341
models (e.g., Haney, 1971; Marotzke and Willebrand, 1991). Notice, however, that eq. (19) connects spatial differences in surface heat flux to
temperature differences. It is readily shown (MS) that the spatial mean
heat flux can likewise be written as a Newtonian law for the mean temperature, but spatial mean and difference are restored with different time
constants. This 'scale dependence' of the restoring coefficient is equivalent
to the heat fluxes, HI and H 2 , not being computable by a purely local
relationship, but involving horizontal atmospheric transports [see eq. (14);
see Marotzke, 1994, for a detailed discussion].
2.3 Surface freshwater fluxes
In analogy to the high-latitude atmospheric heat budget, eq. (11), we write
{ Eda+Fw = 0,
iFol
(22)
where E is the evaporation minus precipitation. The integral in (22) also
comprises precipitation over land, which through river runoff influences
the oceanic freshwater budget. To compute E over the ocean, one has to
consider the ratio, lOW, of the ocean area to the catchment area of the ocean
basin. Taking E to be constant along a latitude circle, the range of lOW is
(23)
meaning that in one extreme (lOW = 1) the ocean basin receives only the
moisture transported in the atmosphere right above it, whereas in the other
extreme (lOW = 10, meaning that the catchment area is the entire latitude
circle) it catches all the river runoff as well. With the Atlantic in mind, lOW
in the range 0.3 to 0.5 is a reasonable value since the Atlantic (including the
Arctic) receives runoff from almost the entire Americas and a large portion
of Asia (e.g., Baumgartner and Reichel, 1975; Broecker et al., 1990). Here,
lOW = 10 = 0.5 is chosen, which corresponds to a single ocean basin of the
size of the Pacific that receives all river runoff, from a continent the same
size as the ocean, through zonal river flow. This choice of ocean area is
made for convenience in the perturbation experiments described in Section
3.
Thus, we obtain for E over the ocean
E_~Fw
- lOW FOI'
(24)
models (e.g., Haney, 1971; Marotzke and Willebrand, 1991). Notice, however, that eq. (19) connects spatial differences in surface heat flux to
temperature differences. It is readily shown (MS) that the spatial mean
heat flux can likewise be written as a Newtonian law for the mean temperature, but spatial mean and difference are restored with different time
constants. This 'scale dependence' of the restoring coefficient is equivalent
to the heat fluxes, HI and H 2 , not being computable by a purely local
relationship, but involving horizontal atmospheric transports [see eq. (14);
see Marotzke, 1994, for a detailed discussion].
2.3 Surface freshwater fluxes
In analogy to the high-latitude atmospheric heat budget, eq. (11), we write
{ Eda+Fw = 0,
iFol
(22)
where E is the evaporation minus precipitation. The integral in (22) also
comprises precipitation over land, which through river runoff influences
the oceanic freshwater budget. To compute E over the ocean, one has to
consider the ratio, lOW, of the ocean area to the catchment area of the ocean
basin. Taking E to be constant along a latitude circle, the range of lOW is
(23)
meaning that in one extreme (lOW = 1) the ocean basin receives only the
moisture transported in the atmosphere right above it, whereas in the other
extreme (lOW = 10, meaning that the catchment area is the entire latitude
circle) it catches all the river runoff as well. With the Atlantic in mind, lOW
in the range 0.3 to 0.5 is a reasonable value since the Atlantic (including the
Arctic) receives runoff from almost the entire Americas and a large portion
of Asia (e.g., Baumgartner and Reichel, 1975; Broecker et al., 1990). Here,
lOW = 10 = 0.5 is chosen, which corresponds to a single ocean basin of the
size of the Pacific that receives all river runoff, from a continent the same
size as the ocean, through zonal river flow. This choice of ocean area is
made for convenience in the perturbation experiments described in Section
3.
Thus, we obtain for E over the ocean
E_~Fw
- lOW FOI'
(24)