The advection speed of the GSA'80s seems to be greater than the one of the GSA'70s: The 1980s anomaly reached the Barents Sea 6 to 7 years after peaking in the West Greenland Current, while the 1970s anomaly traveled the same route in 8 to 10 years. These anomalies, however, seem to be of different origin. The GSA'70s was apparently boosted remotely, by a freshwater/sea ice pulse from the Arctic via Fram Strait. Consequently, the GSA'70s was accompanied by a large sea ice extent anomaly in the Greenland and Iceland Seas, which propagated into the Labrador Sea. In contrast, the GSA'80s was likely formed locally, in the Labrador Sea/Baffin Bay mainly because of the extremely severe winters of the early 1980s, but supplemented with a possible contribution of the Arctic freshwater outflow via the Canadian Archipelago (facilitated by strong northerly winds) which would have enhanced stability and ice formation. This anomaly was also associated with a positive sea ice extent anomaly in the Labrador Sea/Baffin Bay which, however, had no upstream precursor in the Greenland Sea. Thus the GSAs are not necessarily caused solely by an increased export of freshwater and sea ice from the Arctic via Fram Strait. These results are corroborated by the early 1990s data when a new fresh, cold anomaly was formed in the Labrador Sea and accompanied by a large positive sea ice extent anomaly. The harsh winters of the early 1990s were, however, confined to the Labrador Sea/Baffin Bay area while the atmospheric and oceanic conditions in the Greenland, Iceland, and Irminger Seas were normal. The Labrador Sea/Baffin Bay area appears therefore to play a key role in formation of GSAs as well as in propagation of the GSAs formed upstream. A likely contribution of the enhanced Canadian Archipelago freshwater outflow to the GSA formation also seems to be significant. Two major modes of the GSA origin are thus identified, remote (generated by an enhanced Arctic Ocean freshwater export via either Fram Strait or the Canadian Archipelago) and local (resulting from severe winters in the Labrador Sea/Baffin Bay).
See Dickson et al. (1988) and Belkin et al. (1998).
The Greenland-Scotland Ridge extends from East Greenland to Scotland and below a depth of 840 m it forms a continuous barrier between the North Atlantic and the ocean regions north of the ridge. At higher levels, Iceland and the Faroe Islands divide the ridge into three gaps which have different widths and sill depths.
From northwest to southeast the first gap is the fairly wide Denmark Strait with a sill depth of about 620 m. Between Iceland and the Faroe Islands is the Iceland-Faroe Ridge, a broad ridge with minimum depths along the crest of 300-500 m, generally deepening from the Icelandic to the Faroese end. The deepest passages across the Iceland-Faroe Ridge are in the form of four channels, with sill depths between 420 m close to Iceland and 480 m close to the Faroes.
Between the Faroes and Scotland the bottom topography is more complex. The relatively broad, deep Faroe-Shetland Channel is blocked at is southwestern end by the Wyville-Thomson Ridge with sill depth around 600 m. The Wyville-Thomson Ridge joins the Scottish shelf at its southern end and at the northern end joins the Faroe Bank rather than the Faroe Plateau and these two are separated by the narrow, deep Faroe Bank Channel with sill depth around 840 m. This channel, which is a continuation of the Faroe-Shetland Channel, thus exceeds all other passages across the Greenland-Scotland Ridge by more than 200 m in sill depth.
The Ridge separates the basins of the Nordic Seas to the north from the basins of the North Atlantic Ocean to the south. The latter are, from west to east, the Irminger Basin, the Iceland Basin and the Rockall Channel or Trough. See Hansen and Osterhus (2000).
In the summer, the volume of the Greenland Sea consists of about 85% of the deep and bottom water masses (i.e. Greenland Sea Deep Water (GSDW) and Norwegian Sea Deep Water (NSDW)), 9% Arctic Intermediate Water (AIW), and 9% surface water masses, mostly Atlantic Water (AW). See Swift (1986) and Hopkins (1991).
he Gulf of Aden is influenced at depth by the outflow of Red Sea waters moving toward the Indian Ocean, and at the surface by inflow from the Arabian Sea. The signature of the Red Sea outflow is seen throughout the Gulf of Aden and northern Indian Ocean as an intermediate salinity maximum near 600 m depth, spreading southward along the western boundary as far as 20 S. As the Red Sea water spills over the Bab el Mandeb sill, it appears to follow at least two pathways in the western Gulf of Aden, one along the southern boundary of the Gulf in the expected sense, and another along the central or northern part of the Gulf (Federov and Meschanov (1988)). These different pathways appear to be related in part to the complicated topography of the western Gulf, including the Tadjura Rift that extends westward to just outside the Bab el Mandeb. Different mixing behavior along these flow pathways may lead to different penetration depths between 400-1200 m and varying properties of the Red Sea water in the Gulf of Aden. Little detailed knowledge is available on the eastward spreading of Red Sea water in the central Gulf of Aden or how this takes place, either in the form of continuous boundary currents or isolated eddies. The pathways by which surface waters navigate their way westward through the Gulf to provide the required surface layer inflow to the Red Sea are also poorly known.
The upper layer circulation of the Gulf of Aden appears in remotely sensed SST imagery and a few available AXBT survey data to contain large eddies - mostly anticyclones - that are comparable in size to the width of the Gulf (Fig. 3d). These features appear to propagate westward from the mouth of the Gulf toward the Red Sea, and their origin may be linked to the propagation and decay of eddy features generated in the western Arabian Sea. In addition, the seasonally reversing winds over the Gulf may generate localized responses consisting of gyres and seasonal boundary currents along the northern and southern boundaries of the Gulf. Direct evidence for these gyres, eddies, or seasonal boundary currents from in situ observations is almost entirely lacking, however.
Seasonal upwelling with the onset of the SW monsoon is quite pronounced in SST imagery in the western Gulf of Aden. Cool upwelled waters are brought to the surface along the southern coast of Yemen beginning in May and are presumably advected eastward by a wind driven coastal current. The lifting of the thermocline and depression of the sea surface in the western part of the Gulf caused by this seasonal upwelling process is believed to play a major role in the reversal of the surface flow in the Bab el Mandeb Strait in summer, and the associated intrusion of Gulf of Aden thermocline water into the Red Sea. A front is often observed near the mouth of the Gulf (Fig. 3d) during the SW monsoon which marks a water mass boundary between the cool upwelled waters advected northward along the Somali coast and the warmer waters in the Gulf of Aden. Very little is known of the 2-dimensional circulation of the Gulf of Aden or its causes.See Federov and Meschanov (1988).
... the circulation around the Alaska Gyre consists of the eastward flowing Subarctic Current at about 50N, the Alaska Current in the northern Gulf of Alaska, and the southwestward flowing Alaska Stream along the Alaskan Peninsula. Some of the water from the Alaska Stream recirculates into the Subarctic Current, but the strength and location of the recirculation, though poorly described, appear extremely variable. The northward flow of the broad, diffuse, eastern boundary current along the west coast of North America at about 50N is considered to be the origin of the Alaska Current. At the head of the Gulf of Alaska, this flow converges into the swift, narrow Alaska Stream, which has characteristics of a western boundary current. The easternmost extent of the Alaska Stream cannot be rigorously defined, but common nomenclature refers to the extension of the Alaska Current between 150 and 180W as the Alaska Stream.See Favorite et al. (1976), Royer and Emery (1987) and Musgrave et al. (1992).
... as an alternation between a simple two-layer upwelling pattern and more complex basin modes. During steady southerly winds, water flows north out of the bay at 0.05-0.1 m s in a surface layer that is deeper (15 m) on the western side of the bay and intersects the surface on the eastern side, due to upwelling on the east and downwelling on the west (the bay is somewhat larger than the local internal deformation radius). Weak (0.05 m s) return flow occurs beneath the surface layer. The raised layer interface on the eastern side tends to rotate cyclonically to the southeast corner, where it is `arrested' by strong wind-driven vertical mixing, creating the colder temperatures found at the southeast of the bay during strong upwelling. When winds relax or reverse, the raised interface continues to rotate cyclonically to the western side of the basin, where observations of stronger currents, vertical shear and low Richardson numbers indicate intense vertical mixing as the basin mode dissipates. The resumption of southerly winds reestablishes the initial pattern. Enhanced mixing processes during both relaxations and upwelling may help maintain the high primary productivity within the Gulf of Arauco. Furthermore, the geostrophic circulation tends to follow the bathymetry along the shelf outside the Gulf, which may help maintain the high primary productivity rates by confining the waters to stay within the gulf, on the inshore side of the shelf.See Strub et al. (1998).
The Gulf of Carpenteria is a shallow (maximum depth about 70 m) semienclosed body of water. Most of the exchange with the open oceans occurs through its western entraces, as Torres Strait is extremely shallow. In winter, temperatures are at their minimum and salinities at their maximum, and vertically well-mixed conditions predominate. In the austral summer monsoon season, tidal currents are only sufficiently strong to mix the bottom 30 m of the water column, and as a result, the central Gulf is stratified, principally through the input of surface heat but also from lower surface salinities due to rainfall and runoff. In the shallower water near the coast, well-mixed conditions prevail, but there is considerable influence of monsoonal river runoff. It appears that significant changes in gulf water properties are the result of local processes rather than exchange with the surrounding Arafura or Banda Sea waters. A channel model indicates there is little seasonal transport through Torres Strait.
The northwest monsoon winds, density-induced currents and nonlinear tidal rectification all result in a clockwise circulation in the gulf. However, the southeast trades drive a counterclockwise circulation. There is a coastal boundary layer that does not mix rapidly with the central gulf waters, and coastal jets forced by the wind are confined to this boundary layer. The existence of such a layer explains the persistence of the low-salinity regions observed near the coast.See Forbes and Church (1983), Wolanski et al. (1988) and Rothlisberg et al. (1989).
The large-scale water mass distribution in the Gulf of Mexico reflects the limited exchange the gulf basin has with the adjacent oceans. In general, the gulf waters consist of three distinct water masses: subtropical underwater, antarctic intermediate water, and North Atlantic deep water. The subtropical underwater enters the gulf from the Caribbean at depths of 200 to 500 m and is found throughout the eastern portion of the gulf. This water is readily recognized by its high salinity, 37.00 ppt. Antarctic intermediate water also enters the gulf through the Yucatan Strait and is found throughout the gulf between depths of 500 to 1,200 m (in the eastern gulf) and 600 to 800 m (in the western gulf). This water mass is recognized by a distinct minimum, 34.00 ppt, in salinity. North Atlantic deep water is found below 1,200 to 1,400 m throughout the Gulf of Mexico. McLellan and Nowline (1963) suggested that waters deeper than 1,500 m in the Gulf of Mexico have long residence times (300-500 years) are are not frequently exchanged with outside waters. Hydrographic observations indicate that additional water masses - gulf water, for example - are formed locally in the Gulf of Mexico during periods of intense winter cooling.
Numerous studies have shown that the general large-scale circulation in the upper 1,400 m of the Gulf of Mexico is anticyclonic (clockwise). The transport in the northern limb of the anticyclonic gyre is a combination of flow from the Texas shelf and from the southern portion of the gyre. The contribution from the Texas shelf can at times be as high as one-third of the total transport of the easterly flow in this limb of the gyre. The westerly flow in the southern part of the anticyclonic gyre is composed predominantly of water recirculating in the southern gulf, although at times water separating from the Loop Current can contribute to this transport. Average geostrophic velocities and volume transports associated with the large-scale anticyclonic circulation of the Gulf of Mexico are 10 cm/s and 5 . Additionally, large-scale cyclonic (counterclockwise) circulation gyres are found in the Bay of Campeche and over the northern portion of the wester Florida shelf.
Superimposed upon the large-scale circulation of the gulf are two major circulation features, the Loop Current and Loop Current rings. Both of these have considerable influence on the circulation characteristics of the Gulf of Mexico.
The Loop Current is a swift, narrow current that enters the Gulf of Mexico through the Yucatan Strait. This current can be traced as a coherent feature that extends into the northern portion of the eastern gulf, where it turns to the easter and then flows southward along the west Florida shelf. At the southern extent of the Florida shelf, the Loop Current again turns east and exists the Gulf of Mexico through the Straits of Florida. The Loop Current is part of a large circulation system that feeds into the Gulf Stream along the eastern boundary of the United States.
The Loop Current can be readily distinguished in vertical density distributions down to depths of 1,000 to 1,200 m in the region where it enters the gulf. Surface geostrophic velocities into the gulf associated with the Loop Current have been estimated to be 100 to 150 cm/s and the corresponding volume transport has been estimated to be 25 to 35 /s. Surface velocities diminish somewhat as the Loop Current extends into the gulf and widens. Outflow surface velocities are on the order of 50 to 100 cm/s and the corresponding volume transport about the same as into the gulf.
The warm core (anticyclonic) rings that separate from the Loop Current are a major circulation feature. Observations show that these rings typically separate from the Loop Current at the time of the maximum northward penetration of this current into the gulf. On average, one to three rings per year may separate from the current.
Rings are approximately 300 to 400 km in diameter and have a depth signature that extends to approximately 1,000 m. After detaching from the Loop Current, the rings move westward across the gulf, with observations showing them to exist as identifiable features for periods of several months. Geostrophic surface velocities have been estimated to be on the order of 25 to 100 cm/s, with associated volume transports on the order of 5 to 10 /s. These rings therefore represent a major mechanism by which properties such as temperature and salinity are transported from the eastern to the western gulf. One in the western gulf, they encounter the Texas or Mexican continental shelf. The fate of the rings at this time is not fully understood.
On the Texas-Louisiana continental shelf, west of 92.5W, the predominant feature of the circulation is a cyclonic (counterclockwise) gyre, elongated in the alongshelf direction. The inshore portion of this gyre is directed westward (downcoast). An eastward flowing countercurrent at the shelf break constitutes the outer portion of the gyre. Flow in the western extent of the gyre is directed offshore, while that in the eastern gyre - near Louisiana - is directed onshore. The alongshore wind stress is the primary mechanism driving the circulation of this cyclonic gyre. Thus the gyre exhibits seasonal variability in strength and occurrence that reflects the seasonal variability in the wind patterns. In July, when the downcoast (to the west) wind stress is diminished, the cyclonic gyre on the shelf disappears and is replaced by an anticyclonic gyre centered off Louisiana. In August and September, the prevailing wind direction changes abruptly and the gyre is re-established.
The northern Gulf of Oman is strongly influenced by outflow from the Arabian Gulf. From fall through mid-spring, satellite SST's suggest a plume of Gulf water flowing as a coastal current along the Oman and Emirate coast to Ras al Hadd at the edge of the Arabian Sea. This would imply that at least through part of the year the outflow from the Gulf consists of a deep water (PGW) layer and a modified surface layer that must together balance the inflow component. This and the absence of any sill to confine the flow differentiates the Gulf from marginal seas such as the Red Sea and Mediterranean. The manner in which the PGW enters the deep Gulf of Oman is not clear from available data. Data from the U.S. and German WOCE cruises in 1995 and Navoceano AXBT surveys suggest that the PGW layer is dominated by sub-mesoscale eddies. Are these formed at the outfall of the paleo-river channel at the shelf edge or by shelf edge meandering as the plume proceeds southeastward down the shelf break? What are the range of sizes and dynamics of the resulting Peddies? The final issue is the nature of the PGW and associated surface flows. Is the flow a coherent shelf break one or a train of eddies? The interaction of these with the other elements of the Gulf of Oman circulation is also of interest.
Other important elements involved in the Gulf of Oman's circulation are the seasonal upwelling along the coast of Iran to the north and the complicated mesoscale dynamics associated with the extension of the south coastal Oman upwelling system into filaments extending off Ras al Hadd. The latter is complicated by the shallow Murray Ridge that extends across the mouth of the Gulf. The Ras al Hadd jet is highly variable, sometimes extending out to the east as shown in the figure and extending northeastward or southeastward at other times. This feature is also referred to as the Ras al Hadd front because it forms the seasonal boundary between the northern Arabian Sea and the Gulf of Oman. During the SW monsoon the transport of the Ras al Hadd jet is believed to be at least 10 Sv (Elliot and Savidge (1990)). Flagg and Kim (1998) discovered that the Ras al Hadd jet intensified in August 1995 following the reversal of the flow along the northeastern Oman coast from northward to southward, thereby adding to the flow along the Ras al Hadd front. It is speculated that the reversal of the flow along the northeastern Oman coast in August is related to the intensification and/or propagation of a cyclonic eddy in the Gulf of Oman during this period. Similarly, it has been suggested that such an eddy can play a role in the dynamics of the Ras al Hadd Jet, which may become tied to a double vortex as it extends offshore. It is speculated that to the south, an anticyclonic eddy forms, while to the north in the Gulf of Oman, a cyclonic eddy forms, both of which are driven by the extension of the Ras al Hadd jet into the open Arabian Sea. While the anticyclonic eddy to the south has been observed, no direct connection has yet been established between the Ras al Hadd jet and the cyclonic eddy in the Gulf of Oman. The interaction of the Ras al Hadd front with the coastal flow, the Murray Ridge and the eddies are not well understood. In some of the remote SST data, shifts in the dipole lead to its breakup and the propagation of the cyclonic component northwards into the Gulf of Oman. The interaction of these surface intensified features with the thermocline layer Peddies and the PGW outflow is probably complicated. The interannual variations are large, as are those in the interactions with the Murray Ridge and upwelling on the Iranian coast.
The seasonal and interannual variations in the circulation in the Gulf of Oman appear to be significant. The nature of the flow along the southern side of the Gulf appears to be better organized in June through December although this may be tied to the lower thermal contrast in January through May. The northern side of the Gulf has consistent upwelling associated with the SW monsoon along the Pakistani coast. Upwelling along the western, Iran coast is more variable. In 1995, for example, this coast was associated with upwelling filaments that moved to the west and even entered the outer edges of the Strait of Hormuz. Other years suggest less extensive upwelling although there is localized upwelling at the mouth of the Strait in all years examined. One clear need is a better depiction of winds over the Gulf of Oman in relationship to this variability.See Elliot and Savidge (1990) and Flagg and Kim (1998).
The freshwater input causes the entire gulf to be stratified in salinity in the top 20 m. This halocline inhibits tidal mixing, even in shallow coastal areas where tidal currents are greater than 1 m s. The dominant forcing of the circulation is the eastward flowing Coral Sea Coastal Current in the Northwest Coral Sea. It appears to generate a counter-clockwise rotating eddy in the Gulf. Wind forcing is a secondary factor, causing the brackish water to leave the Gulf alternatively at its western and eastern sides. See Wolanski et al. (1995).
As with the rest of the Baltic, winds are most frequently southwestern, giving a general tendency towards cyclonic circulation. In the Gulf the water from the Baltic Sea Proper enters along the southern side of the Irbe Strait and flows around the Gulf. The large amount of river water entering from the south flow along the eastern coast towards the north. In summer months, this circulation may be reversed, with riverine and nutrient-rich intermediate waters transported along the western coast into the nutrient-depleted surface layers of the Irbe Strait, where the outflow is along the northern side. In the Suur Strait to the north, the flow is more or less unidirectional, but changes direction frequently. The currents can be large compared to elsewhere in the Gulf, allowing the transport through the Suur to be comparable to that of the larger Irbe Strait. The transport in the Suur is driven mostly by wind forcing, whereas horizontal density and surface gradients drive the transport through the Irbe.
This circulation pattern explains the observed long-term stable salinity difference between the Baltic Proper and the Gulf. The convergence of in- and out-flowing current support the persistent Irbe Front in the strait area. The slow and steady decrease in the deep water salinity in the Baltic Proper between 1977-1991 is reflected in the deep water of the Gulf. There is a large annual temperature cycle, with autumn cooling and spring warming overturning most of the water column.
Current measurements in the Suur Strait indicate the simultaneous coexistence of several flow regimes. There is a slow regime with surface outflow of gulf water along the northern part of the strait and a deep inflow of Baltic Proper surface water along the southern part. These are separated by a salinity front, i.e. the Irbe Front. Superimposed on this slow regime are high-frequency, unidirectional currents driven by sea level fluctuations. The most energetic movements diurnal and low-frequency oscillations, the the diurnal oscillation part of the eigen-oscillations of the Baltic Sea, Irbe Strait and Gulf of Riga system.