The East Greenland Current (EGC) flows southward along the eastern coast of Greenland from Fram Strait (79°N) to Cape Farewell (60°N) via the Greenland Sea, the Norwegian Sea, and the Denmark Strait (Woodgate et al. 1999). Its low-density water and the conservation of potential vorticity cause the EGC to remain geostrophically constrained to the Greenland Continental Margin (Hopkins 1991). At the Jan Mayen Fracture Zone, the upper part of the EGC is deflected toward the east. Denser water found below 1600 m cannot cross the ridge into the Iceland Sea; therefore, it circulates around the Greenland Sea Gyre, making its way toward the center as it interacts with the Greenland Sea waters (Rudels et al. 1999). Hopkins (1991) analyzed water masses of the EGC and found that it had eastward extensions. The dynamical mechanism behind them is unknown, but it is suspected that they arise from an interaction between the bathymetry and the barotropic component of the EGC. Easterly winds can prevent the formation of the eastward extensions to The East Greenland Current.
The East Greeland current as represented by the Mariano Global Surface Velocity Analysis (MGSVA). The East Greeland Current transports very cold, low-salinity water and Arctic icebergs southward. It is a source water for the West Greeland Current and Labrador Current. Click here for example plots of seasonal averages.
As the only major southward flowing current in the Greenland Sea, the EGC transports recirculating Atlantic Water, Arctic Ocean water masses, and >90% of the ice exported from the Arctic Ocean (Woodgate et al. 1999, Rudels et al. 1999). The EGC is thus an important link between the Arctic Ocean and the North Atlantic Ocean. With its high surface current velocities, the EGC carries sea ice and Polar Water out of the Arctic Ocean through Fram Strait, acting as the main freshwater sink for the Arctic Ocean (Schlichtholz and Houssais 1999). In addition to the cold, low-salinity surface water from the Greenland Sea, the EGC is also fed by warm and saline water from the south via the Norwegian Atlantic Current. This suppresses ice formation in the current because the addition of saline water to this region of intense cooling destabilizes the water column (Aagaard et al. 1985).
According to Aagaard and Coachman (1968a), three major water masses can be found in the EGC. In the upper 150 m there is Polar Water with temperatures between 0°C and the freezing point. This layer has a strong halocline, with salinity of 30 psu or less at the surface and 34 psu or more at 150 m. Under the Polar Water is the Atlantic Intermediate Water, which extends to approximately 800 m. Its temperature remains above 0°C, while its salinity increases with depth from about 34 psu to a value between 34.88 and 35 psu. It usually reaches this maximum value at about 400 m, a depth below which the salinity remains fairly constant. The final water mass, the Deep Water, is found below 800 m. Its temperature is less than 0°C, and its salinity is between 34.87 and 34.95 psu. The Polar Water of the EGC originates in the Arctic Ocean, while the deep water masses circulate cyclonically (Aagaard and Coachman 1968b).
There have been several attempts to measure the current speed of the EGC. Aagaard and Coachman (1968a) found that the EGC increases in speed from 4 to 14 cm s-1 as it flows south. However, they cautioned that at least part of this observed increase could be caused by preferential sampling, since observations in the south were taken mostly from the eastern edge of the current, which is faster than water further > inshore. Typical current speeds were 10 to 15 cm s-1, and there was no decrease in speed with increasing depth. Then Foldvik et al. (1988) found that in the upper 500 m of the EGC at 79°N, the one-year mean speeds of the current were as high as 9.5 cm s-1. Muench et al. (1992) measured the speed of the EGC in the upper layer (at 35 m and 150 m) and found an average speed of 8 cm s-1. Finally, one of the latest estimates comes from Bersch (1995), who calculated geostrophic velocities from CTD measurements and referenced them to velocities recorded in the upper 500 m and found maximum speeds of 20-30 cm s-1. According to Hopkins (1991), transport values for the EGC have varied from 2 Sv to 32 Sv. This discrepancy is primarily attributed to the different speeds assumed for the underlying Return Atlantic Intermediate Water. Aagaard and Coachman (1968a) defined the EGC as the flow in the Greenland Sea occurring west of the 0° surface isotherm. They computed the transport based on the mass and heat budgets and obtained an annual mean of 35 Sv. They suggest that the reason their value was larger than previous estimates (which were an order of magnitude smaller) was that the current varies interannually. Foldvik et al. (1988) used year-long moored measurements from the EGC near 79°N to calculate its transport, and found that in the upper 700 m it was 3 Sv, half of which was barotropic. Hydrographic data collected from the EGC in the summer allowed Schlichtholz and Houssais (1999) to estimate a southward geostrophic transport relative to the bottom of 4 Sv at 78.4°N.
In addition to transporting water, the EGC also transports a substantial amount of ice parallel to the coast. Foldvik et al. (1988) stated that near its northern end the EGC carries about 4-5000 km3 of ice equatorward annually. Aagaard and Carmack (1989) estimated a sea ice flux of 2790 km3 per year, while Mitchelson-Jacob (1993) estimated a transport of 3000 km3 a year. Martin and Wadhams (1999) used satellite data from two different sensor types and a model to obtain a sea ice transport estimate of 1530 kme per year in 1994. This value was smaller than previous estimates. As an explanation, the authors cited the estimation of a smaller ice thickness in Fram Strait and the consideration of a strong zonal gradient in the ice drift velocity. The simulation of seasonal ice flux showed a maximum from October to December and smaller values from January to March.
Eddies in the EGC typically have cross-stream length scales of order 10 km. Although they are abundant, the eddies do not contribute much to the heat flux; thus, local baroclinic instability is probably not a major contributor to the mesoscale eddies (Foldvik et al. 1988). Jonsson et al. (1992), using more than 50 record years of current observations, concluded that, at least in the central and eastern Fram Strait, most of the observed eddy kinetic energy is generated by wind fluctuations. Overall, the kinetic energy is low, and the current is dominated by small-scale phenomena in both space and time (Jonsson et al. 1992).
Aagaard, K., and L.K. Coachman, 1968a: The East Greenland Current north of Denmark Strait, Part I. Arctic, 21, 181-200.
Aagaard, K. and L.K. Coachman, 1968b: The East Greenland Current north of Denmark Strait, Part II. Arctic, 21, 267-290.
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Bersch, M., 1995: On the circulation of the northeastern North Atlantic. Deep-Sea Research, 42, 1583-1607.
Foldvik, A., K. Aagaard, and T. Torresen, 1988: On the velocity field of the East Greenland Current. Deep-Sea Research, 35, 1335-1354.
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Mitchelson-Jacob, G., 1993: Eddies in the Greenland Sea observed from infrared and visible satellite radiometry. Oceanologia Acta, 16, 213-220.
Muench, R.D., M.G. McPhee, C.A. Paulson, and J.H. Morrison, 1992: Winter oceanographic conditions in the Fram Strait Yermak Plateau region. Journal of Geophysical Research, 97, 3469-3483.
Rudels, B., H.J. Friedrich, and D. Quadfasel, 1999: The arctic circumpolar boundary current. Deep-Sea Research Part II, 46, 1023-1062.
Schlichtholz, P. and M.N. Houssais, 1999: An investigation of the dynamics of the East Greenland Current in Fram Strait based on a simple analytical model. Journal of Physical Oceanography, 29, 2240-2265.
Woodgate, R.A., Fahrbach, E., and Rohardt, G., 1999: Structure and transport of the East Greenland Current at 75°N from moored current meters. Journal of Geophysical Research, 104, 18059-18072.