Mesoscale thermohaline variability in the Eastern Gotland Basin following the 1993 major Baltic inflow

Transcription

1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 102, NO. C9, PAGES 20,917-20,926, SEPTEMBER 15, 1997 Mesoscale thermohaline variability in the Eastern Gotland Basin following the 1993 major Baltic inflow V. M. Zhurbas Shirshov Institute of Oceanology, Moscow, Russia V. T. Paka Atlantic Branch of Shirshov Institute of Oceanology, Kaliningrad, Russia Abstract. High horizontal resolution data from a "tow-yo" conductivity-temperature-depth profiling system are analyzed to examine the mesoscale and fine scale response of the Gotland Basin to the 1993 major Baltic inflow. The permanent halocline in the basin is characterized by intensive thermohaline intrusions. A well-defined front of the intrusive region was found to be propagating north from the Stolpe Channel to the Gotland Deep with a speed of 2 cm s 4 or more. A substantial horizontal intermittency of intrusion intensity, related to mesoscaleddies, was observed behind the front. A strongly nonlinear cyclonic eddy, undetectable on the sea surface, was found in the halocline. The cyclone contributed to the transport of salty water arer the inflow. The combined effect of the mesoscale eddies and thermohaline intrusions is shown to be one of the mechanisms responsible for the deep water ventilation by major Baltic inflows. 1. Introduction The water exchange between the North Sea and the Baltic Sea is known to be greatly restricted by a narrow and shallow transition area consisting of the Cattegat and the Belt Sea. The circulation and horizontal mixing of the deep water in the Baltic Sea are hampered by the bottom topography. The Baltic Proper consists of a sequence of basins (the Arkona, Bornholm, and Gotland Basins) separated and linked by sills and channels. The depths of both the sills and the basins increase from the transition area to the Gotland Basin, where the depth is as great as 250 m. Vertical exchange in the Baltic Sea is prevented by the strong density stratification due to the permanent halocline. These factors cause periods of stagnation in the Baltic water that are characterized by increasing phosphate and nitrate concentrations and decreasing salinity and oxygen concentrations. The stagnation sometimes culminates in the formation of considerable hydrogen sulfide concentrations in the deep basins [Matth us, 1993]. Interruption of the stagnation and ventilation of the deep water takes place only with strong inflows of highly saline and oxygenated water from the North Sea into the Baltic Sea. Such events, called the major Baltic inflows, are typical, but relatively rare phenomena. During a "normal" inflow, the volume of inflowing saline water is insufficient to displace the old bottom water in the deep basins of the central Baltic, since Copyright 1997 by the American Geophysical Union. Paper number 97JC /97/97JC ,917 the inflowing water mixes with ambient waters and becomes lighter than the old bottom water. The meteorological precondition for a major Baltic inflow is an anomalous atmospheric circulation with southerly winds over the NE Atlantic. This atmospheric circulation causes enhanced transfer of highly saline water from the open ocean to the North European shelf [Dickson, 1973]. The major inflows occur between the end of August and the end of April with a mean frequency of about one event a year [Matthtius and Franck, 1992; Matth us, 1993]. These inflow events are quite irregular; the longest recorded period of no inflow events lasted nearly 10 years ( ). A recent, major inflow took place in January 1993 and attracted the close attention of oceanographers from many Baltic countries [DaMin et al., 1993; Grelowski and Wojew6dzki, 1993; Hc}kansson et al., 1993; Matthtius, 1993; Matthtius et al., 1993; Ozmidov, 1993; Paka, 1996]. That event interrupted 16 years of stagnation, since the few major inflows between 1976 and 1983 were shown to be inefficient in displacing the bottom water in the deep basins. Compared with all identified events in the past, the 1993 major inflow was classified as a moderate event, with limited effects in the deep central Baltic basins [Matthtius, The meteorological forcing for the 1993 major inflow was due to very strong weather conditions in the transition area. During the 3 weeks between January 6 and 25, three hurricanes passed through the region (on January 14, 22, and 24), and extreme squalls of 45, 39, and 34 m s 4 from 250 ø to 260 ø were recorded at the meteorological station Arkona (Riigen Island) [Matth us et al., 1993]. The

2 20,918 ZHURBAS AND PAKA: TI-IE O E VARIABILITY OF THE BALTIC SEA inflow of highly saline water through the Sound across the Drogden Sill (7 m depth) and through the Great Belt across the Darss Sill (18 m depth) into the Arkona Basin started on January 6 and January 13, respectively, and finished on January Matth ius et al. [1993] report the further propagation of highly saline water in the Baltic Proper as follows: The inflowing water accumulated in the Arkona Basin for a week. During this time and afterward, the 15 practical salinity units (psu) isohaline first rose from 38 m depth to 10 m and then lowered until mid-february to a depth of m. This indicates a flushing time of about 2-3 weeks for the Arkona Basin. Part of the salt water discharged into the Arkona Basin mixed with the surface water and then flowed back into the transition area. Another part was mixed with shallow water and remained in the basin. The main body of highly saline water which entered into the Baltic flowed through the Bornholm Channel (45-m depth at the sill) into the Bornholm Basin and replaced the old bottom water. As a result, the salinity in the bottom layer of the Bornholm Basin increased from about 15 to 20 psu between January and March The old, deep Bornholm water of lower salinity and some part of the new water were lifted above the sill depth of 60 m and overflowed to the Stolpe Channel. Water masses with TS properties of the Bornholm intermediate water were observed in the Stolpe Channel from March to May. The maximum salinity of this water changed within a week, indicating that the outflow from the Bornholm Basin was rather intermittent. A turnover of highly saline water from the Bornholm Basin through the Stolpe Channel into the Eastern Gotland Basin started in the beginning of March. In the beginning of April, the first indications of the bottom water renewal were observed in the Gotland Deep at the station BY15 (57ø20'N, 20ø00'E). At this station, the salinity below 225 m increased from 11.0 psu on March 15 to 11.7 psu on June 4 [Dahlin et al., 1993]. The above description of saltwater propagation in the Baltic Proper was achieved using conductivity- temperature-depth (CTD) measurements with a station spacing of about km on transects through the deepest connection between the Darss Sill and the Gotland Deep. Such sparse spacing of stations cannot resolve the mesoscale structure of the process. However, it is in the mesoscale range that we expect to discover some specific mechanisms of saltwater propagation and mixing which will help us develop a better understanding of deep water ventilation in the Baltic Sea. CTD measurement suitable for studying the mesoscale variability were carried out in the Stolpe Channel and the Eastern Gotland Basin in March-April 1993 during the twenty-ninth cruise of the R/V Professor Shtockman [Ozmidov, 1993]. In this paper, we will use these data to examine the mesoscale and fine scale response of the 2. Data The CTD measurements were carried out in the Stolpe channel and the Eastern Gotland Basin from March 30 to E Figure 1. Map of the experiment, March-April The box (dashed line) is the area of CTD survey of the first stage of the experiment, March 30 to April 2. Solid circles coupled into bold lines of variable thickness are CTD stations of the second stage of the experiment, April A dotted circle is the location of the detected cyclone. April 18, This time period is just after the first indications of bottom water renewal were observed in the Gotland Deep. All observations were taken with a Mark III NBIS CTD profiler. To achieve high horizontal resolution, the CTD profiling was made with a winchdriven "tow-yo" vehicle equipped with a heavy chain to prevent the vehicle touching the ground [Paka, 1996]. Using this method, nearly 2000 CTD profiles with spacing of about m in the entire layer from the sea surface to the seabed (except for the near-surface and near-bottom layers of 1-2 m thick) have been obtained. A map of the CTD measurements in the Eastern Gotland Basin is shown in Figure 1. There were two stages of the experiment. During the first stage, from March 30 to April 2, a detailed tow-yo CTD survey, consisting of six legs, was carried out in the eastern part of the Gotland Deep. In Figure 1, this survey area is bounded by a dashed line. During the second stage, from April 16 to 18, two tow-yo CTD transects, A-B and C-D, were performed. Each tow-yo station (or each CTD profile) of the second stage is marked by a solid circle of different size depending on the value of intrusion intensity at this station. (Our definition and a derailed description of the intrusion intensity are given in the next section.) Solid circles, coupled into bold lines of variable thickness due to small distances between successive stations, show the transects in Figure 1. In addition to the transects, a few largest and deepest basin of the Baltic Proper to the recent separate CTD stations were made during the second stage, major inflow. from the station BY15 toward the south nearly along the deepest connection. These stations are also shown in Figure 1. In this paper, the data of the longest transect, A-B, will be analyzed in detail. The A-B transect was carried out on April The transect was 157 km long and consisted

3 ZHURBAS AND PAKA: THERMOHALINE VARIABILITY OF THE BALTIC SEA 20,919 of 411 stations, giving a mean distance between successive There are no inversions on the salinity profile, and CTD profiles of 383 m. The A-B transect started in the NE periphery of the Gotland Deep (the point B at 57ø29'N, 20ø30'E), went to the SSW to (56ø41'N, 19ø39'E), and than deviated to the WSW, ending at point A (56ø29'N, 18ø52'E). In the following figures, distance is given from point A along the A-B transect. intrusions are almost invisible. Inversions exist only on temperature profiles because density stratification in the Baltic halocline is almost completely determined by salinity. Here the mean density ratio, Rp-- rz/[3 z, satisfies the unequality lrpl<0.04, where ct=-(1/po)0p/0t and 13=(1/po)Op/OS are temperature expansion and salinity contraction coefficients, T and & are mean vertical gradients of temperature and salinity, and p and po are 3. Thermohaline Intrusions: The Spatial water density and reference water density, respectively. Distribution and Dependence on Mesoscale To evaluate qualitatively the degree of development of Dynamic and Kinematic Parameters intrusions on temperature profiles, a simple measure of Thermohaline fine structure in the A-B and C-D intrusion intensity, introduced by Zhurbas et al. [1988], is transects was characterized by abundant intrusions. To used. The intrusion intensity, J, for a layer between chosen illustrate this point, vertical profiles of temperature, isopycnals is defined as the total inversion temperature salinity, and potential density at a distance of km difference in the layer divided by its thickness: from point A (on the A-B transect) are shown in Figure J=Z,= nati/h, where ATi is the temperature difference in 2a. Intrusions are clearly visible on the temperature the ith inversion sublayer and n is the number of inversion profile, with temperature inversions up to øC. sublayers in the whole layer of thickness h. For analysis purposes, two layers have been chosen. The upper layer is between isopycnals of kg m -3, and the lower layer S, psu is between isopycnals of kg m -3. This choice of layers was determined by the following. First, both layers are characterized by well-defined intrusions in both A-B o,T øc 70 and C-D transects. Second, the mean vertical gradient of temperature in each of the layers does not change sign; in the upper layer the mean temperature increases with the 90- depth, while in the lower layer it decreases with the depth. Third, the sea depth is sufficiento discriminate the upper layer in the whole A-B transect. On the contrary, at the southern end of the A-B transect, the potential density near the sea bed is less than 8.6 kg m '3, and the lower layer is impossible to locate. To illustrate the spatial variability of intrusion intensity, CTD stations from the second stage of the experiment have been marked by solid circles of different size depending on the values of d in the upper layer. In Figure 1, the small circles are for d < 0.01øC m 4, the medium P - dbar circles are for 0.01 øc m - < d < 0.02øC m 4, and the large circles are for J > 0.02øC m -. For reference, in Figure 2a, CVo, kg m -3 d = 0.018øC m - for the upper layer, and d = 0.022øC m for the lower layer. To examine the origin of the intrusions, one can use a TS analysis. In Figure 2b, the TS diagram for the temperature and salinity profiles shown in Figure 2a (solid line), as well as TS diagrams from the central part of the Gotland Deep (station BY15, top dotted line) and from the Stolpe Channel (55ø30.6'N, 17ø52.8'E, bottom dotted line), are presented. Isopycnals in Figure 2b are omitted because they are almost indistinguishable from vertical lines. The TS diagram from Figure 2a, with the well- T,øC - defined intrusions, passes between the TS diagrams from station BY15 and the Stolpe Channel. This result confirms 2.0 S, psu a suggestion that the origin of intrusions after the recent Figure 2. (a) An example of vertical profiles of major Baltic inflow is related to the penetration of the temperature, salinity, and potential density anomaly with Bornholm water into the Eastern Gotland Basin. thermohaline intrusions (transect A-B, km from point A). (b) TS diagrams for the profiles shown in Figure The same origin of intrusions is confirmed by the 2a (solid line), from the Gotland Deep (top dotted line), spatial and temporal variability of intrusion intensity. A the Stolpe Channel (bottom dotted line), and the cyclonic derailed tow-yo CTD survey, carried out in the Gotland eddy (dashed line). Deep during the first stage of experiment, did not display

4 20,920 Z AS AND PAKA: THERMO INE VARIABILITY OF THE BALTIC SEA T,øC dbar Figure 3. Sequence of temperature profiles from a fragment of the A-B transect between 57ø14.8'N and 57ø19.0'N including the from of the intrusion domain. Numbers under the profiles are distances from the point A. The temperature scale is correct for the leftmost profile; successive profiles have been offset by 0.3 øc. any significant intrusions. A fortnight later, during the parameter is proportional to the vertical shear in a second stage of the experiment, the area filled with geostrophic flow, (%) N2 /f= N A, and is related to intensive intrusions was found in the southeast part of the the geostrophic Richardson number, Rig = N 2 / (%) 2, as A Gotland Deep, being limited north and west by deep = Rig 'm. Using CTD measurements in the Gulf Stream, J waters of no intrusions (see Figure 1). It appears that this was found to increase with (Th)p and decrease with A. The intrusion domain propagated northward along the eastern interpretation was that double-diffusive intrusions are slope of the Gotland Deep, advancing more than 12 suppressed (or destroyed) by shear instabilities in a nautical miles ( km) in a fortnight, or with a speed stratified flow. Using CTD data in the Subarctic frontal of at least 2 cm s 4. zone of the western Pacific, Kuzmina et al. [ 1994] found a The propagation of the intrusion domain is front-like in regime where J increased with A. This regime was character. To illustrate this, a sequence of temperature explained as the result of the generation of intrusions by profiles from the A-B transect between ' N and baroclinic instability [Kuzmina and Rodionov, 1992] ' N is shown in Figure 3. Despite a 3-km-wide gap In this study, we examine a new parameter, the in CTD profiling, a sharp front of intrusive layering is planetary component of potential vorticity V = f p /p0, clearly visible between and km. When where p is the mean vertical gradient of density. If crossing the front (toward point A), intrusions do not intrusions move as an ideal fluid, the potential vorticity, appear in the whole layer simultaneously but appear first defined as (c% + J) p / p0, where c% is the vertical in some thinner sublayers. These sublayers enlarge with component of relative vorticity, will remain constant for advancing distance into the intrusion domain until nearly material particles. If we suppose that I,%/fl << 1, then V the entire layer shows intrusions. The horizontal will be unchanged. These assumptions, if satisfied, suggest coherence of intrusions, from Figure 3, is of the order of 1 that intrusions move preferably in the direction of the km. minimum change in V and that sites of extreme values of It is also evident from Figure 1 that the intrusion V are unfavorable for intrusive layering. Since f is near intensity behind the front is horizontally intermittent with constant for mesoscales, the conservation of V implies the a typical scale of the order of 10 km. We shall now conservation of p. examine whether this intermittency is associated with the Figure 4 shows the intensity of intrusions, J, mean spatial variability of some mesoscale hydrological feature. values of temperature, Tm, pressure, Pineart, and potential The problem of revealing relationships between the vorticity, V, for upper and lower layers versus the distance intensity of intrusions and local mesoscale dynamic and along the A-B transect. With distance, the values of J kinematic parameters in oceanic frontal zones exhibit a well-defined intermittency, discussed above. We (parameterization of interleaving) was first stated by observe a front of intrusive layering at a distance of 130- Zhurbas et al. [1988]. In that study, two parameters were 133 km, and intrusions are practically absent north of the investigated: termoclinicity, (Th)p(the mean gradient of front in both layers. (This is the same from shown in temperature on isopycnals) and baroclinicity, A = 70 N /f, Figure 3.) Comparing J in the upper and lower layers, one where )t o is the mean slope of isopycnals to isobars (-- can visually identify several horizontally coherent (dpmo /dr) (g p0) 4, where Pmo is the mean pressure and structures of high intrusion intensity in both layers. x is distance) and N and f are buoyancy frequency and However, these structures are, for the most part, Coriolis parameter, respectively. The baroclinicity incoherent vertically, and the correlation coefficient for

5 ZHURBAS AND PAKA: TFIERMOHALINE VARIABILITY OF THE BALTIC SEA 20,921 Pineart, dbar J 100, øc m o 4o loo distance, km Figure 4. Intrusion intensity, d, mean values of temperature, Tmon, pressure, Pmoan, and potential vorticity, V, in the upper (solid lines) and lower (dotted lines) isopycnal layers versus distance for the A-B transect. rows of d for the upper and lower layers is Figure 4 well-defined relation between the intrusion intensity and clearly reveals only two relations between the intensity of the mesoscale parameters. Some trends are displayed, intrusions and selected mesoscale parameters. First, d is however, on the (Tm, d) and (V, 0 r) diagrams. On these low for Tm over some limit, namely for Tmo > 5.1 øc in plots, high values of d are concentrated moderate values the upper layer and for Tmo > 5.25øC in the lower layer. of Tmo and V, while the vicinities of extreme values of Second, at the distance of about 50 km, there is a sharp Tmo and V are characterized by low values of d in both absolute maximum of V in both layers which corresponds layers. The occurrence of high values of d with middle to the absolute minimum of Tmo and low values old. values of Tmo is consistent with the general notion of To analyze the relations between the intrusion intensity intrusion layering: namely, the extreme values of Tmo and mesoscale parameters in more detail, diagrams of correspond to the transformed Bornholm water and to the (Tmoan, 0r), (lyp], 0r), and (K J) have been plotted for both Gotland water, while moderate values of Tmo are due to layers (Figure 5). When estimating the slope of isopycnals the intrusive stirring of these waters. to isobars, ¾0, we found it necessary to eliminate some The relation between d and V in Figure 5 appears to be noise in Pmo due to errors of hydrostatic pressure an interesting new finding. To explain this dependence, measurements. Consequently, the rows of Pmoan(X) Were we first examine the spatial structure of the potential low-pass filtered with a smoothing scale of about 2.5 km, density anomaly, c o, in the A-B transect (Figure 6). The and dpmo / dx was approximated by finite differences over most outstanding features of Figure 6 are the convergence the same scale. None of the diagrams in Figure 5 shows a of isopycnals inside the halocline at the distance of 48-55

6 20,922 ZHURBAS AND PAKA: THE O E VARIABILITY OF THE BALTIC SEA d 100, øc m; ,.:. e- '..,.e e e -,..,.,... e-.,l :.... :. : x %:...'. ' ' ' I. ' *...,..,!. I..,... ;.'... >. ' I ' I ',*'5"; " J 100, øc m '1 " :':,i';.'.. ".. '-....., "- ',.:. line".'... ee a * Tmean, øc I V 109, (m s) '1 Figure 5. Diagrams of (Tmon, d), (1701,, and (V, d) for the upper and lower layers of the A-B transect (top and bottom panels, respectively). km and the divergences of isopycnals at distances of 116- The observed V dependence on d can be explained 129 km and km in the upper layer and qualitatively, if we assume that the inflowing water is km in the lower layer. The convergence corresponds to the mainly transferred by mesoscale eddies: cyclones and absolute maximum of V in both layers at the distance of anticyclones. In the cores of the eddies, which are filled about 50 km. (compare Figure 4). This feature can be with the leastransformed inflowing water, V is extremal, interpreted as a mesoscale cyclonic eddy covering the preventing intrusions from developing. The intensity of whole halocline layer. The divergences of isopycnals, intrusions is enhanced in areas km wide, adjoined corresponding to minima of V in layer, can be interpreted laterally to the eddies, producing exchange between the as anticyclonic eddies covering only part of the halocline. eddies and their surroundings. In Figure 4, a high level of The core of each eddy is characterized by a low intrusion J from km to km, with a miramum in J at 50 intensity, no matter whether it is the cyclone or one of the km, is the region of influence of the cyclonic eddy. The anticyclones. sequence of maximand minima of d between 90 and 135 Figure 6. Potential density versus pressure an distance for the A-B transect.

7 ZHURBAS AND PAKA: THE O E VARIABILITY OF THE BALTIC SEA 20,923 km is the region of influence of several anticyclonic eddies. The above pattern of spatial variability of intrusion intensity in the cyclone or in one of the anticyclones, observed here in the Baltic halocline, is common for oceanic mesoscale eddies and has been observed in Mediterranean salt lenses in the North Atlantic [Armi at al., 1989] and rings in the Subarctic frontal zone of the western Pacific [Zhurbas and Sagdiev, 1992]. 4. Underwater Cyclonic Eddy The eddies described in the last section can contribute to the saltwater transfer and deep water ventilation in the Baltic Sea. In this section, we consider the dynamic properties of the cyclone in more detail. We define the empirical geostrophic flow function as [. 1 1 Fx(p'x)- p p(p,x - p(p, oo) ]dp (1) where p(p,oo) is the background density versus pressure away from the cyclone, p(p, x) is density in the cyclone, and p0 is some reference pressure level with no geostrophicurrent. We took p0 to be equal to 140 dbar, the pressure just above the seabed at the cyclone location. The across-transect component of geostrophic velocity, ug, is written as the x derivative of the flow function: ug-- org/o. In Figure 7, we show Fg for a fragment of the A-B transect, which includes the central part of the cyclone and its northern periphery, where the pressure at the seabed is no less than 140 dbar. This function is a minimum at p 107 dbar, indicating the depth of the core of the cyclone, and increases to near zero over the halocline and away from the cyclone. From this figure, we conclude that the cyclone is located in the halocline and is dynamically undetectable on the sea surface. To estimate the size and velocity of rotation of the cyclone, the empirical geostrophic flow function at the level of p = 107 dbar has been approximated by an analytical function, namely the Gauss curve, (x- x o)2 Fg(x)-- Fg(p,x)[p=lo7 dbar = A exp- 2R 2 where A, x0, and R are parameters to be determined [Zhurbas et al., 1992]. Using a least squares regression, we obtained the following values: A =-1116 m 2 s 4, x0=51.15 km, and R=2.94 km. Since only a single transect across the cyclone is available, the distance between the transect and the center of the cyclone is unknown. We can estimate the errors introduced by this uncertainty by characterizing the cyclone as an axisymmetric flow field described by the two-dimensional Gauss curve, (2) (x- x6) 2 + (y-yo) 2 F x (x,y) = A'exp - (2') 2R,2 where x0', y0, R', and A' are "real" parameters of the flow function. If the transect is made along the x axis at a distance of y -yo from the center of the eddy, (2) transforms to (2') with the following relationships between "real" and "measured" parameters: x0 = Xo', R=R', and A=A'exp[-(y -yo) 2 / 2R '2] < A'. Thus, having a single transect, we would estimate the correct value of the size of the eddy, R, and underestimate the intensity of the eddy, A. Alternately, if the flow function is equal to a nonzero constant inside the eddy and zero outside it, a single transect will give the right value of the intensity of the eddy and underestimate its size. We believe the first case corresponds more closely to reality. From (2), geostrophic estimates of the azimuthal velocity and the frequency of rotation of the cyclone are given by and ux(r)ofx Or - R2 exp- 2R r2 2 (3) distance, km ] t, I ø I I [ T 58,8 C U g = - (4) O) g -=-- - r r=o R 2 loo 12o P, dbar -150 Figure 7. Geostrophic flow function for a fragment of the A-B transect with the cyclonic eddy. where r=x-xo is the distance from the center of the cyclone. Using the above values of A and R, (4) yields to = 1.29x10 '4 s 4, which is comparable with the value of the Coriolis parameter, f= 2 sinq = 1.22x10-4 s 4. Here is the frequency of rotation of the Earth (=0.727x10-4 s 4) and q is latitude (56.7ø). Since % and f are of the same order of magnitude, a centripetal acceleration u2/r must be taken into account, where u is azimuthal velocity. Thus, instead of geostrophic balance, the cyclone has to be described by the balance of Coriolis acceleration, horizontal pressure gradient, and a nonlinear term of the equation of motion, namely the centripetal acceleration: u 2... f ug + f u (5)

8 20,924 ZHURBAS AND PAKA: THE O E VARIABILITY OF THE BALTIC SEA distance, km ',, kg m -3 Figure 8. Temperature versus potential density and distance in the cyclone, the A-B transect. By solving (5) for u, we obtain u(r) f r [1-(1 q Ug )1/2 ] (6) Equation (6) gives the following expression for the frequency of eddy rotation: r=0 4 to g 1/2 -f[1-(l+ ) ] (7) 2 f from the cyclone center and the Gotland Deep. Thus the cyclone contains the least transformed Bornholm water, which overflowed into the Gotland Basin after the recent major inflow. The thermohaline structure of the cyclone is shown in Figure 8, where isolines of temperature in the distancepotential density plane are presented. In the upper layer of the halocline, at cv0 < 6.4 kg m '3, the cyclone is purely baroclinic, and, despite a strong deformation of isopycnals (compare Figure 6), the temperature on isopycnals is nearly constant. On the contrary, in the middle layer of By substituting the above values of % and f into (7), we the halocline, for 6.4 kg m '3 < 0 < 8.4 kg m '3, the calculate 0 = 0.785x104 s ' = 1.29(/72). Thus, including temperature on isopycnals inside the cyclone is 0.2ø-0.3øC centripetal acceleration, the estimate of the frequency of below its value outside the cyclone. Finally, in the lower the cyclone rotation is decreased by a factor of 1.64 layer, adjusted to the seabed, density and salinity inside compared to a purely geostrophic estimate. Nevertheless, the cyclone are greater then their respective maximum c0 remains greater than fi2, the upper limit for the values in the surroundings. Thus the cyclone is frequency of anticyclone rotation, confirming the well- transporting heterogeneous water in the middle and lower known statement that cyclones are more energetic than layers of the halocline, while in the upper layer it appears anticyclones in both the atmosphere and the ocean. As for like a soliton-wave-like bend of isopycnals with no the maximum azimuthal velocity in our cyclone, (3) yields thermohaline anomaly. the geostrophic estimate, ug ma = 23 cm s 'l, at r = 2.94 Finally, to illustrate the transfer of salty water by the km, which decreases to the value of Um x = 16 cm s '1 at r = cyclone, Figure 9 displays the salinity field in the same 3.33 km, from (6), when centripetal acceleration is taken fragment of A-B transect. In the cyclone, the maximum into account. It should be noted that having been salinity is greater than 11.5 psu. We note that the calculated from single transect data, the above values of 0 maximum salinity in the bottom layer of the Gotland Deep and Uma are probably on the low side. was 11.0 psu in the middle of March 1993 and 11.7 psu in The TS diagram from the cyclone center is shown in June, the increase being due to the major Baltic inflow in Figure 2b by a dashed line. This estimate is located January 1993 [Dahlin et al., 1993]. Thus the cyclone is between TS diagrams from the Gotland Deep and from the transferring salty water originating from the recent major Stolpe Channel, while the TS diagram with well- inflow and is contributing to the deep water renewal in the pronounced intrusions is situated between TS diagrams Eastern Gotland Basin.

9 ZHURBAS AND PAKA: THERMOHALINE VAR2ABILITY OF THE BALTIC SEA 20, distance, km P, dbar Figure 9. Salinity versus pressure and distance in the cyclone, the A-B transect. 5. Summary and Conclusions that of the surroundings. In contrast to anticyclonic lenses that were previously observed in the Gotland Deep during In March and April of 1993, we obtained CTD profiles the stagnation period [Lips et al., 1992], our cyclonic eddy of high horizontal resolution in the Eastern Gotland seems to be a new finding. Basin. These data have allowed us to examine the The front of intrusive layering came into the Gotland mesoscale and fine scale response of the largest and Deep as far as 57 ø 19'N on April 17, which was 2-3 weeks deepest basin of the Baltic Proper to the recent major after the first observations of the bottom water renewal. inflow of salty North Sea water. This fact and our study make us think that there are at The important findings of this study are the following: 1. We observed a well-defined front of the intrusive leastwo mechanisms of deep water ventilation caused by major Baltic inflows. In the case of a gradual supply, salty region. This front was propagating north, from the Stolpe water propagates along the sloping seabed as a near- Channel to the Gotland Deep, with a speed of 2 cm s 4 or more. bottom gravitational current, filling up the deepest layer of the Gotland Basin. The salinity of these new waters is 2. We observed a spatial intermittency of intrusion greater than 11 psu, and they are distinctly seen on the TS intensity behind the front. To explain this intermittency, relations between the intrusion intensity and mesoscale curve of station BY15 (Figure 2b) due to a sharp minimum of temperature 11 psu. This mechanism dynamic and kinematic parameters were analyzed. Despite substantial scatter in the data, a dependence of intrusion probably responsible for the renewal of the deepest layer of the basin, presumably below 200 m in the case of the intensity on the mean potential vorticity for isopycnal recent major inflow. Meanwhile, the water column up to layers was found. The dependence impliesome linkage between intrusions and mesoscale eddies that transport the depth of 100 m has been ventilated, as the oxygen salty water after the inflow. concentration increased considerably in the whole layer by June 1993 [DaMin et al., 1993]. Thus a second 3. In the halocline, we observed a strongly nonlinear cyclonic eddy, which was undetectable on the sea surface. mechanism of deep water ventilation has to exist. We believe this second mechanism is mesoscal eddies which This cyclone was shown to be transporting salty water are generated in the halocline due to a time-intermittent originating from the recent Baltic inflow. The cyclone was likely generated by a nonstationary, pulsed supply of the supply of salty water. The eddies are eroded laterally by intrusions, ventilating the upper layer of stagnant water. Bornholm water into the Gotland Basin, with subsequent geostrophic adjustment of the inflowing portion of water, Acknowledgments. This research was made possible in part in which the initial density stratification was stronger than by grants and a from the Russian

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