Many theories on the origin of green-colored layers have been brought forth, and the most common one is to attribute their genesis to diagenetic alteration of volcanic material. Gardner et al. (1986) suggested that the pale green laminae found on sediment from the Lord Howe Rise east of Australia were the product of alteration of volcanic material, as their occurrence correlated with the distribution of volcanic material and their mineralogy corresponded to bentonitic material. Lind et al. (1993) described colored bands from the Ontong Java Plateau sediments of Oligocene to Pleistocene age. They assigned a diagenetic origin to the laminae, as the colored bands crosscut bioturbation. No correlation between colored bands and proxies of productivity and climate was found, but their temporal distribution parallels the distribution of laminae observed at the Lord Howe Rise. Thus, Lind et al. (1993) ascribed their formation at the diagenetic alteration of volcanic material, as suggested by Gardner et al. (1986).
The origin of green granules on the Vietnam shelf was, on the other hand, attributed to diagenetic alteration of terrigenous material brought by the Mekong River into the SCS (Markov et al., 1996). These granules are mainly formed in nearshore environments characterized by low-energy sedimentation conditions. Quartz, feldspars, and clay minerals such as chlorite and micas are present in these sediments. Once settled in marine sediments, reducing conditions due to relatively high contents of organic matter altered the clay minerals, producing the characteristic green color.
There are few points against the volcanic origin for the GCLs recovered in the SCS. The distribution of volcanic layers found in the sediments from the SCS (Fig. F4) shows no striking similarity with the occurrences of GCLs. Several volcanic ash layers intercalated with GCLs were observed throughout the sediments during Leg 184, and, generally, no sign of green color was detected and the volcanic glass was unaltered. Moreover, green layers did not crosscut preexisting sedimentary features and were generally thicker (up to 3 cm vs. a few millimeters) than the colored bands found by Lind et al. (1993). Both Lind et al. (1993) and Gardner et al. (1986) did not notice any remarkable distinction between the colored layers and the matrix. Despite the overall variability of the measured parameters (see Table T3), it is still possible to detect some differences between GCLs and the host sediments: GCLs are relatively richer in iron, nickel, potassium, chlorite, smectite, and mixed-layered clays and are depleted in carbonate, authigenic P, barium, and strontium (Tables T3, T4).
We may explain the differences between GCL and host sediment composition, invoking diagenetic alteration of original detrital material (e.g., glauconitization) as the main process influencing the genesis of GCLs. SCS sediments are mainly constituted by nannofossil ooze mixed with detrital material rich in micas, quartz, and kaolinite. Relatively high contents of organic carbon, generally higher than 0.2 wt% in the investigated sections with peaks up to 1.2 wt%, are reported in SCS sediments recovered during Leg 184 (Wang, Prell, Blum, et al., 2000). These values are higher than the values reported for the Ontong Java Plateau sites, which average 0.1 wt% (Lind et al., 1993).
Reducing microenvironments could have formed close to the water/sediment interface. The active chemical exchange at the water/sediment interface is considered essential for the complete development of the glauconitization process (Chamley, 1989; Odin, 1988). The few observations of GCLs at Site 1144 in the northern margin of the SCS may, in fact, depend on the extremely high sedimentation rates that characterize this location (Table T1); such sedimentation rates would hasten the burial of sediments, impeding the chemical exchange between the seawater and the sediments. Moreover, redeposition processes and focused sedimentation (Wang, Prell, Blum, et al., 2000) could have disrupted the GCLs, making their detection difficult (Berger and Lind, 1997).
Increased microbial activity in the reducing microenvironments could have favored the degradation of organic matter, which is generally depleted in the GCLs compared to the surrounding matrix. These conditions lead to the dissolution of carbonates, which may explain the observed decreases of CaO, barium, and strontium in GCLs, normally associated with carbonates in marine sediments. Clays are altered, and mixed-layered clays are present in the sediments (Fig. F3). Iron is present at higher concentrations, bound both to silicates and to sulfides. Iron in a reduced form is responsible for the characteristic green color that disappears under oxygen-rich conditions.
The existence of a reducing environment during GCL formation is also attested to by relatively high concentrations of nickel that, being chalcophile, tends to bind to iron sulfides. Another indication of reducing conditions is given by low amounts of authigenic P, which is mainly constituted by carbonate fluorapatite (CFA). The precipitation of CFA in marine sediments is dependent on many factors such as pore water concentrations of dissolved phosphate, oxygen content, and alkalinity (Jarvis et al., 1994). Degradation of organic material is the primary source of dissolved phosphate, but with high carbonate alkalinity of pore waters, CFA cannot precipitate (Jarvis et al., 1994; Krajewski et al., 1994).
The extent of glauconitization normally depends on the nature and grain size of the original material (Odin, 1988; Chamley, 1989). The presence of very fine grained particles, as is the case for sediments from the SCS, allows glauconitization to proceed only to the very first stage of the transformation, where glauconite has not appeared yet but some differences between the altered material and the host sediments have already become evident (Odin, 1988). Despite the fact that the SCS could be a suitable environment for verdine formation (Kronen and Glenn, 2000), this authigenic green clay mineral that develops in Fe-rich and shallow environments (between 5 and 60 m water depth) (Odin, 1988) has not been found in sediments recovered during Leg 184.
Glauconitization normally occurs at shallow depths, but the existence of glauconitic layers has already been reported in deeper waters (between 2000 and 3000 m water depth), comparable to SCS site locations (Chamley, 1989). We cannot exclude the possibility, considering the depth of the sites and the nature and geometry of the GCL, that they represent reworked material (i.e., distal turbidites) that have already undergone glauconitization or the higher porosity of which could have promoted glauconitization. Indeed, turbidites and reworked material have been reported at all Leg 184 sites (Wang, Prell, Blum, et al., 2000).
Many authors have used the distribution of glauconitic levels to indicate sea level changes over time (Chamley, 1989), as transgression and regression cycles can expose sediments located at different depths to conditions favorable to glauconitization. All sites except Site 1143 are characterized by nonrandom distribution of GCLs, and a trend in the length of intervals between successive GCL events is evident especially for Sites 1145 and 1146 (Table T5; Fig. F5). Comparison of these trends with the neotectonic curve calculated for the SCS northern margin (Lüdmann et al., 2001) indicates that since 600 ka, when the northern margin started uplifting, the frequency of GCLs was increased (Fig. F5). This observation supports the hypothesis of the reworked nature of the GCL, as tectonic activity and uplift may have increased the frequency of earthquakes and turbidites. Several foraminifer turbidites were observed at all SCS sites, and the relatively higher porosity of these turbidites (see Wang, Prell, Blum, et al., 2000) could have favored active fluid circulation in the sediments, promoting diagenetic alteration of the original material.
A correlation, although poor, seems to exist between interglacial periods characterized by high sea level and longer intervals between GCL events (Fig. F5), especially for Site 1143. This site is located close to the Sunda shelf, a region particularly sensitive to sea level changes (Pelejero et al., 1999). The temporal distribution of GCLs and the results of the autocorrelation point to cyclicity only at Site 1148, with frequencies close to values of primary frequencies and heterodynes of eccentricity. It is difficult to assess if this analysis bears reliable information: we are aware of the uncertainties in the temporal distribution of the GCLs, which arise mainly from difficulties in the visual recording process of the layers (i.e., subjectivity of the different operators in defining a layer) and from the preliminary age models used in the analysis. Moreover, the spectral analysis did recover only the peak at 127 k.y., which is relatively close to a primary eccentricity frequency (e.g., peak at 123.8 k.y.); otherwise, only heterodynes of eccentricity have been observed. Nevertheless, we cannot exclude that changes in sea level resulting from the combination of climatic changes and tectonism could have had any influences on the distribution and genesis of the layers. We know that during the last glaciation the shelf surrounding the SCS was almost completely exposed (Pelejero et al., 1999). Therefore, during low sea level periods the zone of active glauconitization may have moved offshore, and as conditions at the sites became more favorable to diagenetic alteration the intervals between GCLs became shorter. Moreover, during regression and low sea level, turbidites were likely to be more frequent, thus supporting the hypothesis of the possible reworked nature of the GCLs.