subsequent convection of the Atlantic Water at the ice edge at the eastern boundary of the East Greenland Current. Although such deep convection may not occur, a cooling of the Atlantic Water has to be expected, if that water is brought close to the surface in winter. On the seaward side of the polar waters the recirculating Atlantic Water #ows with the East Greenland Current closest to the Greenland Sea gyre. Thus it may easily interact with fractions of Arctic Intermediate Water formed in the western part of the gyre. Strass et al. (1993) found that the #}S characteristics of the Atlantic Water gradually converged towards those commonly assigned to the Denmark Strait Over#ow Water, and they suggested that such mixing process could provide a source for the deep Denmark Strait over#ow in addition to the Arctic Intermediate Water of the Iceland Sea, which has been proposed as the source of the over#ow by (Swift et al., 1980; Swift and Aagaard, 1981). It appears as if the Modi"ed Atlantic Water of the Arctic Ocean only intermittently passes through Fram Strait as an identi"able water mass. Much of the time it becomes incorporated into the recirculation of the Atlantic Water in the strait (Fig. 12). The denser parts of the upper Polar Deep Water, the Canadian Basin Deep Water, and the Eurasian Basin Deep Water survive the mixing that occurs in the strait and are identi"ed on the Greenland Slope by the appropriate #}S slope and as salinity maxima. Of the three salinity maxima recognised by Aagaard et al. (1991), the two deepest would correspond to Canadian Basin Deep Water (type 1 in Aagaard et al., 1991), and the Eurasian Basin Deep Water (type 2 in Aagaard et al., 1991), respectively. The third, less dense maximum could be created by isopycnal mixing between the deeper part of the East Greenland Current and less saline, colder Arctic Intermediate Water, which leads to salinity minima in the out#owing Arctic Ocean water column (Fig. 12). The upper part of the East Greenland Current becomes partly de#ected towards the east at the Jan Mayen Fracture Zone, and no water deeper than 1600 m crosses the ridge into the Iceland Sea. The denser water circulates around the central Greenland Sea gyre, penetrating towards the centre and interacting with the

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Greenland Sea waters. The Arctic Ocean deep waters then become colder and less saline. Aagaard et al. (1985) and Swift and Koltermann (1988) suggested that this circulation and mixing would form the Norwegian Sea Deep Water, which cannot be created locally by deep convection in the Norwegian Sea. Its higher temperature and salinity, compared to the Greenland Sea Deep Water, imply mixing with another warmer, more saline water mass. The Arctic Ocean deep waters have the needed properties. Aagaard et al. (1985) and Swift and Koltermann (1988) assumed that the mixing would occur isopycnally. To create the Norwegian Sea Deep Water by isopycnal mixing with Greenland Sea Deep Water requires dense Arctic Ocean deep water (i.e. type 2 in Aagaard et al., 1991). The characteristics of this water mass are close to characteristics assigned to the Eurasian Basin Deep Water by Swift and Koltermann (1988). Koltermann (1991) followed the characteristics of the deep salinity maximum present in the Greenland Sea on stations from the Greenland slope eastward and from Jan Mayen northward into the central Greenland Sea gyre. He found that the maximum in the centre would line up with the salinity maximum found at the Greenland continental slope with only a weak diapycnal component. However, Aagaard et al. (1991), after examining the evolution of their type 1 salinity maximum, found that the characteristics were approaching those of the Greenland Sea Deep Water along a regression line crossing isopycnals. They therefore suggested that a diapycnal mixing process also must be operating. A look at the stations in Figs. 13 and 14 , which show the contrast between the rim and the central gyre of the Greenland Sea in 1993 (Fig. 13) and the central gyre in 1988 and 1993 (Fig. 14), indicates that internal diapycnal mixing is not dominating. The upper salinity maximum deriving from the Canadian Basin Deep Water does not supply the deep salinity maximum in the Greenland Sea but enters higher in the water column. In periods of weaker deep convection it will develop into a temperature maximum and occasionally into a salinity maximum around or slightly deeper than 1000 m, above the more permanent deep salinity maximum. When deep convection is active this temperature maximum will be cooled and incorporated into the convecting waters. The convection will also erode the deep salinity maximum, which arises from the penetration of Eurasian Basin Deep Water into the Greenland Sea (Fig. 13). The deep maximum thus becomes colder by isopycnal mixing, as it penetrates toward the centre of the gyre. If the convection removes the most saline part, the salinity maximum will become colder, less saline and denser. When the convection breaks through the salinity maximum and renews the bottom water, the density surface corresponding to that of the Eurasian Basin Deep Water will be displaced upwards, and the salinity maximum will penetrate from the rim into the centre of the gyre higher up in the water column. This is not to say that internally generated diapycnal mixing does not occur. When deep convection is weak, or absent, the salinity and temperature of the bottom water increase because of turbulent mixing, probably generated close to prominent topographic features. The salinity maximum is eroded from below and becomes resupplied from the rim of the basin. The colder, less saline bottom water gradually leaks out of

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Fig. 13. CTD pro"les from the rim and the central Greenland Sea taken in 1993 by RV Valdivia. (a) #}S curves, (b) #}S curves (blow-up of the deep data), (c) potential temperature pro"les (#), (d) salinity pro"les (S). (cf. Fig. 1).

the Greenland Basin and the salinity maximum is displaced downward (Meincke et al., 1997). This situation is expected to prevail if the convection in the central Greenland Sea is so shallow that it does not penetrate into the temperature maximum of the Canadian Basin Deep Water. The stations shown in Fig. 14 taken in 1988 and in 1993, respectively, indicate that this has been the case during the 5 years elapsed between the stations. In 1988 the temperature and salinity maxima almost coincide and little Canadian Basin Deep Water can be identi"ed. The winter station was one where convection reaching 1250 m was reported (Rudels et al., 1989), and the station taken in early summer shows an almost homogenous layer reaching deeper than 2000 m indicating further deepening in late winter. The water columns, although weakly strati"ed, were not as homogenous as in the `chimneysa observed recently in the

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Fig. 14. CTD-pro"les from the central Greenland Sea obtained by RV Valdivia in 1988 and in 1993. (a) #}S curves, (b) #}S curves (blow-up of the deep data), (c) potential temperature pro"les (#), (d) salinity pro"les (S). (cf. Fig. 1).

Greenland Sea (Backhaus, pers. comm.). The winter station showed several density compensating inversions in salinity and temperature, and the summer station was slightly strati"ed in salinity and temperature. Both stations were covered by a less saline surface layer. In winter this was cold and homogenous, in summer it was thinner and heated by short wave radiation. The existence of a freshwater lid made Rudels et al. (1989) assume that a recent event of penetrating, haline convection had been observed. A low-salinity layer present also in summer suggests ice melt and is consistent with an assumption of earlier events of haline convection. However, it is not possible to rule out that the freshwater is advected to the station from the sides and that the deep convection was thermal (Backhaus, 1995). The station from 1993 shows a shallower convection (!900 m) and the upper temperature and salinity maxima are not removed. The intrusion of warmer water at

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Fig. 15. CTD-pro"les obtained in 1988 by RV Lance on the continental slope east of Greenland. (a) #}S curves, (b) #}S curves (blow-up of the deep data), (c) potential temperature pro"les (#), (d) salinity pro"les (S). (cf. Fig. 1).

400 m is likely due to lateral interaction with parts of the Greenland Sea, where the strati"cation has been stronger and the convecting water has not attained as high a density, nor reached as deep. This situation, where only Arctic intermediate water is formed also in the central Greenland Sea, is in strong contrast with 1988. Fig. 15 shows isopycnal mixing occurring at the Greenland continental slope between the Arctic Intermediate Water formed close to the rim and the Canadian Basin Deep Water in the East Greenland Current. The Arctic Intermediate Water is as dense as and interacts with the Canadian Basin Deep Water. The mixing is isopycnal and the strong slope of the isopycnals can be deduced from the di!erence in depth between the levels, where the waters are interchanged, 500 m in 50 km. The conditions in the Greenland Sea are almost completely opposite to those of the Arctic Ocean. In the Arctic Ocean the shelves are shallow and ice free in summer. Ice

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formation, melting and brine rejection lead to a stable water column, and slope convection creates a vertical redistribution of heat driving the water column towards a state with constant temperature but salinity increasing with depth. In the Greenland Sea the shelf is covered by thick ice and rapidly moving water. The open ocean convection here leads to a water column homogenous in both temperature and salinity. In the absence of convection advection from the Arctic Ocean begins to stratify the Greenland Sea water column. This does occur not only by an in#ow of low-salinity surface water but also and more e!ectively by the invasion of Arctic Ocean deep waters into the central part of the gyre. When convection prevails, the Greenland Sea acts to partly remove the e!ects of the separation of the Atlantic Water into a high density and a low density mode which occurs in the Arctic Ocean.

8. The Iceland Sea Most of the East Greenland Current above 1600 m crosses the Jan Mayen Fracture Zone and enters the Iceland Sea. This part includes the Modi"ed Atlantic Water, waters in the ranges of the Canadian Basin Deep Water and the upper Polar Deep Water, Arctic Intermediate Water from the Greenland Sea and the Atlantic waters recirculating from Fram Strait, as well as the low salinity surface water. The deep western part of the Iceland Sea is then ventilated from the Arctic Ocean and the Greenland Sea to the north rather than from the Norwegian Sea to the east. This is supported by the higher salinity in the deepest layer and the more stable #}S slope of the intermediate layers, indicating Canadian Basin Deep Water and upper Polar Deep Water, both observed in the western Iceland Sea (Fig. 16). The upper Polar Deep Water #ows high on the continental slope and is not identi"ed in the central and eastern Iceland Sea. It is then likely to cross the sill and leave the Iceland Sea through Denmark Strait. Although waters with Canadian Basin Deep Water as well as upper Polar Deep Water characteristics have been observed at the sill in Denmark Strait (Buch et al., 1992,1996), the Canadian Basin Deep Water is too dense and located too deep to pass through Denmark Strait. It has to turn at the sill and #ow along the Iceland shelf slope eastward into the eastern Iceland Sea and eventually into the Norwegian Sea. High salinities have been reported in the deep water north of Iceland (Malmberg, 1983), and a mid-depth very weak salinity maximum is present in the Norwegian Sea (Clarke et al., 1990). This salinity maximum has already been connected with the Arctic Ocean deep out#ow by Aagaard et al. (1991). If it indeed is a trace of Canadian Basin Deep Water, some deeper part of the boundary current has been identi"ed before it exits through the Faeroe-Shetland channel in the eastern part of the Greenland Scotland Ridge after a complete circle around the pole. The Arctic Intermediate Water found in the central Iceland Sea is more saline than the Modi"ed Atlantic Water and warmer than the upper Polar Deep Water. These waters then cannot be the source of the Iceland Sea Arctic Intermediate Water even if they are partly de#ected toward Jan Mayen at the Jan Mayen Fracture Zone.

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Fig. 16. CTD-pro"les from the Iceland Sea obtained by RV Valdivia in 1991: (a) #}S curves, (b) #}S curves (blow-up of the deep data), (c) potential temperature pro"les (#), (d) salinity pro"les (S). (cf. Fig. 1).

Indications that such a de#ection occurs is seen on station 18 in Fig. 13. It is also warmer and more saline than the Intermediate Water found in the central Greenland Sea and it shows up as a temperature maximum (Fig. 16). However, closer to the rim of the Greenland Sea gyre and the recirculating Atlantic Water, the salinity and the temperature of the Arctic Intermediate Water formed in winter are expected to be higher and could be more similar to the Iceland Sea Arctic Intermediate Water characteristics (compare station 44 in Fig. 13). A possible local source area for the Arctic Intermediate Water in the Iceland Sea is south of Jan Mayen (Stefanson, 1962; Swift and Aagaard, 1981). Part of the Atlantic Water de#ected westward in the Norwegian Sea may be cooled in winter and incorporate some freshwater to account for the reduced salinity. The winter convection would form the lower Arctic Intermediate Water in the Swift and Aagaard (1981) terminology. The upper Arctic Intermediate Water (Swift and Aagaard, 1981) involves

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cooling and convection of waters in contact with the low salinity upper layers of the East Greenland Current and could occur within the Jan Mayen Current and at the eastern boundary of the East Greenland Current. The Arctic Intermediate Water formed south of Jan Mayen may return as an intermediate salinity minimum to the Norwegian Sea and perhaps also spill over to the western Iceland Sea. The hypothesis that the intermediate salinity minimum in the Norwegian Sea partly derives from the Iceland Sea was proposed early by Stefanson (1962), who also pointed out the area south of Jan Mayen as a likely formation area. There are then two sources of Arctic Intermediate Water making up the salinity minimum in the Norwegian Sea. The Greenland Sea forms a colder, denser type of Arctic Intermediate Water which has been described by Blindheim (1990). The salinity minimum in the southern Norwegian Sea (Fig. 16, station 4) is also warmer and more saline than the salinity minimum observed in the northern Norwegian Sea (Fig. 13, station 35).

9. Transports The strength of the boundary current so far has not been discussed. It is not our intention to make a complete review of the exchanges occurring through the di!erent passages in the Arctic Mediterranean, as has been done by Aagaard and Carmack (1989). Nor shall we try to formulate budgets by minimisation in a multi-box model, as was recently attempted by Mauritzen (1996a,b). Rather, we shall assume that the transports through the key passages, Fram Strait and the Barents Sea, are fairly well established and then use our knowledge of the water mass transformations in the Arctic Ocean to estimate the production and #uxes of di!erent water masses. As a constraint we apply estimates of the run-o! to the Arctic Ocean, which limits the amount of Atlantic Water that can be transformed into less dense water masses. From current measurements in the Bear Island Channel, Blindheim (1989) found a net in#ow of 1.9 Sv to the Barents Sea, which he assumed would enter the Kara Sea and then the Arctic Ocean. A similar transport was determined from year long current measurements between Franz Josef Land and Novaya Zemlya (Loeng et al., 1993). Roughly 2 Sv then enter the Arctic Ocean over the Barents Sea. About 0.8 Sv derives from the Norwegian Coastal Current (Rudels, 1987), the rest is Atlantic Water from the Norwegian Atlantic Current. From geostrophic calculations Bourke et al. (1988) obtained an in#ow of 0.6 Sv in the West Spitsbergen Current, and in the mass budget presented by Aagaard and Carmack (1989) an Atlantic in#ow of 1 Sv was adopted. We assume here that 1 Sv enters the Arctic Ocean through Fram Strait. The freshwater interacting with the Atlantic in#ow primarily derives from the Siberian rivers Ob, Yenisey and Lena. Seventy-"ve per cent of the 0.075 Sv supplied by these rivers is exported as sea ice (Aagaard and Carmack, 1989), which leaves 0.02 Sv to mix with the Atlantic in#ow to form Polar Surface Water. If the mean salinity of the Norwegian Sea in#ow is taken to be 35 and that the average salinity of the Atlantic derived Polar Surface Water is 33.7

B. Rudels et al. / Deep-Sea Research II 46 (1999) 1023}1062

1055

(Rudels et al., 1996), this implies that 0.6 Sv of the in#ow is transformed into less dense waters. The Atlantic Layer of the West Spitsbergen Current is about 500m thick. North of Svalbard an approximately 100 m deep mixed layer is formed from ice melt and subsequent freezing and convection. Its salinity is about 34.3 (Rudels et al., 1996). If the velocity of the core is independent of depth, 0.2 Sv of the Fram Strait Branch is then transformed north of Svalbard. This comparatively saline upper layer will be advected along with the boundary current. North of the Laptev Sea, the salinity of the Polar Mixed Layer is less because of the injection of low-salinity shelf water across the Laptev Sea shelf break, and the average salinity of the upper layers (Polar Mixed Layer plus halocline) becomes about 33.7. We assume that this injection supplies the remaining 0.4 Sv low-salinity upper water, and that it derives from water from the Barents Sea branch interacting with the run-o! from Ob, Yenisey and Lena in the Kara and Laptev Seas. Schauer et al. (1997) found that the Barents Sea branch at the slope penetrated at least down to 1200 m. The main part of the Barents Sea in#ow thus enters the boundary current below the Atlantic Layer. However, the strong interleaving observed at the slope (Fig. 4) indicates that at least some part of the Barents Sea branch mixes with the warmer Fram Strait branch. Of the remaining 1.6 Sv of the Barents Sea branch, we assume that 0.6 Sv mixes with the 0.8 Sv of the Fram Strait in#ow, forming the Modi"ed Atlantic Water, while 1 Sv sinks below the warm Atlantic Layer and becomes intermediate and deep water, mainly in the upper Polar Deep Water range. Most of the Modi"ed Atlantic Water recirculates within the Eurasian Basin, but a smaller part remains in the boundary current and enters the Canadian Basin. Slope convection in the Canadian Basin adds little net input from the shelves, but redistributes water from the Atlantic to the intermediate and deep layers (Jones et al., 1995; Rudels et al., 1994). In the entire Arctic Ocean we would expect that 0.2}0.3 Sv of the Modi"ed Atlantic Water is lost to deeper levels. Finally, the in#ow through the Bering Strait, 0.8 Sv (Coachman and Aagaard, 1988), is less than the out#ow through the Canadian Arctic Archipelago, 1 Sv (Rudels, 1987). We propose that 0.1 Sv from the Polar Surface Water and 0.1 Sv of Modi"ed Atlantic Water exit through the Canadian Arctic Archipelago. This implies that 0.5 Sv Polar Surface Water, 1 Sv Modi"ed Atlantic Water and 1.3 Sv intermediate and deep water leave the Arctic Ocean in the boundary current through Fram Strait. In the mid-1980s, the estimates of deep water formation in the Greenland Sea ranged between 0.35 and 1 Sv (Smethie et al., 1988; Heinze et al., 1990). Today, when little or no convective deep water renewal appears to take place (GSP Group, 1990; Schlosser et al., 1991; BoK nisch and Schlosser, 1995), these values are likely to be much too high. However, if little or no deep water is formed, more Arctic Intermediate Water has to be produced. Rudels (1986) estimated the formation of Arctic Intermediate Water in the Greenland Sea to be 1 Sv. If we add what is expected to form in the Boreas Basin, the combined production of Arctic Intermediate Water and Greenland Sea Deep Water would be 2 Sv. At present, perhaps the entire 2 Sv are Arctic Intermediate Water. Greenland Sea Deep Water can only

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B. Rudels et al. / Deep-Sea Research II 46 (1999) 1023}1062

leave the Greenland Sea to the east and the north, but the Arctic Intermediate Water can enter the boundary current to the west, as well as supply the intermediate salinity minimum in the northern Norwegian Sea to the east. We assume that 1 Sv of the Arctic Intermediate Water joins the East Greenland Current and 1 Sv comprising of Arctic Intermediate Water and Greenland Sea Deep Water enters the Norwegian Sea. The surface water, which is transformed in the Greenland Sea, could be the recirculating Atlantic Water. It is closest to the central gyre, and it has been observed to enter the Boreas Basin (Koltermann, 1991). The lower salinity of the Arctic Intermediate Water shows that it contains a fraction of low salinity surface water and/or some melted ice, which could enter the gyre via the Jan Mayen Current. Atlantic Water is not only recirculating within but also south of Fram Strait (Gladfelder, 1964; Quadfasel and Meincke, 1987), and the #ow of recirculating Atlantic Water could be as large as 3 Sv. A fraction of the Atlantic water in the boundary current may also become less dense from ice melt and by mixing with Polar Surface Water. If we assume that 0.5 Sv becomes transformed into low-salinity water in addition to the 2 Sv transformed into intermediate and deep waters, the boundary current, as it reaches the Jan Mayen Fracture Zone, will comprise 1 Sv upper water, 0.5 Sv recirculating Atlantic Water, 1 Sv Modi"ed Atlantic Water, 1 Sv Arctic Intermediate Water and 1.3 Sv Arctic deep waters. The Eurasian Basin Deep Water will remain in the Greenland Sea and eventually, together with 1 Sv of Greenland Sea Deep Water and Arctic Intermediate Water, supply intermediate and deep water to the Norwegian Sea. The part of the boundary current crossing the Jan Mayen Fracture Zone #ows southward to Denmark Strait. The Atlantic Water, the Modi"ed Atlantic Water and the upper Polar Deep Water cross the sill into the North Atlantic, supplying the North Atlantic Deep Water. As mentioned above, the Canadian Basin Deep Water is too dense and resides too deep to cross the sill, and it #ows eastward north of Iceland. If the estimates above, comprising a total of 4.8 Sv of dense water and 1 Sv of surface water, are compared with the presently adopted values of the dense over#ow across the Greenland}Scotland Ridge: Denmark Strait 2.9; Faeroe}IcelandRidge 1 Sv; and Faeroe-Shetland}Channel 1.7 Sv (Dickson and Brown, 1994), about 1 Sv of dense over#ow water is not accounted for. This volume could be supplied by Arctic Intermediate Water formed in the Iceland Sea out of the branch of the Norwegian Current which is detached westward from the boundary current at the latitude of Jan Mayen (Helland-Hansen and Nansen, 1909). The total transport of the boundary current entering the Arctic Mediterranean would then be close to 7 Sv. This "gure does not include the in#ow in the Irminger Current west of Iceland. The estimated out#ow of 1 Sv of low-salinity water is small (McCartney and Talley, 1984; Worthington, 1970,1976). Because of the constraint given by the ice export, it is di$cult to form more low salinity waters in the Arctic Ocean. One possibility is that the West Spitsbergen Current is still stronger and that the amount of recirculating Atlantic Water transformed into less dense waters in the Nordic seas is perhaps 1.5 Sv rather than 0.5 Sv. 2 Sv of low salinity water would then be exported by the East

B. Rudels et al. / Deep-Sea Research II 46 (1999) 1023}1062

1057

Greenland Current through the Denmark Strait, more than half of it formed in the Nordic seas. Of the waters contributing to North Atlantic Deep Water, the Arctic Ocean would supply 40%, and the Nordic Seas 60%. To compare these results with those obtained by Mauritzen (1996b) is di$cult, primarily because the initial assumptions about the in#ow to the Arctic Ocean di!er greatly. The discrepancy is largest for Fram Strait, where Mauritzen postulates an in#ow of 3 Sv and a recirculation of 1 Sv, while we have assumed 1 Sv into the Arctic Ocean and 3 Sv recirculating north of the Greenland Sea. Much of the di!erence in the results arises because of this. However, essentially three di!erences can be noted. (1) In the Arctic Ocean Mauritzen obtains almost twice as large a production of Polar Surface Water and of Modi"ed Atlantic Water (0.5 and 1.0 Sv compared to 1.5 and 1.5 Sv } our estimates are given "rst). (2) In Mauritzen's estimate most of the 1 Sv of recirculating Atlantic Water leaves directly via the East Greenland Current through the Denmark Strait, while here most of the recirculating Atlantic Water becomes transformed, mainly into Arctic Intermediate Water, but a fraction also becomes less dense and supplies the upper water. The amount of recirculating Atlantic Water that is estimated to leave through the Denmark Strait is thus close to that found by Mauritzen (0.5 and 0.8 Sv, respectively). The stronger recirculation and transformation lead in our case to a much larger production of intermediate and deep waters in the Greenland and the Iceland Seas, about 3 Sv as compared to 0.6 Sv. This includes the recirculation directly to the Iceland Sea. (3) The amount of water entering the Arctic Ocean over Barents Sea that becomes transformed into denser water is about the same (1.6 and 1.3 Sv). However, Mauritzen assumes that all this #ows through the Greenland Sea into the Norwegian Sea to supply the Norwegian Sea Deep Water and the out#ow through the Faeroe-Shetland Channel. In the present estimate we see that about 1 Sv leaves through the Denmark Strait and only that part which has become transformed into Canadian Basin or Eurasian Basin Deep Water enters the Norwegian Sea and exits through the FaeroeShetland Channel.

Acknowledgements We wish to thank Norbert Verch and Felix Yilderim for technical assistance, and an anonymous referee for preventing us from presenting a parochial view of the research in the Arctic Ocean. This work has partly been supported by the European Commission MAST II programme ESOP-I (MAS2-CT93-0057). Economic support for one of us (BR) has been received from the Deutsche Forschungsgemeinschaft (DFG-SFB-18, TP B3) and from the European Commission MAST III Programme VEINS, through contract MAS3-CT96-0070.

Appendix A The hydrographic stations' positions are given in Table 2.

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B. Rudels et al. / Deep-Sea Research II 46 (1999) 1023}1062

Table 2 A listing of CTD-pro"les used in this text Figure

Ship/cruise

Station(s)

Position

Date

01

All of those below

All of those below

All of those below

1984}1993

02

(a) Polarstern (b) Polarstern

541}569 196}215

Fram Strait, 803N Fram Strait, 803N

21}25.07.1985 13}17.06.1987

03

Lance/86 Valdivia/048

245-272 002 003 004

Barents Sea, &203E 70.4983N; 017.9983E 71.0023N; 017.9863E 71.4993N; 017.9903E

14./15.08.1986 02.08.1986 03.08.1986 03.08.1986

04

Polarstern

025 027 029 031

81.0983N; 81.2223N; 80.8353N; 80.7273N;

105.3573E 106.5773E 104.1683E 103.3023E

07.08.1995 08.08.1995 10.08.1995 11.08.1995

05

Oden

010 013 016 026

85.7303N; 86.7373N; 87.6033N; 88.0153N;

048.0523E 056.9753E 069.7453E 163.6033E

25.08.1991 27.08.1991 29.08.1991 04.09.1991

06

Oden

043 046 047 048 049

85.0533N; 84.5833N; 84.2203N; 83.9953N; 83.5883N;

014.8553W 006.1933W 002.4423W 000.2223W 004.3383E

19.09.1991 21.09.1991 22.09.1991 23.09.1991 24.09.1991

07

Valdivia/048

043 068 070 072

78.8323N; 77.1293N; 77.2533N; 77.3503N;

008.0003E 007.5713E 008.9803E 010.3223E

11.08.1986 14.08.1986 15.08.1986 15.08.1986

08

Valdivia/071

009 010 011

77.7533N; 009.9783E 77.7523N; 008.7433E 77.7523N; 007.9003E

05.06.1988 05.06.1988 05.06.1988

09

Oden Valdivia/136

005}061 002}076

North of FramStrait South of Fram Strait

Aug./Sept. 1991 May/June 1993

10

Lance Valdivia Polarstern

Fram Strait sections at &80.33N and &78.93N

MIZEX &84 MIZEX &84 MIZEX &84

12

Lance

78.8503N; 004.1503W 78.8533N; 003.4553W 78.9103N; 002.6473W

26.08.1984 26.08.1984 26.08.1984

251 252 253

B. Rudels et al. / Deep-Sea Research II 46 (1999) 1023}1062

1059

Table 2. Continued. Figure

Ship/cruise

Station(s)

Position

Date

13

Valdivia/136

008 012 018 035 044

73.9993N; 74.5013N; 71.5003N; 75.0023N; 76.2493N;

004.3273W 005.9943W 009.0003W 008.1403E 006.2223W

21.05.1993 22.05.1993 24.05.1993 29.05.1993 31.05.1993

14

Valdivia/067 Valdivia/071 Valdivia/136

065 033 013

74.7483N; 002.0183W 75.0003N; 002.0183W 74.0023N; 006.4993W

14.02.1988 11.06.1988 23.05.1993

15

Lance

110 111

76.0173N; 003.9333W 75.6173N; 008.8503W

27.09.1988 28.09.1988

16

Valdivia/104

004 008 018

67.4603N; 007.0123W 68.7163N; 010.4983W 69.1453N; 020.0673W

19.10.1990 20.10.1990 22.10.1990

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