Temporal evolution of the deep western boundary current where it enters the sub-tropical domain

Deep-Sea Research I 45 (1998) 1053—1083

Temporal evolution of the deep western boundary current where it enters the sub-tropical domain Robert S. Pickart!,*, William M. Smethie, Jr." ! Physical Oceanography Department, Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA " Lamont-Doherty Earth Observatory, Palisades, NY 10964, USA Received 28 October 1996; received in revised form 13 May 1997; accepted 27 June 1997

Abstract Four repeat hydrographic sections across the Deep Western Boundary Current (DWBC) at 55°W, occupied between 1983—1995, are used to investigate the inter-annual variability of the deep flow. The sections include measurement of tracers (oxygen, CFCs) and absolute geostrophic velocity. All properties are interpolated onto a regular grid, both in depth space and density space. The analysis focuses on two of the water masses of the DWBC: the Denmark Strait overflow water (DSOW) and classical Labrador Sea water (CLSW), both of which are clearly revealed in the property sections. The mean volume flux of water denser than p " h 27.8 kg/m3 is 13.3($4.2) Sv, comparable to that measured south of Greenland in the DWBC, suggesting that this is an accurate measure of the deep throughput into the sub-tropical North Atlantic. The largest property variability over the 12 yr period occurs in the CLSW, which in the 1990s became markedly colder, fresher, more weakly stratified, and higher in oxygen and CFCs — all indicative of new ventilation. By contrast, over this same period the deeper DSOW became less well-ventilated. Such opposite behavior of the two water masses is consistent with the large-scale atmospheric forcing in the North Atlantic. The deep absolute velocities were decomposed into a barotropic (CLSW) and baroclinic (DSOW) contribution. The DSOW flow intensified with the appearance of the new CLSW in 1991. An attempt is made to explain this, but it remains unclear to what extent the change could be due to local versus remote forcing. ( 1998 Elsevier Science Ltd. All rights reserved.

* Corresponding author. Fax: 001 508 4572181; e-mail: [email protected]. 0967-0637/98/$—See front matter ( 1998 Elsevier Science Ltd. All rights reserved. PII: S 0 9 6 7 - 0 6 3 7 ( 9 7 ) 0 0 0 8 4 - 8

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1. Introduction In recent years there has been an increased emphasis towards repeat measurements in the ocean in order to help sort out long-term, climate related variability. For instance, re-occupied hydrographic surveys in the North Atlantic have shown significant differences in the makeup of both the sub-polar and sub-tropical gyres. Multiple occupations of a section near 48°N (Koltermann, 1995) revealed significant cooling on the order of 0.4°C within the intermediate layers of the western sub-polar gyre since the early 1980s. In contrast to this, the mid-depth portion of the sub-tropical gyre has warmed by 0.1—0.2°C over a similar time period (Parrilla et al., 1994). Bryden et al. (1996) have demonstrated that this latter trend is not simply due to vertical displacement of isopycnals, but is due to variation of the water masses themselves. They speculate that it is caused largely by a reduction in import of Antarctic Intermediate water to the northern hemisphere. It is ultimately the goal in such analyses to explain why such long time scale changes are taking place, and in doing so better elucidate the ocean’s adjustment and response to climate forcing. In the North Atlantic, the Deep Western Boundary

Fig. 1. Schematic depicting the origin and pathways of the four sub-thermocline water masses of the DWBC. The 55°W section is located where the DWBC enters the sub-tropical North Atlantic.

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current (DWBC) is a major conduit for climate signals to travel equatorward from their high latitude sources. It is thus important to examine the variability of the hydrographic structure of this current; after all, climate signals should appear here first and perhaps be most unambiguous to sort out. This should in turn make it easier to interpret some of the changes that are being revealed within the interior basins. With this in mind, we have analyzed four repeat hydrographic crossings of the DWBC over a recent 12 yr period. While there are numerous locations along the western boundary where one might argue this should be done, a particularly pertinent spot is equatorward of the Grand Banks of Newfoundland. This represents the direct input into the sub-tropical gyre system and thus, in a sense, is a true measure of export away from the formation regions. Our analysis is unique in that the repeat sections contain measurements of both anthropogenic tracers and absolute geostrophic velocities, in addition to the standard ¹—S variables. This paper addresses the sub-thermocline water masses of the DWBC. There are four such water masses, originating from four different geographical areas in the sub-polar domain (Fig. 1). Our emphasis is on two of these, the water entering the current from the Denmark Strait and that from the interior Labrador Sea. The most striking changes over the period of observation occur within the Labrador Sea water layer (see also Pickart et al., 1996). However, the inclusion of tracers in our analysis has revealed a curious out of phase relationship between this water and the deeper Denmark Strait overflow water. The velocity data indicate substantial variation in the transports of both of these water masses as well. We employ a quantitative approach: the four occupations are interpolated onto an identical grid, and the analysis is performed in both depth space and density space. After discussing the variability so revealed, we attempt to shed some light on the cause of the observed changes, in particular with regard to remote versus local forcing.

2. Data and methods Since 1983 there have been four crossings of the DWBC at 55°W (Fig. 1), the most recent in 1995. Although they were parts of different experiments, the sections were oriented similarly across the topography in order to facilitate the study of inter-annual variability (Fig. 2). Two of the sections were taken during the spring season (1991, 1995), one in summer (1983), and one in fall (1994) (Table 1). We have not attempted to compensate the data for seasonality. The belief is that, for the part of the deep water column considered in this study, if any seasonal signal exists in the property fields it is not large enough to alias significantly the inter-annual trends. All cruises included a Mark III CTD system, calibrated for salinity using water samples, and measurement of dissolved oxygen and chlorofluorocarbons (CFCs). Typically there were 24 bottle samples per station. On three of the cruises (1991, 1994, 1995) acoustic transport floats were deployed during the CTD casts to measure the vertically averaged velocity, which enabled construction of absolute geostrophic velocity sections. This is important since in each case there is no deep ‘‘reference level of no motion’’, so any assumptions about such a level using water mass information

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Fig. 2. Station locations of the four repeat sections and their projections onto the line of best fit used for the gridding.

Table 1 55°W DWBC sections Cruise

Dates of section

Project

R/» Oceanus 134

4—6 July 1983

R/» Endeavor 223

26—30 March 1991

R/» Endeavor 257

10—14 November 1994

R/» Oceanus 269

29 May—1 June 1985

Gulf Stream recirculation (N. Hogg, WHOI; W. Smethie, LDEO) DWBC water mass origins (R. Pickart, WHOI; W. Smethie, LDEO) DWBC variability I (R. Pickart; W. Smethie) DWBC variability II (R. Pickart; W. Smethie)

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(for example, following Worthington, 1976) would result in error. For the 1983 occupation we have no direct velocity data — this was before the advent of either the transport floats or vessel mounted ADCPs. However, since we decompose the velocity field into barotropic and baroclinic components, the 1983 section at least affords a comparison of the baroclinic signal. The edited CTD data were used to derive the following variables: potential temperature, planetary potential vorticity, dynamic height, and neutral density (McDougall, 1987). The dynamic height was used to compute relative geostrophic velocities, which were then made absolute using the acoustic transport data. To remove the highwavenumber signal in these sections (such as topographic waves), they were low-pass filtered along density surfaces using the technique employed by Pickart and Smethie (1993). In each case the resulting filtered section clearly reveals a DWBC signal consistent with the tracer information. In order to distinguish between changes caused by local heaving of density surfaces versus water mass changes, our analysis was performed using both depth and density as independent variables. This was accomplished as follows. Using the average neutral density—depth relationship for the 1983 section, we introduced a quantity called scaled depth, which is simply the depth corresponding to a given value of neutral density using this average curve. Then the individual density values for each station on all cruises were converted to scaled depth using this relationship. The scaled depth variable serves two purposes: it enables sections in density space to be plotted in the familiar context of a depth axis (Fig. 3), and, for the density interpolation, it allows for a constant grid spacing and hence reasonable resolution throughout the water column.

Fig. 3. Potential temperature in 1991, gridded in depth space and density space.

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The variables used in this study are, therefore, depth (z), cross-stream distance (x), potential temperature (h), salinity (S), neutral density (c), potential vorticity (q), absolute filtered geostrophic velocity (v), density-scaled depth (z@), oxygen (O ) and 2 CFCs (F-11). To quantify the analysis, each section was interpolated onto a standard z grid and z@ grid. This was done by projecting the location of all stations onto the line of best fit (Fig. 2), then using spline-Laplacian interpolation. The resolution was 10 km in x and 50 m in z and z@. Gridded sections were carefully compared to analogous contoured sections using the original station spacing to assess the accuracy of the interpolation; the overall agreement was excellent (it was virtually impossible to distinguish the re-gridded sections from the originals). The consequences of density-gridding are seen by considering a potential temperature section (Fig. 3). The pronounced cross-stream tilt of the deep isotherms near 3000 m in the depth grid is the familiar shear of the Denmark Strait overflow water. This feature is absent in the density grid where these isotherms are level (i.e. isopycnals coincide with isotherms in the deep water). This is crucial for interpreting differences between sections. For instance, large deep temperature changes will occur in the depth grid if the DWBC moves laterally. Hence, if we are interested in h—S changes of the DWBC water mass, these will be revealed unambiguously only on the density grid. On the other hand, if we want to consider something such as layer thickness change, this is obviously done using the depthgridded sections. Throughout the study we use either grid, depending on our particular focus. 3. Average fields It is useful to present first the mean sections, in order to elucidate some of the general features of the DWBC. This provides a context in which to understand better the significance of the observed inter-annual variability. The depth-averaged fields are shown in Fig. 4a—h (the average potential density section is included as well for comparison to historical data). We’ve divided the sub-thermocline water column into four water masses, corresponding to previous studies (see also Fig. 1). These are, from deep to shallow, (1) Denmark Strait overflow water (DSOW), which enters the North Atlantic through the Denmark Strait (e.g. Ross, 1984); (2) Iceland—Scotland overflow water (ISOW), which passes through the Charley Gibbs Fracture Zone (e.g. Smethie, 1993); (3) classical Labrador Sea water (CLSW), which originates in the interior Labrador Sea (e.g. Lazier, 1973); and (4) upper Labrador Sea water (ULSW), which is formed along the western boundary of the Labrador Sea (e.g. Pickart et al., 1997). The density limits for each of these water masses are listed in Table 2. As mentioned in the introduction, our focus here is on the classical Labrador Sea water and Denmark Strait Overflow water. 3.1. Denmark Strait overflow water The deepest northern-source component of the boundary current, the DSOW, is clearly evident in the mean tracer fields, where it appears as a core of high oxygen

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(Fig. 4c) and high F-11 (Fig. 4d), banked against the western boundary. Its dynamic signature consists of tilted isopycnals sloping deeper offshore, indicative of increased equatorward flow of this water mass relative to the water above (Fig. 4f ). Pickart (1992) noted both features in data collected throughout the mid-Atlantic Bight but could not quantify the flow. The mean absolute geostrophic velocity section (Fig. 4g) shows nicely the DSOW velocity core, which is '5 cm/s and is aligned with the tracer core. The mean volume flux of the DSOW in our sections (1991, 1994, 1995) is 8.0 Sv. This represents a first quantitative calculation of this quantity downstream of the Irminger Sea (Dickson and Brown, 1994).1 To afford a more meaningful comparison to the Dickson and Brown (1994) study, we computed the equatorward transport of water denser than p "27.8 kg/m3. Note h that this corresponds roughly to our limits of DSOW#ISOW. Dickson and Brown (1994) described the large increase in transport of this flow between Denmark Strait (2.9 Sv) and the southern tip of Greenland (13.3 Sv). However, they were unable to partition this flow precisely between ‘‘through-put’’ and recirculation; this was partly due to the absence of appropriate measurements further downstream. Our average transport at 55°W is comparable to the value observed at the tip of Greenland (though the standard deviation is quite large, Table 3). The implication is that this value represents the through-put of the fully entrained DWBC into the sub-tropics (i.e. that portion denser than p "27.8 kg/m3). h In some ways this idea is appealing. For instance, overflow models (Smith, 1975; Price and Baringer, 1994) suggest that the entrainment process occurs over a short distance downstream from the source; this is qualitatively similar to the Dickson and Brown (1994) observations where the bulk of the entrainment happens quickly. This suggests that further downstream there should be little variation in transport, provided that the measurements are not ‘‘aliased’’ by a recirculating component or wind-driven component. At the southern Greenland site such aliasing, if any, is unknown; Dickson and Brown (1994) suggest a small recirculation, but it is largely Table 2 Water mass boundaries chosen for this study Density range Water mass

Potential density (kg/m3)

Upper Labrador Sea water

(ULSW)

Classical Labrador Sea water

(CLSW)

Iceland—Scotland overflow water (ISOW) Denmark Strait overflow water

(DSOW)

p "27.68—27.74 h p "27.74—27.82 h p "34.62—34.70 1.5 p "45.66—45.76 4 p *45.76 4

Neutral density (c) 27.81—27.90 27.90—27.98 27.98—28.04 *28.04

1 Admittedly, the average consists of only three realizations, but recall that each section is void of its high wave number (and presumably short time scale) contribution. The resulting velocity fields are thus indicative of the more slowly evolving DWBC.

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Table 3 Distribution of transport (in Sverdrups) by water mass and by year Transport by water mass Year

ULSW

CLSW

ISOW

DSOW

p *27.8 h

Total

1991 1994 1995

3.2 1.5 1.5

6.5 3.3 7.5

3.6 1.8 3.8

11.9 6.0 6.0

17.6 9.1 13.2

25.2 12.6 18.8

Average

2.1$1.0

5.8$2.2

3.1$1.1

13.3$4.2

18.9$6.3

8.0$3.4

speculative. At 55°W we also cannot definitively partition the flow. However, this is perhaps the least likely location along the western boundary between the Labrador Sea and the tropics to contain a substantial recirculating or wind-driven component. The 55°W section is positioned at the point of greatest separation between the DWBC and Gulf Stream/North Atlantic Current (Fig. 1). While there is a well-established cyclonic (largely barotropic) Gulf Stream recirculation in the mid-Atlantic Bight, it is confined to bottom depths greater than 4400 m (Hogg, 1983; Hogg and Stommel, 1985). Not only is our section shoreward of this, but the velocity core we observe is aligned with the tracer core, consistent with the absence of recirculation. This is in contrast, say, to the well-studied DWBC site near 24°N, which contains substantial recirculation at mid-depth (Lee et al., 1994; Leaman and Harris, 1990). While we do not wish to exclude the possibility of recirculation, it is certainly intriguing that our mean transport at the entry point to the sub-tropics is the same as the value obtained south of Greenland. 3.2. Classical Labrador Sea water As described by Talley and McCartney (1982), perhaps the two best tracers of CLSW are salinity and planetary potential vorticity. This is because during convective formation in the Labrador Sea, low salinity surface water is mixed downward so that the final product is a weakly stratified (low q), fresh water mass. Our average sections show a clear signature of CLSW in the DWBC at 55°W, both as a low salinity layer (Fig. 4b) and low q layer (Fig. 4e), centered near 1500 m. Note that the layer extends across the entire section, though both the salinity and q cores are intensified at the western boundary. The average oxygen section (Fig. 4c) also shows a core of high concentration near 1500 m, but this is partly an artifact of the pronounced oxygen minimum layer at shallower depths, which truncates the upper part of this core (Pickart, 1992). That such a truncation occurs can be seen in the CFC section (Fig. 4d), which shows high F-11 concentrations in both the CLSW and ULSW layers (CFCs are inert in seawater and hence are not subject to biological consumption, as is oxygen). The tracer cores of the CLSW in the average sections are as expected based on historical observations (though the strength of the signals is significantly larger than previously observed, as discussed below). Unlike the DSOW, however, there is little

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Fig. 4. The mean sections in depth space.

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Fig. 4. (continued).

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shear associated with the flow of CLSW (Fig. 4f ). This makes our measurements of absolute velocity particularly important. For instance, with a small baroclinic contribution, the choice of a reference level would be especially difficult. Using water mass considerations, Worthington (1976) used the core of the CLSW layer as a reference level. Our absolute velocity section (Fig. 4g) shows that in reality there is no such feature (at least during our time period of observation); rather, there is significant equatorward transport of this water mass. Also note the intensification of the CLSW flow near the western boundary, coincident with the shoreward enhancement of the tracer cores. The mean volume flux in the CLSW layer is 5.8 Sv, only slightly less than the transport of DSOW.

4. Variability Over our period of observation, 1983—1995, there have been pronounced changes in the properties and structure of the DWBC. These changes have not been confined to a particular water mass; there is intriguing evidence that some of the variability is coupled throughout the sub-thermocline, in spite of the fact that the different water masses have different source regions and mechanisms. It is beyond the scope of the present study to address fully the implications and causes of the variability. To understand these coordinated changes will require a broad perspective, including consideration of the atmospheric forcing of the source regions and perhaps basin-scale numerical modeling. Our intention here is simply to document the temporal changes in the deep inflow to the sub-tropical gyre, and, where appropriate, offer some insight as to the nature of the variability. We present first the property data, followed by the velocity sections. 4.1. The property fields For completeness we have included the four occupations for each variable listed in Section 2 (Figs. 5—11). On each figure it is clearly marked whether the section is in density space or depth space. The most obvious change in the temperature field is the pronounced broadening of the CLSW layer between 3.0°C and 3.5°C between 1983 and 1991 (Fig. 5a). As discussed in Pickart et al. (1997), the turn of the decade was marked by a transition towards harsh winters in the Labrador Sea, which resulted in a rapid flushing of CLSW away from the Labrador Sea. The 55°W sections continue to show a broad CLSW layer well into the 1990s. The evolution of temperature in the DWBC is more easily appreciated by computing the difference sections (in density space, Fig. 5b). The net change over 12 yr has been a cooling throughout the three deepest water masses, with the largest change ('0.25°C) occurring in the CLSW. The incremental changes are revealing as well: over the first 8 yr the upper part of the CLSW experienced the most enhanced renewal (i.e. transition to colder temperatures), whereas over the following 3 yr the renewal occurred at the densest horizons of CLSW. This is consistent with ¹—S data collected in the Labrador Sea (Rhines and Lazier, 1995), which shows that newly formed CLSW

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became progressively denser into the 1990s and has recently reached its densest level in the modern record. Finally, the anomalies over the last time period (less than a year) were different in character than the previous changes, oriented across density surfaces rather than along them. Thus the shoreward part of the CLSW layer cooled while the seaward part warmed; the reasons for this remain a puzzle. It is not necessary to compute difference sections to visualize the changes in the salinity field (Fig. 6). Again the most obvious change is in the CLSW. In the 1980s the salinity minimum of this water mass was barely evident, but by the early 1990s there is a pronounced low salinity core, representing a reduction in salinity of 0.06. Consistent with the temperature evolution, the density of this core increases further as time progresses. In regard to the DSOW, this water mass cooled and freshened over the first time period, and has since remained nearly constant. In contrast to the overall ¹—S changes in the DWBC, in which both the CLSW and DSOW evolved in similar fashion (cooled and freshened), the oxygen cores of these water masses varied out of phase with each other. Fig. 7 shows this curious relationship: the CLSW core sharply increased into the 1990s, while the DSOW core weakened. There is no clear-cut explanation for this, though in general it seems consistent with the following notion. The transition towards enhanced buoyancy forcing in the Labrador Sea noted above is tied to large-scale changes in the atmospheric circulation known as the North Atlantic Oscillation (NAO, e.g. Dickson et al., 1996). A strengthened NAO results in strong, cold winds over the Labrador Sea, while at the same time the Nordic Seas are subject to weak, variable winds. The highly coupled nature of the North Atlantic atmospheric/oceanic system with regard to this phenomenon is nicely reviewed by Dickson et al. (1996). For example, the high value of the NAO as of late has apparently helped shut down the formation of Greenland Sea Deep Water (see also Schlosser et al., 1991). While the precise formation mechanisms of DSOW remain unclear, it is nonetheless strongly influenced by air—sea interaction within the Nordic domain. Thus, one might expect a less rigorous ventilation of this water mass to accompany the enhanced production of CLSW, which is consistent with the opposite trends in oxygen we see downstream in the DWBC. Note that the F-11 sections show the marked enhancement of CLSW production (Fig. 8), but a contrary trend in the deep water will be much harder to discern, simply because of the increasing atmospheric CFC source function. However, by considering percent saturation of the DSOW F-11 core relative to the atmospheric concentration at the time of formation, we can crudely account for this secular increase in the deep F-11 data. Assuming a 12 yr tracer-age of the DSOW core (see Smethie, 1993; Doney and Jenkins, 1994), the corresponding percent saturations are 39% (1983), 22% (1991), 27% (1994), and 25% (1995). The large decline in saturation into the 1990s corresponds with the evolution of the deep oxygen signal. 4.1.1. Density layers versus tracer cores As useful as it is to examine changes in density space, in some ways it is more enlightening to consider variations of the water mass property cores; i.e. to work in tracer space (e.g. Pickart, 1992). This perhaps offers more insight into the variability of the water mass sources. Recall, for instance, that the lighter portion of the CLSW layer

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Fig. 5. (a) Individual sections of potential temperature (density space). (b) Potential temperature difference sections (differencing done in density space).

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Fig. 5. (continued).

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was renewed from the 1980s to the 1990s, followed by renewal of the denser portion of this layer (Fig. 5b). There is no a priori reason to expect that the formation of the DWBC water masses is occurring at the same density year after year (for the case of the CLSW we know this is not true); it is worthwhile then to consider variation other than within the broad density classes defined by Table 2. For these reasons we examined the water masses of the DWBC from a tracer perspective, i.e. using the tracer field to identify the location of the water mass in each section. For the DSOW we chose oxygen as the defining tracer (Fig. 7), and for the CLSW we used salinity (Fig. 6). The extent of the tracer core was defined as a percentage of the extreme value.2 Various properties were then integrated over the core region, and compared from section to section. We did the analogous integration within the DSOW and CLSW density layers defined in Table 2, for comparison. A different interpretation of the DWBC variability emerges when viewed from the water mass core perspective (Figs. 12 and 13). The densification of the CLSW in the mid-1990s noted above is clearly depicted in tracer space, due in part to continued cooling of the source product. When viewed in the context of the pre-defined density limits, one would conclude instead that the CLSW as a whole is warming over this period. For the case of the DSOW the difference between perspectives is even greater. Recall, for instance, the ¹—S evolution of the DSOW density layer, which cooled and freshened a bit over the first 8 yr then remained nearly constant. By contrast, the DSOW tracer core has large ¹—S variations year to year, including significant warming over the first 8yr. Note also the difference in oxygen signals: the DSOW oxygen core is weakest in 1991, but the density layer as a whole continues to decrease in oxygen beyond this point. This type of pronounced inter-annual variability of DSOW is reminiscent of that noted by Pickart (1992); it would be of great interest to determine definitively the cause. In Section 5 we offer some insights. 4.2. The velocity field As was the case for the property distributions, the velocity structure (and transport) of the DWBC shows significant variability from section to section. In each case there is a well-defined bottom-trapped velocity core of DSOW, but it varies in intensity and location. Interestingly (and perhaps re-assuringly), the depth of maximum speed varies similarly to the oxygen core depth, i.e. it is found shallower in the 1991 and 1995 sections, and somewhat deeper in the 1994 occupation (compare Figs. 11 and 7). In an effort to simplify the information in the velocity sections, we computed the average DWBC profile for each year (where the average was taken laterally over the e-folding width of the deep core, Fig. 14a). Each profile is similar in character, i.e. uniform over the depth of the CLSW, with increased deep flow of the DSOW. We thus defined a barotropic and baroclinic contribution as, respectively, the uniform mid-depth value

2 For oxygen we used 0.992]the maximum, and for salinity 1.0007]the minimum. These were determined (subjectively) so as to cover a reasonably large portion of the water mass. Tests with different percentages did not significantly change the results.

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and the deep excess (note: using this division we can compute the baroclinic contribution from the 1983 section, Fig. 14b). It should be noted that despite the lack of newly ventilated CLSW in 1983, this does not imply minimal transport in this density layer at that time. It remains an open issue whether the appearance of undiluted CLSW is associated with significant transport variation (for example the CLSW may act more as a dye in a pre-existing current). Unfortunately our one section devoid of new CLSW is also the only section for which we cannot calculate absolute geostrophic velocity. The variation of the baroclinic and barotropic velocity modes is intriguing. The first thing to note is that the baroclinic mode is strongest in 1991, corresponding to the arrival of CLSW in the early 1990s. This may or may not be coincidental (see next section). With regard to the 1991 and 1994 sections, one might conclude that the two velocity modes tend to vary in phase with each other; this of course was disproven during the 1995 occupation when the barotropic and baroclinic modes were comparable in magnitude. One further point of interest is that the two large values of the barotropic mode occurred during the spring season, whereas the small value occurred in the fall. Recently Pickart et al. (1997) have suggested that the 1990s vintage of CLSW is formed outside of the cyclonic gyre of the Labrador Sea, which drastically shortens the time it takes for this new water to get entrained into the DWBC. They conclude that it takes only months for the water to appear equatorward of the Grand Banks, whereas prior to the 1990s it apparently took years to a decade. Such a short time scale suggests the possibility of a seasonally varying volume flux, and one wonders if there might be a pulse of this water that appears downstream after each convective season (hence the two large springtime barotropic values). Based on our historical understanding of the CLSW circulation this would certainly be surprising, so the notion warrants careful investigation. A slopewater moored array presently in the water near 70°W will shed some light on this issue.

5. Implications The apparently coupled nature of some of the variability described above is, in many ways, puzzling (for instance the out of phase nature of the mid-depth and deep oxygen trends, Figs. 12 and 13). The history and mechanisms by which CLSW enters the DWBC are very different from those of the DSOW: the former involves open ocean convection while the latter occurs via an overflow. There is no reason to think that the time-scales involved should be similar. Thus a basic question arises: are the changes we see at 55°W locally influenced, or are they driven by variations in the source inputs (or some combination of both)? Obviously this is a complex issue, and here we consider only a couple of aspects of the observed variability in this context. 5.1. Nature of the velocity variations Is it a coincidence that the DSOW velocity core intensified with the advent of CLSW in 1991 (Fig. 14b)? One idea to consider is that the broadening of the

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Fig. 6. Individual salinity sections (density space).

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Fig. 7. Individual oxygen sections (depth space).

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Fig. 8. Individual CFC sections (depth space).

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Fig. 9. Individual potential vorticity sections (depth space).

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Fig. 10. Individual density sections (depth space).

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Fig. 11. Individual absolute geostrophic velocity sections (depth space). The sections have been filtered along density surfaces as described in the text. Note that there is no such section for the 1983 occupation.

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Fig. 12. Time series of CLSW properties, where the averaging is done (a) within the density layer (see Table 2), and (b) within the region of the property core.

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Fig. 13. Same as Fig. 12, except for DSOW.

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Fig. 14. (a) Average DWBC profiles. (b) Decomposition of the profiles into barotropic and baroclinic components (see text).

mid-depth CLSW layer (Fig. 5a) somehow impacted the deeper portion of the water column. To help investigate this we computed the layer thicknesses of the three deep water masses, the CLSW, ISOW, and DSOW (note: for this calculation we chose a slightly deeper interface between the ISOW and DSOW, c"28.07, to capture the isopycnal tilt of the latter; see Fig. 4f ). The simplest view one could imagine is that the lower water column compresses in response to a widening of the CLSW (and vice versa). However, to our surprise there is apparently a higher-order relationship in the layer thicknesses variability: whenever the CLSW layer broadens, the layer directly beneath it (ISOW) compresses, but the deepest layer (DSOW) actually stretches in concert with the CLSW layer (Fig. 15). Associated with the stretching in 1991 is an increased cross-stream tilt of the DSOW interface (Fig. 10). To help decide if such behavior of the deepest flow is a local response to the change in the CLSW, we considered a potential vorticity conserving framework to determine

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Fig. 15. Time series of layer thickness for the three deep water masses of the DWBC.

if the magnitude of deep stretching in 1991 is consistent with its increased velocity (layer tilt). If so, then perhaps the CLSW is driving some kind of local deep recirculation. In any event, equating the potential vorticity of the stretched and non-stretched states gives, f#»/¸ f#(»#*»)/¸ " , H H#*H

(1)

where H\1000 m, »\5 cm/s are the initial DSOW layer thickness and speed, f\9]10~5/s~1 is the Coriolis parameter, and *H, *» are the thickness and speed changes. For small Rossby number (»/f¸;1), which is certainly the case for the above scales, the above equation can be re-arranged: *» *H \ . f¸ H

(2)

This says that the increase in Rossby number is given by the fractional change in layer thickness. The observed increase in DSOW layer thickness between 1983 and 1991 (Fig. 10) is \50 m. This is quite large, and the above equation predicts an increase in DSOW core speed of \50 cm/s, an order of magnitude larger than observed (Fig. 14a). Thus, even though the sense of the relative vorticity change is consistent, it’s magnitude is not. This inconsistency isn’t completely surprising, since the higher mode variability in layer thickness is not easily reconciled by forcing due to the CLSW layer alone. With this in mind we mention one other intriguing observation regarding the ISOW. The decrease in layer thickness of this water mass into the 1990s is visualized nicely in the sections of potential vorticity (Fig. 9). Near 2500—3000 m note the sharp enhancement in 1991 of the potential vorticity maximum between 100—150 km; high values persisted in this region in the next two sections as well. It also happens that this portion of the layer has been more recently ventilated in the 1990s, as revealed by the CFC sections (Fig. 8). Note the appearance of high F-11 concentrations (clearly above the density of the DSOW) coincident with the high-potential vorticity region. This suggests more than a simple local response to the layer broadening of the CLSW, in particular, that the source history of the ISOW might have changed as well.

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Fig. 16. Schematic of the overflow process modeled by Price and Baringer (1994), where o and ¹ denote density and transport. Source water (o , ¹ ) flows through the Denmark Strait and entrains oceanic water 4 4 (o ), forming the final product (o , ¹ ). The location of Ocean Weather Station Alpha is indicated by the 0 1 1 solid square.

5.2. Nature of DSOW core variations Using the simple framework of Price and Baringer’s (1994) overflow plume model, we investigated the nature of the DSOW property core variability (Fig. 13). We focused on the most striking change in the core, that which occurred from 1983 to 1991. Over this time period the DSOW became less dense and lower in dissolved oxygen, and its transport increased. Price and Baringer (1994) used the terminology ‘‘source water’’ to denote the overflow water exiting Denmark Strait, ‘‘oceanic water’’ as the water entrained into the current downstream of the Strait, and ‘‘product water’’ as the final mixture appearing downstream (Fig. 16). We assume that our 55°W section depicts the final product, and ask if the changes seen between 1983 and 1991 are easily explained by variation of the source water, or oceanic water. One of the basic results from the Price and Baringer (1994) analysis is that, in many ways, the downstream product is more sensitive to changes in oceanic entrainment than to source variability. This is due to the damping influence of the entrainment, whereby significant anomalies in the source are strongly moderated downstream. Applying Price and Baringer’s (1994) end-point model, we first considered changes in the overflow source. The observed density change at 55°W implies a temporal change

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Fig. 17. Average summer density profiles at OWS Alpha (see Fig. 16). The sill depth of the Denmark Strait is indicated by the dotted line. The variation in density of the oceanic water (*o ) implied by the 55°W data 0 is shown.

in source density of \0.09 kg/m3, which is comparable to the total downstream change in a given year. While possible, this does not seem likely. Furthermore, if the source density decreases this leads to less entrainment and hence a decrease in product transport, which is the opposite of what we observe. Finally, less entrainment suggests less dilution of the oxygen signal, i.e. a higher product concentration. Contrary to this, we observe a significant decrease in the DSOW oxygen core concentration. (Note: about 50% of this decrease can be accounted for by the change in solubility due to the temperature increase). It seems unlikely, therefore, that the observed DSOW change is simply due to a change in the source water. Regarding the surrounding oceanic water, our data require a temporal change in the density of this water of \0.018 kg/m3. To determine if this is feasible we used time series data from Ocean Weather Station Alpha, occupied until 1970 in the Irminger Sea. This is near the region of large entrainment of DSOW (Fig. 16) and near the station used by Price and Baringer (1994) as their oceanic profile. Using an average over the summertime months (most data coverage), a comparison of density profiles over a 17-year time period shows substantial inter-annual variability near the level of the deep overflow (Fig. 17), enough to account for the observed change at 55°W. Also, the sense of the corresponding transport change and oxygen change downstream are in line with what we observe; i.e. smaller oceanic density means more entrainment (because of the larger initial density

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difference between source water and surrounding water), hence larger transport and less oxygen. The major shortcoming with this scenario, however, is the magnitude of the transport enhancement. The Price and Baringer (1994) model implies a transport increase which is over an order of magnitude smaller than that observed. Thus, the intensification of the DSOW flow into the 1990s remains a puzzle.

6. Summary Four repeat hydrographic/tracer/velocity sections across the DWBC at 55°W, from 1983 to 1995, have better elucidated the properties and transport of the current and revealed significant variability throughout the deep water column. The average fields clearly depict the two major water masses, the DSOW and CLSW. The mean deep transport of water greater than p "27.8 kg/m3 is 13.3 Sv, which is comparable to the h value observed further upstream south of Greenland (Clarke, 1984). Placing this value in the context of historical measurements (Dickson and Brown, 1994) implies that it is a true measure of the export into the subtropical domain. The most conspicuous inter-annual changes in the DWBC have occurred in the mid-depth CLSW layer. In 1983 this water mass was barely evident, but in all three sections occupied in the 1990s there is a strong signature of cold, fresh, newly ventilated, weakly stratified water. This is consistent with the rapid flushing hypothesis of Pickart et al. (1997), who believe that CLSW has recently been formed outside of the Labrador Sea cyclonic gyre and hence more easily enters the DWBC system. The 55°W tracer data have revealed that the deeper DSOW has decreased in oxygen content, out of phase with the CLSW. This seems reasonable in light of the large-scale atmospheric forcing of the northern Atlantic; the NAO has recently favored deep ventilation of water in the Labrador Sea, while apparently diminishing such ventilation in the Nordic domain, from which the DSOW originates. The DWBC velocity field at 55°W also showed significant variation between sections. Splitting the flow into a barotropic (CLSW) and baroclinic (DSOW) mode revealed an increase of the DSOW velocity coincident with the appearance of the new CLSW in 1991. It also suggested that there might be a seasonal flux of CLSW, with stronger flow in the spring and weaker flow in the fall. The latter notion is consistent with the short CLSW flushing time recently computed by Pickart et al. (1997). The strengthening of the DSOW flow, however, remains a mystery. Acknowledgements Such an analysis of multiple data sets requires the assistance of many people. The authors gratefully acknowledge the help of Terry McKee, who contributed in numerous ways to this project. Daniel Torres helped with the computation of velocities. The CTD group of the Woods Hole Oceanographic Institution acquired and processed all of the hydrographic data sets in their typical expert fashion. Bathysystems built the transport floats that allowed the determination of absolute velocity. We thank the

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crews of R/» Oceanus and R/» Endeavor for helping make the fieldwork so successful. Finally, we acknowledge the National Science Foundation which supported each of the experiments, the latest grants from which are OCE-9301448 (R.P.) and OCE9302243 (W.S.).

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