, 2005 and Montes et al., 2010). A comparison of the two right panels of Fig. 6a illustrates PD-0332991 molecular weight the striking differences that can occur in the propagation of dynamical (δ′TSEδ′TSE) and spiciness (δ″TSEδ″TSE) signals, in this case with the spiciness signal extending more prominently equatorward. Similar differences occur in our other regional solutions, and have been noted previously by Nonaka and Xie, 2000 and Taguchi and Schneider, 2013. Equatorial response. Fig. 6b illustrates the vertical structures of the dynamic and spiciness anomalies along the equator, plotting δTSE,δ′TSEδTSE,δ′TSE, and
δ″TSEδ″TSE fields averaged from 1 °S to 1 °N. Consistent with Fig. 6a (top panels), the deep dynamical signal δ′TSEδ′TSE ( Fig. 6b, middle panel) is spread throughout the equatorial ocean. There is also a near-surface, positive anomaly that is locally generated (see below). It is noteworthy that δ′TSEδ′TSE has fewer zero
crossings at the equator than it does in the forcing region ( Fig. 4b, middle-left panel south of 8 °S), an indication that either Rossby waves associated with higher-order vertical modes are preferentially damped or the large change in stratification modifies the structure of the modes. Also consistent with Fig. 6a, there is a strong, negative spiciness signal δ″TSEδ″TSE within the pycnocline selleck inhibitor ( Fig. 6b, bottom panel), which is advected to the equator from Region Mephenoxalone SE along the two pathways noted above. Below the pycnocline, there is a positive anomaly (bottom panel) near the western boundary. Most of it flows out of the basin as a deep part of the ITF (not shown; e.g., McCreary et al., 2007), with some bending eastward to join the southern
Tsuchiya Jet and the lower part of the EUC. There are also negative δ″TSEδ″TSE and positive δ′TSEδ′TSE signals above the pycnocline. Because of surface fluxes, however, these signals cannot be interpreted as arising solely from the remote forcing region. The negative δ″TSEδ″TSE signal is advected along the equator within the pycnocline by the EUC and is mixed upward into the surface layer in the eastern Pacific. The heat flux into the ocean is increased there, reducing the negative temperature anomaly (Fig. 6b, top panel) and leaving behind a negative salinity anomaly. At the same time, evaporation is reduced owing to the lower SST while precipitation is not affected (Section 2.1), enhancing the negative salinity anomaly. This anomaly is advected westward by the surface South Equatorial Current while the negative temperature anomaly is almost erased by the surface heating before reaching the western Pacific. Since the dominance of negative salinity anomaly implies a negative density anomaly in the western Pacific, the vanishing temperature anomaly there is projected onto δ′T>0δ′T>0 (see Eq. A.2c), resulting in δ″T<0δ″T<0 since δT=δ′T+δ″TδT=δ′T+δ″T.