Cold anticyclonic eddies formed from cold pool water in the southern Middle Atlantic Bight

ContinentalShe(fResearch,Vol.17,No. 15.pp. 1839 1867.1998

Pergamon

PII: S0278--4343(97)00038--1

(!"!1998ElsevierScienceLtd.Allrightsreserved Printedin GreatBritain 02784343/98$19.00+ 0.00

Cold anticyclonic eddies formed from cold pool water in the southern Middle Atlantic Bight CHARLES N. FLAGG,*t DOUGLAS WALLACE* and ZBIGNIEW KOLBER (Received 14 December 1995; in revised.form 7 April 1997; accepted 24 April 1997) Ahstract--AVHRR satellite imagery of the southern Mid-Atlantic Bight during May 1993 revealed a large area of cold water over the shelf break and slope that appeared to spin up into a series of southward propagating anticyclonic eddies. The eddies had diameters of 35-45 km at the surface and moved southward at about 20 cm/sec. A radial TOYO CTD (to 50m) and ADCP velocity (to 400m) transect was conducted across the southern-most of these eddies. The upper 50 meters had minimum temperatures of less than 7°C and salinities of about 33 pss, characteristics similar to cold pool waters usually found over the continental shelf. ADCP velocity data from one of the eddies revealed anticyclonic flow extending to a depth of about 250m. The transport of cold pool water by the eddies was estimated to be 0.1 to 0.2 Sv which is of the same order as the annual mean alongshore transport of shelf water in this region. The origin of the deeper water within the eddy is unlikely to be the continental shelf because the shelf break is less than 100 m. The depth and velocity profiles along the TOYO transect were consistent with the constant potential vorticity eddy model of Flierl (1979) although the source of the eddy kinetic energy is uncertain. The cause for the exodus of cold pool water from the shelf, which extended northward to at least 38°N, is unclear but must involve the establishment of an alongshore baroclinic pressure gradient against the usual southwestward shelf flow. It is possible that the intrusion of Gulf Stream waters onto the shelf near Cape Hatteras was a precursor of this off-shelf transport. The southern-most eddy was marked by high biological productivity and very high oxygen supersaturation. The phytoplankton bloom detected within the exported cold pool water, located over the continental slope, suggests a mechanism whereby production fueled by nutrients derived from the shelfcan be locally exported into deep water. © 1998 Elsevier Science Ltd. All rights reserved

1. INTRODUCTION The mean currents on the shelf of the Middle Atlantic Bight (MAB) range between 5 and 10 cm/sec toward the west and southwest with a total alongshore transport of approximately 0.4 Sv south of New England decreasing to about 0.2 Sv off Maryland (Beardsley and Boicourt, 1981; Biscaye et al., 1994). There is considerable exchange of water across the shelf-slope frontal boundary so that by the time water reaches the offing of Maryland and Delaware approximately half the water on the shelf has its origins in the slope water the remainder being water that entered the Bight past Nantucket shoals (Fairbanks, 1982; Falkowski et al., 1988; Biscaye et al., 1994). Approaching the southern terminus of the Bight at Cape Hatteras, there must be an even more dramatic export of shelf water. The * Oceanographic and Atmospheric Sciences Division, Department of Applied Sciences, Brookhaven National Laboratory Upton, NY 11973, USA t Corresponding author. 1839

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shelf's cross sectional area at Cape Hatteras (36°N) is 50 percent of that south of New England and 70 percent of that off Maryland while the mean currents are approximately the same. As a result, the alongshelf transport must decrease rapidly as it approaches Cape Hatteras. In addition, the mean northeastward flow from the South Atlantic Bight (Stefansson et al., 1971; Lee and Atkinson, 1983) serves to block a permanent southward flow past the Cape. Therefore on average, all the water flowing to the south on the shelf must exit, by one means or another, by the time the flow reaches the Cape. This paper focuses on one method by which part of the shelf water mass, the cold pool, was observed to exit the shelf during the spring of 1993. Cold pool water, first described by Bigelow (1933), is a key component of shelf waters in summer. The cold pool waters are formed in the winter when the shelf waters reach their annual temperature minimum and are relatively well mixed vertically. During the vernal warming a portion of this cold water over the outer half of the shelf becomes isolated from the surface by the seasonal thermocline. Further inshore, vertical mixing due to tides and wind serves to transfer some of the heat downward warming the deeper waters relative to the cold pool. The cold pool is separated from the warmer, more saline and denser waters offshore by the sub-surface expression of the shelf-slope front. The cold pool water moves alongshore at typical speeds of about 5 cm/sec (Houghton et al., 1982) replenishing the water to the south during the summer with colder water from the north. Because of this alongshore movement and the limited mixing of heat across the seasonal pycnocline (e.g. Gordon and Aikman, 1981; Wallace, 1994), the minimum bottom water temperatures in the southern MAB occur in June or July rather than during the coldest months of the winter. The cold pool waters must also leave the continental shelf "upstream" of Cape Hatteras. There has not been a concerted effort to date to describe this phenomena but several authors have commented on where the cold pool must be exiting the shelf. Because of the large density difference across the seasonal pycnocline, significant baroclinic motions are possible and the waters of the cold pool can act quite separately from the warm near surface waters. Boicourt (1973) made a number of closely space hydrographic observations in the southern MAB and noted that the location and intensity of the cold pool was variable from month-to-month and from year-to-year. Garwarkiewicz et al. (1992) made a similar comment noting that the extent of southward penetration of cold pool water along the shelf was variable, as was the apparent location of the "detachment" of the water from the shelf. More recently, Garwarkiewicz et al. (1996) looked at the departure of shelf waters and noted that Gulf Stream intrusions at the southern end of the Bight totally changed the offshore baroclinic pressure gradients and could act as a "blocking" agent forcing the shelf water offshore. There have been a number of studies of the incorporation of shelf waters along the north wall of the Gulf Stream (Ford et al., 1952; Kupferman and Garfield, 1977, and Lillibridge et al., 1990; Wood et al., 1996) where the temperature and salinities of the incorporated waters indicate not only that they came from the shelf (Stommel, 1966), but that some of the water came from the cold pool. A study that is particularly relevant to the observations reported here is that of Houghton et al. (1986) and Marra et al. (1990) in which they observed a bolus of cold pool water south of New England apparently exiting the shelf and forming an anticyclonic filament. Even though these results came from an area far removed from Cape Hatteras and the filament was isolated and had not formed itself into a complete eddy, it had a number of features that were similar to the cold eddies described here. The relationship between these studies will be discussed further below.

Cold anticycloniceddies formedfrom cold pool water in the southern MiddleAtlantic Bight

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The motivation for the collection of much of the data reported in this paper came from real time satellite AVHRR images acquired during a hydrographic survey of the southern MAB and Cape Hatteras region in May 1993. The images indicated the presence of a series of cold anticyclonic features progressing southward along the edge of the continental shelf (Fig. 1). At least three cold eddies were identified, the southern most of which was located at approximately 36.1 ° N, 74.5 ° W on May 15 1994, within range of our survey program. We were able to conduct a TOYO CTD and shipboard ADCP transect starting in the center of this eddy and progressing along a radial line to the northwest. This transect was the final activity of a five day survey of the continental shelf between 35.5 ° N and 36.5 ° N, which also included three cross-isobath transects and a north-south TOYO section approximately along the 40 meter isobath (Fig. 2). The cruise was a multi-disciplinary effort during which a number of physical, chemical and biological parameters were measured. In this paper we describe the cold eddy in terms of temperature, salinity, density, dissolved oxygen, fluorescence, and surface nutrient distributions relative to those same parameters observed on the shelf and in nearby waters. We then discuss the dynamics of the eddy and make some inferences as to the origin of the eddies and their significance in terms of the export of shelf waters and carbon to the continental slope environment.

2. OBSERVATIONS 2.1. A V H R R sea surface temperatures

The near-realtime AVHRR images (Fig. 1) were acquired through Ocean Imaging Corp's facilities in California. The cruise took place from May 10 through May 16, 1993 and images were available from a few days before the cruise started. All the images shown are from local morning satellite passes which are less subject to atmospheric interference and the masking of features due to diurnal thermocline formation. The images cover the southern MAB from 35 ° to 38°N showing the very warm waters of the Gulf Stream, the intermediate temperature waters from the slope, a relatively warm outflow from the Chesapeake, and low temperature waters just off the edge of the shelf. In hindsight, the sequence of sea surface temperature images shows the formation of a series of three to four anticyclonic eddies although at the time the nature of the evolving field only became obvious after the passage of a cloudy period that lasted from May 12 through 14. Figure la shows a broad rather amorphous region of low surface temperature centered over the shelf break (as indicated by the 100 meter isobath). North of 37.5 ° N this low temperature region covers the outer half of the continental shelf and extends out into the slope region by some 50 to 100 km. This cold surface water extends southward as far as the north wall of the Gulf Stream, with its leading edge being swept off to the east. Drier atmospheric conditions on May 10 (Fig. lb) revealed more detail and we began to see cold circular, eddy-like features and extended cold filaments with temperatures less than 10°C. The centers of the cold features were still located over the shelf break except at the southern end near the Gulf Stream. By May 15 (Fig. I c) and 16 (Fig. I d) the anticyclonic nature of the eddies was clear from the way different temperature filaments spiralled out from the center of the eddies. At this time there appeared to be three eddies at about 100 km intervals. In contrast to the pict-,we on May 10, by May 15 the cold water had moved offshore and was no longer located over the shelf break. From a combination of the clear images shown in Fig. 1 plus some partially obscured images not shown, the eddies in Fig. lb appeared to move

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C.N. Flagg et al. C a p e Hatteras R e g i o n

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~ 85 k m south between M a y 10 and 15 and the southern-most feature in Fig. Ib (Eddy #1) was stretched out along the north wall of the G u l f Stream in Fig. lc. This indicates a southward eddy propagation speed of about 17 km per day, or 20 cm/sec. It is noticeable that as the eddies reached the G u l f Stream their surface expression became distorted, perhaps as a result of being over-run by warm surface waters of Gulf Stream origin. This layering is possible because the near-surface sigma-t of the eddies was about 2 5 . 5 whereas the G u l f Stream's near surface density was one full sigma-t unit lower (see below). These densities and density differences are reminiscent of the subsurface low salinity filaments in the G u l f Stream's north wall observed by K u p f e r m a n and Garfield (1977). The satellite images also show the presence of relatively warm shelf water originating from the Chesapeake Bay outflow. In Fig. l a + b the warmer Chesapeake Bay plume occupied much of the shelf from Cape Hatteras to about 20 km north of the m o u t h of the Chesapeake Bay at 37°N. Hydrographic data from near shore stations indicated that the plume waters had salinities less than 30 pss. The wind close to the Chesapeake mouth (Norfolk, Virginia) during the period leading up to the cruise was generally from the south at between 2 and 10 kts. At Cape Hatteras, the winds were stronger and from the southwest

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Fig. I. A succession of AVHRR sea surface temperature images of the southern Mid-Atlantic Bight during the study period showing the development of the cold anti-cyclonic eddies and their interaction with the shelf and Gulf Stream. In the clear sky images shown here the blanked-out pixels are the result of ground fog, not clouds per se, brought about by contact between the cold surface waters and the warm moist air from the southwest. The succession of three cold eddies observed during this period have been labeled. The heavy black line on each of the figures is the 100 m isobath indicating the approximate edge of the continental shelf.

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during the cruise. The wind direction is indicated by a string of thin low-level clouds trailing offthe coast to the northeast, north of 36°N in Fig. lb. During the cruise the southerly winds appeared to have retarded the normal southward near-coastal flow of low salinity water at the surface, and the near surface Ekman transport had driven this low-density water well out over the shelf. Extensive flooding of the shelf with low salinity water appears to be common. It was clearly evident during the 1989 SEEP-II experiment, conducted further north, when salt marsh grass and very low salinities were observed at the shelf edge at 37.5 ° N (Flagg et al., 1994). The effect of offshore Ekman transport of near surface water is also reflected in the coastal upwelling of colder subsurface water near shore between Hatteras and the Chesapeake (Fig. ld). The entrainment of shelf water by Gulf Stream frontal eddies was also observed during this period. Although not completely evident from Fig. 1, the entire series of images during this period shows that there were four Gulf Stream frontal eddies that passed through the region during this period. The remnants of the first eddy are visible at the eastern portion of Fig. la. The second and third eddies are visible in Fig. lb with the third eddy almost curling back on itself and disappearing to the east in Fig. lc. The last and largest eddy of the set is just visible in the southwest corner of Fig. lb and in its mature form in Fig. lc and ld. Streamers of relatively cold, low salinity shelf waters visible along the north wall of the Gulf Stream are clear examples of the "Ford water" (Ford et al., 1952; Kupferman and Garfield, 1977; Lillibridge et al., 1990). For the earlier and smaller frontal eddies, the entrained water came from the slope which at this time was occupied by cold pool derived water (see below). The last and largest eddy formed further south and crested in the vicinity of Cape Hatteras with the result that it preferentially entrained, at least at the surface, low salinity shelf water from the Chesapeake, as well as colder, upwelled, nearshore shelf water. The initial width of the shelf water filament drawn offshore by this last eddy was about 20 km but within 30 km it narrowed down to ~ 10 km, suggesting either a doubling of its thickness or, more likely, a doubling of its eastward current speed. No hydrographic sections were made through the eddy during this period but it is reasonable to assume that it bore a structural resemblance to the Gulf Stream frontal meanders and "shingles" observed in the South Atlantic Bight (e.g. Bane et al., 1981). Hence along the north wall of the Gulf Stream there was a layering of low salinity shelf water over Gulf Stream water which in turn had overridden cold pool water of intermediate salinity that was moving southward along the shelf edge (see also Lillibridge et al., 1990). 2.2. TO YO section and surface sampling through cold eddy The realtime AVHRR imagery encouraged us to alter the cruise plan in order to conduct a five hour, 35 kilometer TOYO CTD transect from the center of the southern most cold eddy to the northwest beginning May 5, 1993 at 2200 GMT (Fig. 2). The TOYOed CTD was a NBIS MK-III equipped with a Beckman dissolved oxygen sensor and a custom built fluorometer (see below). A stabilizing fin was attached to the instrumentation and the package was raised and lowered at 40 meters per minute from the surface to a depth of approximately 50 meters while the vessel steamed at three to four knots. The CTD's oxygen sensor was calibrated (Owens and Millard, 1985) based upon Winkler titrations from 311 discrete samples analyzed during three cross-shelf transects occupied immediately prior to the TOYO transect. The standard error of the fit to the bottle oxygen data was less than 6 #mol/kg.

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Chlorophyll biomass and photosynthetic parameters of the phytoplankton such as variable fluorescence, photosynthetic quantum yield, and functional absorption crosssection were measured using a Fast Repetition Rate (FRR) fluorometer (Kolber and Falkowski, 1992). Chlorophyll concentration was assessed from variable fluorescence using a calibration coefficient calculated from a linear regression between extracted chlorophyll and FRR fluorescence from stations occupied prior to the TOYO deployment. Instantaneous primary production was calculated from the FRR derived photosynthetic parameters as in Kolber and Falkowski (1993). The fluorescence-based estimates of primary production were converted to net primary production (carbon fixation) using an empirical calibration between FRR and 14C measurements from stations occupied prior to the TOYO deployment. This calibration, based upon 39 data points, yielded an r-squared of 0.81. Daily primary production along the TOYO section was calculated assuming that the photosynthetic parameters were constant during the day with the radiance profile in the water column calculated at one hour intervals based upon the surface PAR (photosynthetically active radiation). The light propagation in the water column was calculated from an empirically derived extinction coefficient of the water plus chlorophyll biomass (estimated from FRR fluorescence) at the time of the TOYO deployment. Horizontal resolution along the TOYO track was about 0.6 km at 25 meters depth, decreasing to about 1.2 km near the surface and at 50 meters depth. A 150 kHz ADCP was also collecting data during the cruise using five minute ensembles with 4m bin lengths and 4m or 6m pulse lengths depending on the water depth. Starting over approximately the 100 meter isobath at ~ 35.7°N and proceeding along the cruise track to the beginning of the TOYO transect and thence along the transect, samples for dissolved nutrient (NO3, NO:z, NH4, PO4, SiO4) analyses were collected from the ship's clean pumped surface seawater supply and immediately frozen for later analysis. The seawater intake was at a depth of 2 meters below the surface. Samples were collected approximately every 15 minutes during steaming, corresponding to a horizontal resolution of every 6 km south of the TOYO and 2.5 km along the TOYO transect. In the laboratory, samples were thawed in a room temperature water bath, shaken, and promptly analyzed on a Technicon AAII autoanalyzer system (Whitledge e t al., 1981). Sections from the TOYOed CTD along the cold eddy transect are shown in Fig. 3 (sections are oriented looking south), with the associated temperature-salinity (T-S) diagram in Fig. 4. The clearest depiction of the cold eddy is in the temperature section, Fig. 3a, which shows a near surface front across which the temperature changed by about 3.5°C between the eddy (temperatures of 10°C to 11°C), and the shallow shelf waters further inshore (temperatures > 14°C). This temperature range is consistent with the AVHRR SST image for the morning of May 16, 1993, Fig. ld. Below the thermocline the cold eddy signature was also quite evident, with a large bolus of cold water less than 8°C and with minimum temperatures of 6.6°C. There was also a subsurface front between the cold eddy water and warmer water ( > 9°C) further onshore which separated the cold water from the shelf break. It appeared that the deepest extent of the cold core of the eddy, as demarcated by the 8°C isotherm, was about 50m. Note that the approximate depth of the shelf break south of Chesapeake Bay is also ~ 50m, as is the vertical extent of the cold pool waters found at the edge of the shelf further to the north. The salinity section, Fig. 3b, shows that there was also a shallow salinity front between the cold eddy and the less saline shelf water, which was coincident with the thermal front. The salinity difference between the center of the cold eddy and the shelf water was about 1 pss.

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Fig. 4. Temperature/Salinity correlation plots from two portions of the radial TOYO section, a) from the cold core of the eddy between the beginningof the transect at 36.07°N, 74.50°W, and a point 22.25 km to the northwest at 36.20°N, 74.680W, and b) from the region of intruding shallow shelf water between a point 24.0 km to the northwest at 36.21°N, 74.70°W, and the end of the section at 36.28°N, 74.79°W. This difference decreased with depth and below 30 meters the cold eddy waters were only slightly less saline than the waters closer to shore (at the same depth). While the temperature section indicated a thermostad between 20 and 50 meters (AT <0.2°C) capped by a thermocline, the salinity increased nearly linearly with depth from a surface value of 32.8 to 33.4 at 50 meters. This vertical T-S structure is a characteristic of cold pool water found along the outer shelf of the MAB. Figure 4a presents the T-S relation for T O Y O data collected within 22.25 km of the beginning of the transect (i.e., within the cold eddy). These data indicate minimum temperatures below 7°C at about 33 pss with a very tight correlation toward higher slope water salinities and a broader correlation rising to surface temperatures of 11 °C to 13°C at about 32.8 pss. Shoreward of the cold eddy, the T-S relation (Fig. 4b) is dramatically different and indicates minimum temperatures of about 8°C, which were broadly distributed between salinities of 32.3 pss to 33.4 pss, whereas shallow water had salinities of 32.0 to 32.3 and temperatures of 13°C to 14°C. At depth there was some overlap of water types at salinities of 33.5 pss but in general the nearshore water was almost two degrees warmer than the cold eddy. The density section of Fig. 3c shows an upward doming of the pynocline above 20 meters centered about 7 k m into the section. Below 30 meters, isopycnals were displaced downward within the cold eddy. The presence of lower salinity, warmer water derived from the shelf seems to have depressed the isopycnals over the nearshore third of the section. The impression given by the density section is of a cyclonic feature, at least within the upper 30 to 40 meters, but this impression is misleading. A D C P data taken over this transect showed a strong anticyclonic flow extending over the eastern two-thirds of the section, in agreement with the flow pattern inferred from the A V H R R images. The 26 sigma-t isopycnal between 35 and 50 meters also suggests an anticyclonic feature, but the geostrophic reference level is unknown.

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The dissolved oxygen distribution within the cold eddy section, Fig. 5a, shows rather different features than do the temperature, salinity, and density plots. There was a very strong maxima in both oxygen concentration (see Fig. 6a) and percent saturation (not shown) within the thermocline. The oxygen maximum shoaled towards the northwest (i.e. shorewards) along with the isotherms. In contrast, the isopycnals sloped downward towards the northwest. The depth of the oxygen maximum appeared to track the local maximum in the vertical density gradient within the cold eddy while it was located just below the maximum density gradient within the fresher water to the northwest. Within the pycnocline, the oxygen maximum was strongest within the cold eddy. Oxygen concentrations were supersaturated by up to 115 #mol/kg within the cold eddy (at depths of 10-25m, sigma-t of 25.4 to 25.8), whereas to the northwest, outside the eddy, peak saturation anomalies were < 95 #mol/kg (at depths of 10-15m; sigma-t of 24.6-25.2). This pattern was also seen at the sea surface, with cold eddy saturation anomalies of 75-90 #mol/kg, whereas to the northwest, anomalies of 60-80 #mol/kg were encountered. Surface pCO2 (D. Chipman, pers. comm.) was undersaturated throughout the TOYO transect and inversely correlated with surface oxygen supersaturation, with the lowest pCO2 found within the cold eddy, and higher pCO2 to the northwest. An inverse correlation between surface oxygen and CO2 saturation indicates that gas saturations were strongly controlled by net biological production rather than by temperature changes. This implies that the location of the oxygen maximum noted above is most likely a consequence of the vertical distribution of primary productivity. Co-location with a local maximum in the vertical density gradient may reflect a balance between the supply of nutrients from below via cross-isopycnal mixing, the (lightrelated) nutrient demand of phytoplankton and the loss of oxygen via air-sea gas exchange which can take place from the surface layer. The contrast in oxygen saturation between the cold eddy and its surroundings was, however, reversed in waters deeper than ,-~25m, for which the supersaturation was found to be higher inshore (20-30 #mol/kg) than within the eddy (0-20 ktmol/kg). At the inshore edge of the cold eddy (23 to 27 km from the beginning of the transect) there

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40

was a strong signal of highly oxygenated waters at depths of 40-50m. Within this feature, high oxygen supersaturation was found at significantly greater densities and salinities than anywhere else in the TOYO transect. The oxygen concentrations found within this feature were also higher than any measured within the same density range in any of the cross-shelf transects occupied during the cruise. The high oxygen core was correlated with a local upwards displacement of the 33.2 to 33.8 isohalines and the 26.0 sigma-t surface. The FRR fluorometer mounted on the TOYO indicated that primary production was higher in the cold eddy than in the warmer waters to the northwest, confirming the impression given by the oxygen distribution. In Fig. 5b the fluorescence based chlorophyll distribution was very similar to the dissolved oxygen distribution although its peak was located 5-10m below the O2-maximum. The highest chlorophyll concentrations were generally found just below the thermocline/pycnocline, although the peak concentrations were found within the anomalous feature at the edge of the eddy discussed above. There were slightly higher chlorophyll concentrations under the warm shelf water than in the center of the eddy itself. The FRR fluorometer showed, however, that the greatest potential for photosynthesis was actually within the cold eddy as shown by the distribution of a'A, Fig. 5c. Somewhat simplistically, a, the functional absorption cross-section of photosystem II, is a measure of the efficiency of light utilization by the phytoplankton, whereas A is a

Cold anticycloniceddiesformedfromcold pool waterin the southernMiddleAtlanticBight

1853

measure of the quanta of photosynthesis. The product of a*A describes the bulk chlorophyll-normalized photosynthetic potential; its product with irradiance and chlorophyll biomass yields a measure of the total photosynthesis per unit volume. Thus, Fig. 5c shows that the greatest photosynthetic potential and thus the greatest potential for oxygen production was centered within the cold eddy. That the high surface oxygen values were produced locally through photosynthesis is also illustrated by a comparison of salinity/oxygen correlations from within the cold eddy and from a nearly contemporaneous cross-shelf section at 36.5°N, Fig. 6. Figure 6a shows an abrupt rise in dissolved oxygen concentration of about 80 #moles/kg (corresponding to a saturation value of about 135 percent) at the oxygen maximum (just above the depth of the maximum chlorophyll concentration), followed by a decline toward the surface. The lower near-surface oxygen levels reflect the effects of out-gassing although diminished primary production in the upper layer because of nutrient depletion probably also contributes. A comparison of this curve with a similar curve from stations sampled over the shelf during the same cruise (Fig. 6b), and with historical data from the Middle Atlantic shelf collected at the same time of year (not shown, see below), showed that the elevated oxygen concentrations found in cold eddy waters with salinities between 32.8 and 33.1 pss (i.e. from depths > 30 m) were completely absent from the shelf. We therefore hypothesize that a high level of photosynthetic activity and oxygen production took place within the cold eddy after its waters had risen to the surface and were exposed to photosynthetically active radiation. The distinct quality of the water within the cold eddy is evident in a plot of the surface nutrient concentrations during the northeastward run up to the beginning of the cold eddy TOYO and out along the TOYO section, Fig. 7. All the nutrient values were low, as one would expect for near surface waters, nevertheless coherent spatial variations were discernable. The distributions (Fig. 7), particularly that of phosphate, demark three key water types which were clearly visible in satellite imagery (Fig. 8). Surface nutrient sampling began in water of Gulf Stream (and perhaps South Atlantic Bight) origin with low phosphate and N + N (NO3 plus NO2) concentrations but relatively high silicate (B in Fig. 8). About 25 km south of the beginning of the TOYO the ship passed into lower salinity water of shelf origin (C in Fig. 8) which was marked by significantly increased phosphate, slightly higher N + N and lower silicate. The low silicate is likely a result of the typically diatom dominated northeast shelf phytoplankton community, with the higher silicate values to the south resulting from either (1) an ecosystem which is less dominated by diatoms, or (2) subject to higher silicate supply from underlying waters. Near identical nutrient levels were encountered 25 km northwest of the beginning of the TOYO section (E in Fig. 8) reflecting another band of water of shelf origin. The surface water found within the cold eddy itself (D in Fig. 8) had lower phosphate than the immediately adjacent shelf water, and levels that were only slightly higher than those in the Gulf Stream. The N + N was variable, silicate was less so, and neither were greatly different than levels found in the adjacent shelf waters. Without much hydrographic data below 50 meters (XBTs were not available) it is difficult to determine the source of the deep waters involved. The deepest data we have is from the first TOYO cycle when the ship was moving slowly and the TOYO reached 80 m, Fig. 9a. The profiles of temperature, salinity, and dissolved oxygen all contain values typical of those found over the outer shelf and well above the sub-surface salinity maximum just offshore of the shelf-break. Typically, the 34 pss isohaline is found within the shelf/slope front at depths less than 50 m as shown in a cast taken in May 1989 at the edge of the shelf at 36.5°N, Fig.

1854

C.N. Flagg et al.

(a)

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T e m p e r a t u r e (C) Fig. 9. a) Temperature, salinity, and dissolved oxygen profiles to 80 m at 36.08°N, 74.51°W, from the first up-down cycle of the TOYO section, b) Temperature, salinity, and dissolved oxygen profiles from a May 1989 station at 36°52.4"N, 74°42.4'W in approximately 90 meters of water.

Cold anticyclonic eddies formed from cold pool water in the southern Middle Atlantic Bight

RN

Gyre.

16 May

1855

1993

25

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bh0516am.sv Fig. 8. Sea surface temperatures from an AVHRR image from the morning of May 16, 1993, just after the completion of the northwestward TOYO section of the cold eddy. Upper layer current vectors from the Gyre's shipboard ADCP have been overplotted to show the relation between the hydrographic structure and currents. The position of the surface nutrient observations used in Fig. 7 are shown as open circles. The heavy black line indicates the 100 m isobath, the approximate location of the shelf break.

74

(a) 74.~ W

36.~N 74,79W

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Fig. 10. a) C o n t o u r plots of A D C P current components for the 35 km northwestward T O Y O section down to 400 m (looking southwestward). The coordinate system has been rotated to normal to (U42) and along (V42) the T O Y O section as shown in Figure 2. The depth profile appropriate for a modeled constant potential vorticity eddy (Flicrk 1979) has been ovcrplottcd on the U42 section plot (see text), b) C o n t o u r plots of A D C P current components for the 75 km run up from 35.5°N on the shelf to the beginning o f the T O Y O section (looking northwestward). The current coordinate system has again been rotated into components along (U37) and normal (V37) to the cruise track, see Fig. 2. The A D C P data were averaged over l 0 m in the vertical and 10 minutes in time. Blanked pixels reflect the bottom or limits on thc A D C P ' s range. The current isopleths arc at 10 cm/scc intervals and correspond to the velocity scale on the right.

Cold anticycloniceddies formedfromcold pool water in the southern MiddleAtlantic Bight

1857

9b. At the eastern end of the TOYO transect, by contrast, the 34 pss water was found more than 30 km offshore of its usual location and some 20 to 30 m deeper than normal. Below this level we have no hydrographic information, leaving the source of the deep water associated with the eddy to be surmised from the ADCP velocity measurements. 2.3. Velocity data from a shipboard A D C P An RD Instruments 115 kHz hull-mounted acoustic Doppler current profiler was in operation throughout the cruise. The data were collected using RD Instruments' DAS version 2.48 and processed using the CODAS data processing and archiving system developed by Eric Firing's group at the University of Hawaii (Firing et al., 1992). The velocity data were collected in 4 meter vertical bins with 5 minute ensembles using GPS for navigation and the Gyre's gyro compass for heading. Bottom tracking was used in waters of less than 100 meter depth and several long straight shallow water runs were used to establish the ADCP's angular offset and speed calibration. Navigation data were edited using a procedure that smoothed the ship's velocity relative to a reference layer velocity between 14 and 22 meters with a Blackman filter whose half-width was one-half hour. Because the ADCP was hull mounted at a depth of about 5 meters together with a 4 meter blank-aftertransmit, the first 4 meter bin was centered at about 11 meters. For this study a 10.5 hour segment of the ADCP current data was extracted covering the run up from the south to the approximate center of the cold eddy and from there to the northwest for the duration of the TOYO, approximately 5 hours, and then up onto the shelf. The velocity data were vertically and temporally averaged for presentation over ten meters and ten minutes, respectively. The velocity vectors for the layer between 10 and 20 meters have been over plotted on a portion of the May 16, 1993 morning AVHRR image from Fig. ld to show the correspondence between the velocity and the physical features made visible by the sea surface temperatures, Fig. 8. The velocity data were not detided because tidal constituents are unavailable for this area in any detail and the duration of the cruise was too short to establish the constituents based upon the ADCP data themselves. In any event, tidal velocities are not expected to be large in the 1000 to 2000 meter water depths in which the cold eddy was located. Shoreward of the 100 meter isobath, however, the Mz tidal current amplitudes are between 5 and 10 cm/sec based upon two month current records collected by SAIC for the Minerals Management Service during 1992 (data provided courtesy of Tom Berger). The near surface ADCP velocity data showed a clear indication of an anticyclonic flow pattern coincident with the cold surface expression of the cold eddy (D in Fig. 8). The westward currents immediately south of the eddy center and the northeastward currents along the northwestward cruise segment indicate that the TOYO section started just southeast of the center of the surface expression of the cold eddy. Maximum velocities in the l0 to 20 meter depth interval illustrated in the figure were about 25 cm/sec occurring 5 to 10 km southwest of the eddy center. Maximum near surface velocities on the TOYO transect were about 20 cm/sec ten kilometers northwest of the beginning of the transect. To the northwest of the eddy the current vectors rotated to the east in what appears to be a filament of warmer shelf water being drawn offshore (E in Fig. 8). Progressing south of the cold eddy the current vectors rotated toward the south and then east in another filament of warmer water drawn from the shelf (C in Fig. 8), although its relation to the cold eddy is unclear. There was also a band of warm water of Gulf Stream origin (B in Fig. 8) moving

1858

C.N. Flagget al.

onshore further south as part of the cyclonic Gulf Stream frontal eddy or "shingle" that propagated into the area from the southwest (see Fig. 1b and lc). South of this warm band was a long filament of "Ford" water (A in Fig. 8) although the ADCP velocities do not exhibit a strong cross-isobath flow, possibly as a result of tidal current interference. The 2-dimensional structure of the horizontal currents associated with the cold eddy was rather complicated (Fig. 10). A clear correlation between the currents and sea surface temperatures along the TOYO section (Fig. 8) diminished with depth although the anticyclonic sense of rotation was retained. In Fig. 10a (viewed looking southwestward), note l) that the northeastward velocity, upper panel, associated with the eddy extended down to about 250 m, 2) that there was a subsurface maximum at 17-22 km from the beginning of the section, and 3) that the center of the eddy, indicated by a change in sign of the northeastward current, appears to be near the 5 km mark. There was a region of relative current minimum between 22 and 25 km, extending from the surface to 400 m. Shoreward of this minimum there was a second northward current maximum down to 250 m flowing along the shelf break. Given the limited hydrographic evidence available (see Figs 3, 4 and 5), it seems that this second northeastward flow was a separate entity from the eddy and corresponds to filament E in Fig. 9. Below 300 m, the ADCP shows that the currents were to the southwest across the entire section, more southerly at the offshore end and more westerly near the shelf-break. The ADCP velocity section south of the eddy, Fig. 10b (viewed looking northwestward), reinforces the character and depth range of the eddy and suggests an origin for the large northeastward shelf-edge currents seen in Fig. 10a. The northwestward currents (lower panel) associated with the surface cold eddy (centered about the 70 km mark) extended down to about 250 m as in Fig. 10a, with nearly uniform magnitudes of between 25 and 30 cm/sec in a band that was 10 to 15 km wide. The center of the eddy, indicated by a change in sign of the northwestward current component, is visible at the extreme right of the section. Immediately south of the core of the eddy current was a southeastward flow of 20 cm/sec within the upper 50 meters which corresponds to the shelf water drawn off to the east just north of the Gulf Stream eddy (C in Fig. 8). South of the shelf water filament the current of the cyclonic Gulf Stream frontal eddy (B in Fig. 8) is clearly evident with onshore flow of 10 to 20 cm/sec that extends from near the surface to a depth of about 200 m. The maximum northeast current component (upper panel) associated with the Gulf Stream eddy was located just north of the maximum onshore flow. Together, these suggest that there may have been a bifurcation of the Gulf Stream eddy current with a significant portion flowing along the shelf break to the northeast. This then could be the source of the strong northeastward shelf-edge current in Fig. l 0a. It is interesting to note also that the penetration depth of the Gulf Stream frontal eddy was nearly the same as that of the cold eddy. 3. DISCUSSION As with all serendipitous investigations, the data collected in the vicinity of the cold eddy did not address all aspects of the eddy and its dynamics. There are however, a few issues that seem clear. The source water for the upper 50 m of the cold eddy water clearly was the cold pool and possibly the lower portion of the shelf thermocline. The SST images indicated that the cold water surfaced along a large section of the shelf edge extending at least as far north as 38°N. Historically, minimum water column temperatures less than 8°C only occur north of about 37.5°N at this time of the year (Houghton et al., 1982) while numerous sections

Cold anticycloniceddies formedfromcold pool waterin the southern MiddleAtlantic Bight

1859

including those from the SEEP-II program (Biscaye et al., 1994) indicate that waters less than 10°C are found all along the shelf north of Chesapeake Bay. Even though the cold pool waters are colder than slope waters of equivalent depth, they are less dense because of their lower salinity. As a result, in those same sections the isopycnals slope up and offshore where they either come to the surface or merge with the seasonal thermocline of the slope waters. The evidence thus suggests that the pathway of the cold eddy surface water was a general offshore and upward movement of sub-thermocline waters between 37°N to 38°N, approximately following the isopycnals. Figure 11 shows a temperature section from a TOYO transect on the shelf along the 40 meter isobath (see Fig. 2.) In it, one can see first the 8°C, then the 9°C isotherms intersect the bottom progressing from north to south. Thus, it appears that more and more of the cold pool waters were shunted offshore progressing from north to south. We surmise that the deflected cold pool waters then rose to the surface approximately along isopycnals which typically shoal in the offshore direction. From this evidence, the less than 7°C water mass, marking the coldest water of the cold pool and cold eddy, had departed from the shelf well north of the start of the alongshore TOYO section at 36.72°N. The satellite and ADCP data suggest that cold pool water emerged as a series of cold near surface anticyclonic eddies. The earlier SST images showed the cold water oriented along the outer shelf while later images indicated that the cold water had moved offshore all along the shelf by ~ 20 km. The eddies had a diameter of some 45 kilometers (e.g. eddy #3 in Fig. ld) and moved southward alongshore at a speed of about 20 cm/sec. As they moved south, the eddy diameter at the surface decreased and they became less distinct as they incorporated filaments of both shelf and slope waters. The eddies showed clear evidence of high biological productivity. The chlorophyllnormalized primary productivity (a'A) was higher in the center of the eddy than within the water sampled further inshore. Oxygen saturation was also higher within the eddy than further inshore, reflecting higher net production even though the chlorophyll concentration was lower. The bloom within the eddy was probably stimulated when nutrient-rich, cold pool waters became exposed to light as they surfaced. Both a*A and oxygen indicated that a phytoplankton bloom was occurring within the eddy. The higher chlorophyll concentrations, but lower normalized productivity and oxygen supersaturation, found within the shelf waters to the northwest reflect conditions of more balanced growth and

36.72"N 74.90°W

35.$0~ 75.03°W

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1860

C.N. Flagget

al.

lower net community production. These inshore waters, which appeared to be closer to a state of equilibrium between nutrient supply and primary production, are suggestive of a later stage of bloom development and therefore a longer residence time within the euphotic zone. Since the surface N + N and silicate data (Fig. 7) show that surface nutrients concentrations within the eddy had been depleted to near-background levels (i.e. not significantly elevated relative to the shelf waters) by the time of sampling, the eddy may have been present within the euphotic zone for a period sufficient to strip out nutrients. Nitrogen is invariably the limiting element for production in Middle Atlantic Bight waters at the sea surface. Thus, the presence of slightly higher phosphate but nearly identical N + N within waters surrounding the eddy (Fig. 7) likely reflects slightly higher preformed phosphate within the surface shelf water mass, relative to cold pool water. This, in turn, could be a consequence of localized riverine influence. There is evidence for some warming of the near surface water of about l°C to 2°C,as indicated by the spread of the T/S relationship at salinities of about 32.8 pss, Fig. 4. It would be possible to estimate the length of time since the cold pool water had surfaced if we knew the net air-sea heat transport during the previous couple of weeks. While we do not have sufficient data to compute the heat budget accurately, the Bunker climate atlas (Isemer and Hasse, 1987) indicates that the daily average net air-sea heat transport for the Mid-Atlantic Bight for May is about + 100 W/m 2. Over a l0 m thick layer, this translates into a temperature increase of about 1.5°C per week which is the same order of magnitude suggested by the T/S diagram. Thus, it seems likely that the cold eddies were about one week old by the time they reached the region of the Gulf Stream. At 20 cm/sec, the eddies would have traversed about 120 km during that period which is consistent with the series of satellite images. Another approach to estimating the lifetime of the cold eddies is to examine the oxygen distribution within the eddy to see whether it is consistent with the cold pool water surfacing and sustaining a phytoplankton bloom over a period of about a week. Pre-surfacing and pre-bloom nutrient (N + N) and oxygen concentrations for the water type found in the eddy were estimated from historical hydrographic and chemical data collected along the Middle Atlantic Bight between 1975 and 1996. The data were collected during the Atlantic Coastal Experiment, the SEEP II project and the Ocean Margins Program. We selected only stations from the months of April and May which sampled water with salinities between 33 and 33.4 and temperatures below 8°C. A total of 303 stations matched those criteria: all of these had nutrient data, however only 28 stations had oxygen data. Based on these stations we estimate the initial NO3 concentration for cold pool water found south of Delaware Bay, to be ~ 1 to 4/~mol/kg, with oxygen concentrations averaging ,-~310 #tool kg-1. The latter represents slight undersaturation of ~ 10-15 #mol/kg (cf. Sharp and Church, 1981). In contrast, the upper 30m of the central portion of the cold eddy had supersaturated oxygen values with depth-averaged oxygen concentrations of ~ 360 ~mol/kg. Thus, the integrated (0-30m) oxygen increase within the eddy since surfacing was ~1.5 m o l m "2. The photosynthetic quotient (moles of 02 produced/moles of carbon consumed) (Laws, 1991) typically ranges from 1 to 1.45 for net biological production. Thus, ignoring any loss of oxygen to the atmosphere, the oxygen increase implies a net organic carbon production of 1 to 1.5 mol (C) m "2. If this production took place over a period of one week as suggested by the heat budget, it would imply a net community production rate of 0.15-0.2 mol (C) m 2 day -1 or 1.8- 2.4 gC m -2 day -I . This estimate can be compared with independent estimates of gross primary production derived from the Fast Repetition Rate fluorometer.

Cold anticycloniceddiesformedfromcold pool waterin the southernMiddleAtlanticBight

1861

Note however, that the estimated oxygen inventory increase of 1.5 mol m -2 cannot be supported by the presumed initial standing stock of nitrate within the cold eddy waters, if classical algal stoichiometry is assumed. Using a classical Redfield ratio (-AO2/ANO3) of 8.6, removal of a nitrate concentration of 4/~mol kg n over the upper 30 m would support an average net oxygen change of 1 mol m -2. An alternative value of this ratio (Takahashi et al., 1985) of 10.75 gives an oxygen increase of 1.3 mol m -2. Because the initial nitrate inventory given above is probably an upper bound, the oxygen evolution appears to have outstripped in-situ nitrate utilization. Some additional nitrate may have been supplied from below via mixing. It is also possible that phytoplankton continued to photosynthesis after all nitrate had been removed, thereby producing oxygen and organic material with relatively high C:N ratios. We also estimate the total daily gross primary production of the water column using the average chlorophyll concentration and (tr*A) profiles from within the eddy together with a photosynthetically active light profile calculated from surface light values attenuated by water and chlorophyll. This estimation procedure relies upon a regression with measured J4C production over 6-hour shipboard incubations from stations occupied earlier in the cruise (see Kolber and Falkowski, 1993) and gives an estimate of 0.15 mol (C) m -2 day -I or 1.8 gC m -2 day -1. Net primary production is no greater than 80% of gross primary production (Falkowski and Raven, 1997), while net community production would be lower. Thus, either (1) the observed oxygen excess developed over a period greater than the one week suggested by the heat budget (e.g. 1 to 2 weeks), or (2) phytoplankton production at the time of measurement had declined from higher values which existed earlier in the lifetime of the eddy (e.g. prior to nutrient depletion). Overall, there is a reasonable correspondence between two independent estimates of eddy lifetime, given the uncertainties underlying both estimates, and suggests that the scenario of a bloom having occurred over a period of 1 to 2 weeks prior to the TOYO section is consistent with the biogeochemical evidence. The dynamics behind the apparently downwelled highly oxygenated and high chlorophyll water between 23 and 27 km from the beginning of the transect (Fig. 5) are unclear. There was a negative along transect gradient of the radial velocity component between 20 and 30 km (approximately -1.5 × 10-5 sec-1). If this approximated the divergence, that is there was no variation in the azimuthal direction, it would correspond to a vertical velocity difference over the upper 50 meters of -0.075 cm/sec, or -65 meters per day. The difficulty with assuming that the anomalously high oxygen and chlorophyll water was the result of downwelling is that the salinity and density values and T/S and O2/S relations are unlike any of the water found above or on either side of the eddy. We note that the temperature and especially the salinity of this water (9°C and 33.2 to 33.8 pss) are similar to those found within the shelf/slope front. In general, historical data show that high oxygen supersaturation and chlorophyll concentrations extend into higher density water within the front. It seems reasonable to assume, therefore, that this parcel came from a portion of the front which was offshore of the location where the cold pool water surfaced, and that it was drawn around the cold eddy as the latter propagated southward. There also may have been some subduction as this heavier water encountered the lighter, low salinity surface water that was being drawn directly off the shelf. We do not have much information about the deeper portion of the eddies except that provided by the ADCP. This data set clearly indicates that the vertical extent of the southern eddy was about 250 meters. Less precisely defined was the radius to the maximum currents which was 8 to 10 km on the southwest ADCP transect and about 18 km on the TOYO

1862

C . N . F l a g g et al.

transect, thus this southern eddy (eddy #2 in Fig. 1) at least, appeared to be distorted into an ellipse elongated in the offshore direction. The maximum velocities within the eddy were about 29 cm/sec along the southern transect and 17 cm/sec along the TOYO transect. Despite the lack of complimentary deep hydrographic data, we can speculate on the dynamical balances within the eddy by making use of these gross parameters. From the combined TOYO and ADCP data sets, we can estimate the potential vorticity of the upper 50 meters of the southern eddy using: r

~ (v,. + - +.)ON2,

(1)

r

where v is defined to be normal to the ADCP transects, r is the radial distance from the apparent eddy center, and N is the Brunt-Vaisalla frequency. In using this definition, we make the assumption that the ADCP transects were along approximate radials from the eddy center and that the currents normal to those radials were a reasonable representation of the eddy's azimuthal velocity. Calculations indicated that there was a constant potential vorticity region ( ~ 10.8 sec -3) below the weak thermocline within the eddy, Fig. 3a, that was contained laterally within the cold water portion (_<8°C) of the TOYO transect. These results lead then to the not unreasonable possibility that the entire eddy might be treated as a constant potential feature in the same manner that Flied (1979) used to model warm and cold rings generated by the Gulf Stream. The Flierl model for anticyclonic warm rings assumes a steady state, one and a half layer system with an axis-symmetric, constant potential vorticity eddy resulting in an interface that rises up from the eddy center to intersect the surface. The non-dimensional equations for the azimuthal velocity, v, and the eddy depth, h, are: 1 1 0 rOv_~v_E(r)

r2

= 0,

r Or

(2)

h = 1 + ~(v,. + v ) , r

where r is the radial distance, e = Vo/fR, where Vo is the maximum velocity which occurs at the maximum eddy radius, R is the deformation radius, R = (g'Ho)J/2/f, and Ho is the central eddy depth. The variables r, v, and h have been non-dimensionalized by R, Vo, and Ho, respectively. This set of equations forms an eigenvalue problem for e, an order one Rossby number, given the boundary conditions that:

h=0

v=0 (v=l)

at at

r=0,

(3)

r=ro,

where ro is the maximum non-dimensional eddy radius (ro = Ro/R) where v is set to one to preserve the definition of Vo. The eigenvalue problem was solved numerically using a fourth order Runge-Kutta method. When Flierl applied the model to warm rings he had measurements that included all the necessary parameters: ring diameter, central depth, maximum velocity, and, in particular, a value for reduced gravity, g'. While we also had estimates for the first three parameters, we did not have an estimate of g' with which to directly scale the problem. So instead, we solved the equations for a range of ro'S and obtained a functional relation between ro and the ratio:

Cold anticyclonic eddies formed from cold pool water in the southern Middle Atlantic Bight

1863

rO .fRo --

- -

e

(4)

Vo'

where Ro is the maximum dimensional radius of the eddy. Using this relation to estimate ro from Ro and Vo, we could solve the eigenvalue problem for e, the eddy velocity and depth profiles, and by inverting the definition of the Rossby number, determine the value for the reduced gravity:

g' = (V°)ZlHo.

(5)

6

A radius of 18km, a maximum speed of 17 cm/sec, and a central depth of 250m for the TOYO section yielded an eigenvalue of-0.98, a reduced gravity of 0.012 cm/secz, and a depth profile that has been over-plotted on the northeastward velocity section in Fig. 10a. In the parameter range that seems to typify these eddies, significant velocities are confined to the outer portion of the eddy and the azimuthal velocity profile looks very similar to the observed profile. The depth profile in Fig. 10a where the center of the eddy was assumed to be 4 km from the start of the transect, appears to match the observed velocity distribution rather well lending some confidence in the application of the model and the reduced gravity estimate that resulted. The reduced gravity suggested by the model, 0.012 cm/sec2, applies to the depth averaged density difference between the eddy and the surrounding waters. The value is quite small, a factor of twenty less than a typical warm ring might have, and is equivalent to a sigma-t difference of similar numerical size. Unfortunately, this small difference does not allow us to distinguish the deep source waters of the eddies because the temperature/salinity relationships are quite variable above about 300 meters in this area.

4. CONCLUSIONS From the size and movement of the eddies and the vertical extent of the cold waters we can estimate the volume transport of cold shelf water associated with the eddies. The diameter of the southern eddy (Eddy #2) was about 34 km, the depth over which salinities were lower than 34 pss was about 60 m, thus, the volume of shelf water within the eddy was about 54 km 3. An estimate of cold shelf water volume for the well defined central eddy (Eddy #3) whose diameter was about 45 km, Fig. lb, assuming a similar vertical structure, is about 95 km 3. The interval between the cold eddies in Fig. 1 was about 100 km, or more than twice the estimated diameter of the eddies. At 20 cm/sec, an eddy would travel this distance in about 5.8 days which leads to transport estimates of between 0.1 Sv and 0.2 Sv. These results are similar in size to the annual mean alongshelf transport estimate of shelf water of 0.2 Sv from the SEEP-II cross-isobath array located just south of 38°N (Biscaye et al., 1994) and suggest a near complete diversion of the alongshore shelf flow. Since it appears that the water that escaped the shelf consisted primarily of cold pool water, with the exception of the surface filaments at the southern end of the Bight, most of the upper layer waters must be accumulating within this region. This is in qualitative agreement with the diversion of subthermocline water shown in the alongshore TOYO results from Fig. 11 although it is not evident from the three cross-isobath hydrographic sections.

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The mechanisms and rates of exchange of shelf water with the open ocean are important for ocean biogeochemistry, particularly the carbon cycle (e.g. Walsh, 1983; Falkowski et al., 1988; Biscaye et al., 1994). Particularly important is the fate of sub-euphotic zone shelf water masses, of which the Middle Atlantic Bight cold pool is a good example. These deep shelf waters act as a seasonal focal point for nutrients, with the sub-euphotic zone portions tending to retain their initial, late-winter, nutrient (and CO2) content in a similar fashion to their retention of their temperature and other dissolved gas signatures (e.g. Wallace, 1994). In an Eulerian sense, nutrients increase in the southern Middle Atlantic Bight cold pool during Summer due to the southwestward flow of shelf water. This flow, when coupled with a seasonal reduction in vertical mixing associated with a strong summer pycnocline, allows the higher nutrient levels found in the northern portions of the Middle Atlantic Bight to penetrate southward. This is analogous to the well-known decrease of cold pool temperatures during summer (described in the Introduction). Finally, nutrients and CO2 tend to become focussed within the cold pool during Spring and Summer, with accompanying reductions in dissolved oxygen (Sharp and Church, 1981; Kemp et al., 1994), as a result of the sinking of organic material from the overlying euphoric zone followed by interception and remineralization at the sediment-water interface. Lillibridge et al. (1990) and Wood et al. (1996) have demonstrated that high-chlorophyll, cold pool waters found over the inner continental shelf north of Cape Hatteras can be advected and mixed into adjacent Gulf Stream waters in the immediate vicinity of Cape Hatteras. The shelf water and the organic carbon exported via this process can be traced as low temperature and low salinity intrusions with relatively high chlorophyll content within the upper 200m of the Gulf Stream (see also Ford et al., 1952). Wood et al. (1996) argued that the organic material exported via this process, which is mixed into low-nutrient Gulf Stream waters, is likely to be oxidized in-situ and is unlikely to represent an immediate sink for organic carbon to the deeper layers of the ocean and slope sediments. Similarly, Lillibridge et al. (1990) pointed out that once entrained into the Gulf Stream, the shelfderived carbon can be exported far into the open ocean on a time scale of a few days. According to this scenario, remineralized nutrients would become available for further cycles of production and export along the axis of flow of the Gulf Stream. However the implication of this scenario is that any sinking of organic carbon associated with the export of organic material which was initially fixed on the shelf and sequestered within the cold pool, would be non-localized and distributed over a wide region. It is worth noting that the cold pool water sampled by this study, within eddies located beyond the shelf break (core salinities of 32.8 to 33.4), was significantly more saline than cold pool water (salinities of 32.4 to 32.9) sampled on the shelf by Wood et al. (1996). Because salinity increases with depth over the shelf, the cold pool water found within the offshore eddies must have originated at greater depths over the continental shelf further to the north than the water sampled over the shelf by Wood et aL. The high biological productivity within an eddy of cold pool water located over the continental slope described here has striking similarities with a scenario described by Marra et al. (1990) based on a TOYO study from the northern end of the Middle Atlantic Bight (3941 N; 71 W). The Marra et al. (1990) study included high-resolution nutrient data and showed that the chlorophyll was concentrated within a nutricline. Marra et al. (1990) hypothesized that migration of high nutrient shelf waters offshore along upward-sloping isopycnals into the euphotic zone provides a mechanism whereby nutrients (or a "production capacity") derived from the shelf, can be fixed over deep water with

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consequent export of organic carbon to deep water. Our study suggests that this may be an important process all along the Middle Atlantic Bight. As Marra et al. (1990) noted, the importance of this process is that it provides a mechanism to export shelf production to deep water without exporting chlorophyll or particulate organic material directly across isobaths. Export of nutrients would not have been detected, for example, by the fluorometers deployed at the shelf edge by Wirick (1994). Similarly, even if a large proportion of shelf production is oxidized on the shelf (Falkowski et al., 1988), rather than being exported as organic carbon as had earlier been hypothesized (Walsh, 1983), the oxidation products which become focussed within the cold pool, could be exported and subsequently repackaged for sinking into the deep ocean via this process. The question of the formation of the anti-cyclonic eddies remains to be answered satisfactorily. Initially, based upon the satellite and TOYO data the supposition was that cold pool water had been forced offshore, rising along converging isopycnals until the water reached to surface. The vertical squeezing of the cold water mass would have produced an anticyclonic relative vorticity generating eddies potentially similar to those observed. The difficulty with this scenario is the evidence from the ADCP which clearly shows that the anticyclonic motion extended much deeper, ca. 250m, than it is plausible for shelf water to penetrate and with more kinetic energy than obtainable simply from the migrating shelf water. This whole scenario is remarkably similar to the situation found by Houghton et al. (1986) (the physical companion paper to Marra et al., 1990) in an anti-cyclonic filament that they tracked off the shelf south of New England. In that case, shelf water moved up and offshore along isopycnals and either produced or was the result of, a deep anti-cyclonic velocity feature that ADCP data showed to be about 200 m deep. The overall size of their eddy-like feature from AVHRR photos was also similar although the high current regime that they measured had a much smaller radius of curvature (~ 7 km) than seems appropriate in our case. While Houghton et al. attributed the eddy spin-up to local geostrophic adjustment and vertical pressure work, they were unable to suggest a sufficiently large energy source. If on the other hand, we do not suppose that the eddies are formed from the offshore movement of cold pool water itself, we must assume a pre-existing eddy system within the slope waters just offshore of the shelf break. Thus, after the cold pool water was driven offshore by whatever means, it surfaced on top of a series of anticyclonic eddies and picked up its rotation from the underlying water masses. Presumably, the underlying eddies had two-dimensional characteristics similar to that evidenced by the surface expression of the cold eddies, that is diameters of 30 to 45 km, separation distances of 100km, and alongshore propagation speeds of about 20 cm/sec. The 250m depth of the eddies is about half that found within warm core rings about which much has been written (eg. Saunders~ 1971: Morgan and Bishop, 1977; Joyce et al., 1984; Auer, 1987). The typical warm core ring also has a diameter of about 100kin, azimuthal speeds of 100 cm/sec, a reduced gravity of about 0.2 cm/sec 2 and alongshore propagation speeds of 6 cm/sec. Lastly, six warm rings per year in the western slope would be a major ring year. These warm ring characteristics all differ markedly from the apparent characteristics of the deep portion of the cold eddies. It seems implausible therefore that warm rings, even moribund warm rings, would be the source for the deep portion of the cold eddies. Another possibility is that the deep eddy water was formed from intermittent pulses of Gulf Stream frontal eddy water such as appeared to be flowing northeastward along the shelf edge in Fig. 3a. It is difficult to reconcile this mechanism with the appearance of anti-cyclonic eddies south of New England, some

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800 km to the northeast. It may be that they are two separate processes, one originating from the Gulf Stream, the other from warm rings, and that for some reason stimulate features of similar horizontal and vertical scale. The actual formation mechanism of the eddies, how they form themselves into apparently constant vorticity eddies, and why there was so little apparent interaction between the deep eddies and their source waters remain to be investigated. Acknou,ledgements--The collaboration and assistance of Paul Kemp in the collection of the data is greatly appreciated. Neither the data collection nor its reduction would have been possible without the able assistance of Bill Behrens, Yan Shi, Rick Wilke and Steve Boose. We thank Dave Chipman for his help during the cruise and for sharing his data with us. We thank Claire Lamberti for her editorial assistance. This manuscript has been authored under Contract No. DE-AC02-76CH00016 with the U.S. Department of Energy. Accordingly, the U.S. Government retains a nonexclusive, royalty-free license to publish or reproduce the published form of this contribution, or allow others to do so, for U.S. Government purposes.

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