The Three-dimensional Temperature Distribution and its Variation in Time 121 



Table 50. Heat transport downwards assuming a temperature gradient of 1 °C/ 1 00 m 



Vertical exchange coefficient (/Ij g cm~^ sec"^) 20 10 5 2-5 1 



Heat amount (g cal cm-2 day-i) 172 86 43 21-5 8-6 



Beneath the thermocline from about 200-300 m the water masses of the sub- 

 troposphere are remarkably constant in their nature and geographical distribution. 

 The vertical temperature gradient in these waters rapidly decreases with depth and 

 gradually changes its magnitude into that of the stratosphere. Considerable amounts 

 of heat are transported by dynamic convection through the layer immediately be- 

 neath the almost isothermal top layer to the layer below. Table 50 gives an idea of the 

 quantities of heat involved; it assumes a mean temperature gradient of 1°C/100 m. 

 These amounts of heat are surprisingly high. Even for small values of A^, the down- 

 ward heat flux amounts to 10-40 gcal cm "May ""^. Since there is always a tem- 

 perature gradient, this raises the very natural question of where all this heat goes to. 

 In the lower layers of the troposphere the temperature gradient is again smaller and 

 therefore the downward heat flux becomes smaller again in the middle layers of the 

 troposphere ; the accumulation of heat in these layers should soon destroy the vertical 

 temperature gradient and thus also the thermocline. It must therefore be true that the 

 vertical temperature gradient in the troposphere can only be maintained if the lateral 

 influx of colder water compensates the flow of heat from above and indeed the heat 

 from above and the horizontal advection must compensate each other exactly. The 

 vertical temperature distribution in the troposphere is thus maintained in a stationary 

 state by the oceanic circulation (Defant, 1930). 



The cause for formation of the thermocline below an almost isothermal top layer 

 in the tropics and the subtropics is therefore as follows: The top layer is certainly 

 more or less in thermal equilibrium with the atmosphere above. The lower tempera- 

 tures of the lower subtroposphere and of the stratosphere are essentially of polar 

 origin; as they flow towards the equator these water masses mix with warmer water and 

 thereby gain heat, but are continually renewed and are thus kept at a relatively low 

 temperature. It would be expected that the diff"erence between the high temperature 

 at the top and the low temperature of the deeper layers would give rise to a roughly 

 linear vertical temperature gradient in the middle layer; instead a homogeneous top 

 layer is formed and the transition to the lower temperatures of the subtroposphere 

 takes place abruptly in a well-developed transition layer (thermocline). 



The explanation of this thermal stratification in the tropics and the subtropics lies 

 in the same circumstances that give rise to the summer transition layer in lakes as well 

 as in the ocean. The turbulence induced by the wind and the waves will slowly trans- 

 port the heat from the upper layers downwards and the temperature diff"erences thus 

 formed will work their way down into deeper and deeper layers. However, further 

 rise in temperature in the top layers will also increase the vertical density gradient. 

 The downward transfer of heat from above by turbulence will cease when the increase 

 in vertical stability diminishes the intensity of the turbulence. If the vertical density 

 gradient is very strong the turbulence of the flow cannot overcome the great stability 

 of the stratification and a transfer of heat to a deeper level through the thermocline 

 can no longer occur. In the top layer the turbulence leads finally to a complete 



