By frictional mixing we mean that which is caused by the uneven movement of contiguous 

 water masses; i.e. , by the presence of vertical and horizontal velocity gradients. At the surfaces 

 of separation of water masses, these gradients create eddies, penetrating from one water mass to 

 another, and in this manner, mixing them. 



By convective mixing, we mean that which arises as a result of either a decrease in the den- 

 sity of the deep layers, or of an increase in the density of the surface layers of the sea. In either 

 case, there arise, in the water mass, vertical currents which cause mixing of the superposed 

 layers. 



The main difference between convective and frictional mixing is that convective mixing may 

 occur whether or not the given layers are in motion, and that it takes place only in a vertical direc- 

 tion. Frictional mixing depends on the presence of vertical and horizontal velocity gradients, and 

 in this regard, they may occur both horizontally and vertically. 



If the horizontal velocity gradient is found at the surfaces of separation of water masses of 

 different density, as happens, e.g. , when sea currents flow into basins with different physical- 

 chemical characteristics, convective and frictional mixing may occur simultaneously. Frictional 

 mixing may also cause convective mixing when the horizontal layers differ little with respect to 

 density and the mixing causes "a density increase" of the layers. 



The rate of vertical mixing depends most strongly on the resistance which individual layers 

 display to mixing. This resistance is determined by the stability of the layers and, according to 

 Hesselberg and Sverdrup, this stability has the value: 



dt dz^dS dz dt dz' ^^^ 



where a = the specific volume, 

 t = the temperature, 

 S = salinity, 



I = the adiabatic temperature change. 

 Formula 1 can be represented in another way: 



'^ dz dt dz- ^' 



The first member of the right hand side of this formula is the vertical specific volume gradient, 

 and the second member is the specific volume corrected for adiabatic temperature change. 



Since the adiabatic correction is very small, it may be disregarded when judging the stability 

 of the upper ocean layers, (where the vertical specific volume gradients are large). At the lower 

 depths, where the layers are extremely uniform with regard to temperature and salinity and where, 

 accordingly, the vertical specific volume gradient is very close to 0, the adiabatic correction may 

 play a decisive role. 



Usually, in the ocean, the stability is positive, i.e. , the lighter layers lie above the heavier 

 ones. But in certain regions, negative stability, i.e. , density inversion, can be observed in the 

 intermediate layers. This is explained by the presence of sea currents, consisting of waters of 

 different origin superimposed. Finally, in individual cases, during mixing caused, e.g. , by strong 

 cooling during the winter, stability may be negative even in the upper layers. This indicates that 

 cooling of the ocean surface occurs more rapidly than convection. 



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