A good correlation exists between planktonic foraminifer changes and the oxygen isotopic record (Figs. F2, F3, F4, F5). If the isotope record indeed chiefly reflects the waning and waxing of polar ice, the planktonic foraminifer results reported here must have resulted from global ice volume changes through glacial–interglacial cycles. Prior to 0.9 Ma (below 420 mcd), however, this relationship was vague for many individual species. Abundant warm-water species occur below 420 mcd, indicating a much warmer climate regime before the MPR. This warm climate regime caused relatively smaller glacial–interglacial contrasts and smaller summer–winter SST differences than after 0.9 Ma. The estimated SST using transfer function FP-12E shows an average difference of 7.5°C between summer and winter SST (28.5°–21.0°C) for the 300- to 420-mcd interval. The strongest winter SST differences occurred between 410 and 422 mcd, in MIS 23 to lower 22, bracketing the MPR event (Fig. F6). The pre-MPR warm climate regime changed gradually, however, as indicated by increases upsection of cool-water species including G. bulloides and G. inflata.
In spite of being relatively warm before the MPR, glacial–interglacial transitions appear to have been rapid. The total warm- to cold-water species ratio and G. menardii to G. inflata ratio both show sudden jumps, indicating rapid transitions across glacial/interglacial boundaries (Fig. F4). These rapid changes from one climate stage to another continued into younger periods, although the overall climate has shifted to one dominated by much cooler conditions with more abundant cool-water species after the MPR.
The SST differences enlarged after the MPR, accompanied by stronger glacial–interglacial contrasts in planktonic foraminifer abundances. This shift appears to be mainly influenced after the MPR by much cooler winters driven by stronger glaciations. The cooler climate regime started affecting planktonic foraminifers immediately after the MPR, and climate deteriorations, began early even in the later part of interglacials and continued across interglacial/glacial boundaries. For instance, abundance of cool- to cold-water species such as G. bulloides, G. pachyderma, and G. inflata increased in the later part of MIS 21, 19, and 17, although the increases were not always simultaneous or at a similar amplitude (Fig. F4). These progressive changes in planktonic foraminifer response to interglacial–glacial transitions mimic the pattern of changes from MIS 5–2 (Wang and Wang, 1990; Wang, 1999; Wang et al., 1999; Jian et al., 2000b), indicating that such typical late Pleistocene interglacial–glacial transitions first began ~0.9 m.y. ago, immediately after the MPR. In contrast, faunal changes from glacial–interglacial cycles were rapid after the MPR, especially between MIS 22 and 17, as indicated by the warm- to cool-species ratio (Fig. F4).
As recorded elsewhere, the first strongest glaciation in the last 1 m.y. occurred at MIS 16, which dwarfed all tropical–subtropical species while encouraging an exceptionally high production of cool-water species at Site 1144 (Fig. F4). The impact of this glaciation continued into the subsequent interglacial MIS 15 and glacial MIS 14 with minimal recovery of warm-water species.
Several planktonic foraminifer species are good indicators of upper water stratification because they are found living mainly in certain layers of the world ocean. G. ruber is a typical surface water species living in the upper 50 m of the water column, whereas globorotaliid forms are more frequent in deeper waters with relative heavy 
18O (Fairbanks et al., 1982; Helemben et al., 1989). Surface-water species increased in abundance when the thermocline deepened, and vice versa for those deepwater dwellers including Globorotalia, Pulleniatina, and Sphaeroidinella (Anderson and Ravelo, 1997; Jian et al., 2000a, 200b).
Figure F5 shows abundance variations of the deep-dwelling group and its major constituents, apart from G. aequilateralis and Orbulina spp., which are ubiquitous in the tropics and subtropics. The total abundances of the deep-dwelling group fluctuate mainly between 30% and 50% throughout the studied interval, indicating that the overall upper water structure in the northern South China Sea has been stable during the early and mid-Pleistocene. The abundance variations of these deep dwellers, however, did not responded well to glacial–interglacial cycles until MIS 21, immediately before the MPR. Only after MIS 21 did all deep dwellers exhibit a close relationship with isotopic fluctuations. The increase in P. obliquiloculata and G. menardii groups during interglacials at and after MIS 21 likely indicate the development of a well-constrained thermocline that shoaled during interglacials and deepened during glacials. In contrast, S. dehiscens decreased substantially after MIS 23, and since then its abundance has remained extremely low. The reductions of S. dehiscens after MIS 23 and G. tumida after MIS 17 were probably due to a cooler climate regime and weaker warm currents, as indicated also by the decrease in many shallow-dwelling warm-water species close to the MPR (Fig. F3).
Therefore, planktonic foraminifer response to the mid-Pleistocene transition was gradual and progressive, with species abundance changes occurring before, at, and after the MPR. It is not clear, however, whether and how much these faunal variations were influenced by the dynamics of Asian monsoons and/or of the Western Pacific Warm Pool during the mid-Pleistocene climate transition.
The MPR marks the transition from 41- to 100-k.y. cyclicities, shaping the general pattern of the two-moded Quaternary climate as recorded in deep-sea oxygen isotopes (Prell, 1982; Berger et al., 1993; Raymo et al., 1997; Schmieder et al., 2000; Wang et al., 2001). The orbital forcing is also reflected in individual planktonic foraminifers. Figure F7 shows the results of spectrum analysis on four of the most common species using the method described by Schulz and Stattegger (1997). A prominent eccentricity band represented by 80- to 100-k.y. cycles occurs in all four species, most pronouncedly in G. ruber. The obliquity (39–41 k.y.) and precession (19–23 k.y.) cycles and many shorter cycles are also strong, indicating orbital and local climatic forcings including tropical monsoons. For instance, the missing 41-k.y. cycle in G. sacculifer may have been caused by the disturbance of monsoons. Although the MPR is not obvious in Figure F7 because the studied interval is too short to be separated into two sections (before and after 0.9 Ma) for spectrum analysis, these results still demonstrate that (1) faunal responses to orbital forcing climatic changes are different in both frequencies and amplitude, and (2) local factors represented by shorter cycles could have also played an important role.
Jian et al. (2000b) reported foraminifer responses to the mid-Pleistocene climate transition in Core 17957 (10°53.9´N, 115°18.3´E; water depth = 2195 m), in the southern South China Sea. Similar to the results presented above, planktonic foraminifers from Core 17957 show immediate changes in water temperature and thermocline depth at 0.9 Ma, but the benthic foraminifers did not change much until the Brunhes/Matuyama reversal. In the same core, radiolarian abundances also changed close to the MPR, characterized by a decrease in tropical species (Wang and Albermann, 2002). These authors attribute the changes in plankton to a southward shift of the North Equatorial Current likely induced by variations in the northern trade wind system during the MPR, about 0.9 m.y. ago.
At Site 1143 (9°21.72´N, 113°17.11´E; water depth = 2772 m), planktonic foraminifer changes were also in path with a progressive mid-Pleistocene climatic transition (MPT). The most striking feature, however, is that the abundance of P. obliquiloculata became reversed from high in interglacials before the MPR to high in glacials after the MPR (Xu et al., 2005). As P. obliquiloculata prefers high salinity, its high abundance in glacials after the MPR could have been because of high salinity in the southern South China Sea when sea level was as much as 120 m lower than today's, the basin was semienclosed, and evaporation was high. This phenomenon, however, does not exist in the northern South China Sea, including Site 1144, probably because a consistent influence of the west Pacific water through the Bashi Strait (–2600 m) and more frequent winter monsoons during glacial and interglacial periods.
