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top and another at the bottom allowed circulation between the two test tanks. One storage tank containing hot salty water along with its test tank represented the tropical ocean. The other storage–test tank pair with cold and freshwater represented a polar region. The model illustrated the properties of convection with combined heat flux and freshwater flux boundary conditions. As the governing parameters (temperature and salinity differences between the two storage tanks) were slowly changed
top and another at the bottom allowed circulation between the two test tanks. One storage tank containing hot salty water along with its test tank represented the tropical ocean. The other storage–test tank pair with cold and freshwater represented a polar region. The model illustrated the properties of convection with combined heat flux and freshwater flux boundary conditions. As the governing parameters (temperature and salinity differences between the two storage tanks) were slowly changed
multidecadal memory of the PMV may reside in the slow westward propagation of the first baroclinic-mode Rossby waves in the subpolar North Pacific Ocean. Furthermore, the subsurface thermocline or halocline variability may affect the surface climate through local convective feedback as well as oceanic advection and interaction with the Kuroshio Extension (KOE). In particular, salinity variability is suggested to provide an indispensable ingredient for both the time-scale selection and convective growth
multidecadal memory of the PMV may reside in the slow westward propagation of the first baroclinic-mode Rossby waves in the subpolar North Pacific Ocean. Furthermore, the subsurface thermocline or halocline variability may affect the surface climate through local convective feedback as well as oceanic advection and interaction with the Kuroshio Extension (KOE). In particular, salinity variability is suggested to provide an indispensable ingredient for both the time-scale selection and convective growth
temperature and salinity structure in an indirect manner.) Figure 1 displays the approximate time required for the baroclinic Rossby wave with the fastest group velocity to cross a 5000-km-wide ocean as a function of latitude. The fastest wave is, in the basic theory, the one with the basin-scale wavelength. The multidecadal time scale in the modeling results of Cessi et al. (2004) is fully consistent with expectation. These times (and basin widths do shrink with latitude) are not the adjustment time
temperature and salinity structure in an indirect manner.) Figure 1 displays the approximate time required for the baroclinic Rossby wave with the fastest group velocity to cross a 5000-km-wide ocean as a function of latitude. The fastest wave is, in the basic theory, the one with the basin-scale wavelength. The multidecadal time scale in the modeling results of Cessi et al. (2004) is fully consistent with expectation. These times (and basin widths do shrink with latitude) are not the adjustment time
inferred eddy stress is computed as the residual necessary to minimize the difference between the observed global temperature–salinity distribution of the ocean and the model. The magnitudes of mixing rates found in that study are consistent with other studies (e.g., Marshall et al. 2006 ) and are strongly depth dependent. Here we draw on additional evidence that is strongly suggestive of depth-dependent mixing. Mesoscale eddies are widely believed to be generated by baroclinic instability of the mean
inferred eddy stress is computed as the residual necessary to minimize the difference between the observed global temperature–salinity distribution of the ocean and the model. The magnitudes of mixing rates found in that study are consistent with other studies (e.g., Marshall et al. 2006 ) and are strongly depth dependent. Here we draw on additional evidence that is strongly suggestive of depth-dependent mixing. Mesoscale eddies are widely believed to be generated by baroclinic instability of the mean
the separated boundary current transport and moves the location of the outcrop southeastward: Here, h 1 u 1 † is the nondivergent time-mean transport. Generally, layered tracer equations take the form for the tracer concentration ϕ , which could be salinity, potential temperature, potential density, etc. After averaging, the upper interior layer density equation is 14 As in Nurser and Williams (1990) , the mass in the upper layer is diminished by entrainment into the mixed layer by w 1
the separated boundary current transport and moves the location of the outcrop southeastward: Here, h 1 u 1 † is the nondivergent time-mean transport. Generally, layered tracer equations take the form for the tracer concentration ϕ , which could be salinity, potential temperature, potential density, etc. After averaging, the upper interior layer density equation is 14 As in Nurser and Williams (1990) , the mass in the upper layer is diminished by entrainment into the mixed layer by w 1
