General Characteristics of Density-Turbidity Currents in the Ross Sea (Antarctica)
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since oceanic forcings lasting 103 s are rather frequent. So a fondamental différence is found between the kind of diffusivity that plays a rôle in laboratory simulations and real oceanic eddy diffusion where a much larger value plays an important rôle. Ail these considérations finally suggest, as a practical rule, that it is a
realistic value of v that is critical. It is consequently of interest that t* can be realistically estimated by comparing the two most popular expressions of bottom
stress, namely i=p0CdU2 and r=povôU/dz, in this way giving rise to
PoCdU2 ~ povU/h,
where a realistic Cd = 3,5 x 10'3 (MKS units) can be assumed and, consequently,
v ~ CdhU
t* = h02 / (tc2v) -> h0/(n2CdU),
a fairly easily applicable relation.
5 Application to Antarctic Currents
Let us apply these considérations to hydrologie sections along the Antarctic coast
off Cape Adare, as shown in Fig. 1. The bottom depth ranges from about 2200 m
to about 500 m, the shelf. It is immediately apparent that a very dense salty water
is présent over the deepest part of the shelf, with S = 34.75-34.80 ppt and T = -1.9
to -1.5 °C. Its geostrophic velocity can be estimated as 5-10 cm s1 and the shelf
layer is about 100 m thick. This looks like the classical geostrophic situation of any
dense water présent over the shelf.
It is also of interest that a vein of dense water is évident at the bottom of the
shelf break (Fig. 3). Interestingly, this water is very cold and its density is 27.87,
slightly denser than the bottom water observed in the adjacent stations. This
denser layer has a thickness that can be estimated as h0 = 100-200 m, but the velocities are rather small, namely 1 cm s1. This bottom water is not particularly salty
but is very cold: one is led to assume that its origin is different from that of the salty
shelf water, probably due to ice melting, and it arrives at the bottom of the shelf
break via a different path, probably along a submarine dépréssion skirting the
shelf.
It is of some interest, consequently, to estimate the corresponding values of t*.
For this deep layer we assume that the hydrologie characteristics observed at station 116 are due to dense water reaching this deep layer flowing along a submarine canyon. We thus estimate the friction v = CdhU ~ 5 x 10'3 (MKS units), in
agreement with classical values. It results that t* = h20/(ir2v) = 2 x 106 s and so the
observed deep, slow current must essentially be considered stable. This is not surprising since for this deep flow no hydrodynamic instability is évident.
We also apply the same considérations to the salty shelf water discussed earlier. For geostrophic velocity u ~ 5 x 10'2 ms'1 and h0 ~ 100-200 m, we obtain an
“ignition” time t* = 5 x 104 s. This estimate of an eventual “ignition” time is
smaller by a considérable factor than the previous estimate of t* for the deep
layer. Note that the water velocity over the shelf is much smaller than an even-
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since oceanic forcings lasting 103 s are rather frequent. So a fondamental différence is found between the kind of diffusivity that plays a rôle in laboratory simulations and real oceanic eddy diffusion where a much larger value plays an important rôle. Ail these considérations finally suggest, as a practical rule, that it is a
realistic value of v that is critical. It is consequently of interest that t* can be realistically estimated by comparing the two most popular expressions of bottom
stress, namely i=p0CdU2 and r=povôU/dz, in this way giving rise to
PoCdU2 ~ povU/h,
where a realistic Cd = 3,5 x 10'3 (MKS units) can be assumed and, consequently,
v ~ CdhU
t* = h02 / (tc2v) -> h0/(n2CdU),
a fairly easily applicable relation.
5 Application to Antarctic Currents
Let us apply these considérations to hydrologie sections along the Antarctic coast
off Cape Adare, as shown in Fig. 1. The bottom depth ranges from about 2200 m
to about 500 m, the shelf. It is immediately apparent that a very dense salty water
is présent over the deepest part of the shelf, with S = 34.75-34.80 ppt and T = -1.9
to -1.5 °C. Its geostrophic velocity can be estimated as 5-10 cm s1 and the shelf
layer is about 100 m thick. This looks like the classical geostrophic situation of any
dense water présent over the shelf.
It is also of interest that a vein of dense water is évident at the bottom of the
shelf break (Fig. 3). Interestingly, this water is very cold and its density is 27.87,
slightly denser than the bottom water observed in the adjacent stations. This
denser layer has a thickness that can be estimated as h0 = 100-200 m, but the velocities are rather small, namely 1 cm s1. This bottom water is not particularly salty
but is very cold: one is led to assume that its origin is different from that of the salty
shelf water, probably due to ice melting, and it arrives at the bottom of the shelf
break via a different path, probably along a submarine dépréssion skirting the
shelf.
It is of some interest, consequently, to estimate the corresponding values of t*.
For this deep layer we assume that the hydrologie characteristics observed at station 116 are due to dense water reaching this deep layer flowing along a submarine canyon. We thus estimate the friction v = CdhU ~ 5 x 10'3 (MKS units), in
agreement with classical values. It results that t* = h20/(ir2v) = 2 x 106 s and so the
observed deep, slow current must essentially be considered stable. This is not surprising since for this deep flow no hydrodynamic instability is évident.
We also apply the same considérations to the salty shelf water discussed earlier. For geostrophic velocity u ~ 5 x 10'2 ms'1 and h0 ~ 100-200 m, we obtain an
“ignition” time t* = 5 x 104 s. This estimate of an eventual “ignition” time is
smaller by a considérable factor than the previous estimate of t* for the deep
layer. Note that the water velocity over the shelf is much smaller than an even-