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The primary sediment composition and the early diagenetic reactions determine subsequent diagenetic pathways. Early diagenetic processes are contemporaneous with deposition and can locally affect the physical properties by cementing depositional porosity and make the mudstone locally resistant to mechanical compaction. These early diagenetic processes are typically redox processes where organic matter becomes oxidised by microbes (Aplin and Macquaker, 2011).

Sulphate reduction was locally important during early diagenesis in the studied Formation E mudstones as evidenced by the common occurrence of framboidal pyrite. Framboidal pyrite likely precipitated from iron-rich porewaters near sites of sulphide supply, usually near organic matter as a result of bacterial sulphate reduction, where sulphide production rates were high enough to reach supersaturation with respect to FeS. Euhedral pyrite likely precipitated directly from iron-rich porewaters when FeS2

134 sites of sulphide supply due to lower sulphide production rates from organic matter (Taylor and Macquaker, 2000).

The presence of siderite indicates that Fe(III) reduction was also important (at least locally) during early diagenesis (Taylor, 1998). Cementation by siderite is likely to have occurred prior to compaction as a result of bacterially mediated respiratory processes. This indicates that sites of cementation were close to the sediment-water interface for prolonged periods of times in order to allow for cementation to take place (e.g. Taylor and Macquaker, 2000, 2014). Although siderite precipitation had locally affected mudstone porosities its overall effect on porosity reduction in the studied Formation E mudstones is negligible.

The Bravo-Y CEC and clay fraction XRD results reveal a diagenetically mature clay mineral assemblage consistent with the high geothermal gradient. It should be noted that the question whether the illitic nature of the samples is depositional or diagenetic in origin cannot be comprehensively answered by this study as shallower samples were not available. However, the observed negative correlation of plagioclase with quartz and with Kfeldspar and its positive correlation with illite-smectite in Bravo-L mudstone samples does give some indication of a diagenetically mature clay mineral assemblage. Since in these mudstones sodium resides exclusively in the plagioclase feldspar (Figure 4.18e, f) positive correlation between plagioclase and illite-smectite strongly suggest that sodium was originally a constituent in the clay minerals most likely in smectites or in smectitic interlayers in illite-smectite. The lack of such correlation in Bravo-Y might be attributed to the presence of small amounts of detrital plagioclase in those samples since Bravo-Y mudstones are slightly more silt-rich. Bravo-Y mudstone samples are also more heterogeneous due to the poor quality of the provided cuttings which might be another contributing factor.

135

Figure 4.30. Published examples of illite % in illite-smectitie with new results from the Central Malay Basin Bravo-Y well plotted against maximum burial depth (a) and maximum burial temperature (b).

Another indication which shows that clay mineral diagenesis in the Bravo-Y mudstone samples is in the final stages is the lack of pure smectite and smectitic illite-smectite, and additionally the dominance of an ordered (R>1) mixed layer illite-smectite phase. As it is shown in Figure 4.30a when plotted against maximum burial depth the Central Malay Basin mudstones appear considerably more mature than those from other locations. Note that illite-smectites in mudstone samples reported from the Gulf of Mexico and Podhale Basin at similar burial depth are considerably more smectitic and are all pre- transition (R0 to R1). However, when plotted against temperature as opposed to burial depth the Central Malay Basin mineralogy results fit in with reported examples from different locations (Figure 4.30b).

The Bravo-Y results correlate well with clay mineralogy data from Duffy (2011) (Figure 4.30), from two south Malay Basin wells, from Formations H,I,J,K and L. The bulk mineralogy of those mudstones is fairly similar to Formation E with the only main difference being the lack of siderite from the older formations (K and L) which were

136 deposited in a fluvial-lacustrine environment. The illitic nature of those South Malay mudstones was interpreted to be a result of diagenesis based on the presence of a glycolated 16.8 Å peak (expandable/smectitic interlayer) in the shallowest samples. The presence of smectitic layers in the shallowest South Malay samples and the slightly more illite % in illite-smectite in the deeper samples (Figure 4.30) compared to Formation E mudstones can be attributed to differences in geothermal gradient (present day geothermal gradient in the South Malay Basin is 45-50o_{C/km) in combination with}

differences in starting mineralogy (e.g. smectite content, Kfeldspar content) and age. In overall the mineral composition of the studied Formation E mudstones show limited variation and is consistent throughout with a digenetically mature assemblage in which smectite has been illitized.

4.7.3 Evaluation with wireline log-based compaction profiles and implication for overpressure development

Figure 4.31 shows smoothed cosine-weighted density and sonic log responses through the mudstone intervals for wells Bravo L, X&Y. Below the top of Formation Unit D sonic transit time decreases and density increases monotonously with increasing depth until the top of overpressure in all three wells indicating ongoing compaction. Small fluctuations in the smoothed data are the result of changing lithology, most likely related to variations in grain size, heavy mineral and/or clay mineral content. These changes cannot be resolved by the mineralogical results presented in 4.6.1 due to the poor resolution. Small lithological changes cannot be seen since the presented results represent averaged mineralogical composition of several meters of mudstones.

137

Figure 4.31. Smoothed density (blue) and sonic (red) logs through the mudstone intervals for (a) Bravo-X, (b) Bravo-Y and (c) Bravo-L. Top of overpressure is indicated by horizontal black dotted line.

The onset of overpressure is associated with rapid increase of the sonic log response (sonic reversal) while density continues to increase below the top of overpressure in all the wells which is indicative of unloading. Note that the sonic reversal is slightly masked in Bravo-X and L due to the presence of an approximately 15-20 m thick mudstone layer with anomalously low sonic transit time (“high velocity layer”; Figure 4.31a,c and 4.32a,c). Similar layers can be observed in Bravo-Y although slightly below the top of overpressure (Figure 4.31b). High velocity layers are shown in Figure 4.7c from Bravo-Y prior to mudstone discrimination and smoothing. Without prior evidence, they might be interpreted as evidence for lateral drainage, but no other log types show similar

138 excursions. The nature of these layers is uncertain because of the lack of available mudstone samples from these depth intervals.

Figure 4.32. Sonic-density crossplots for Bravo-X (a), Bravo-Y (b) and Bravo-L (c) with smoothed logs. Colour coding is by depth below seafloor (m). Blue and red dots denote the last hydrostatic RFT and the first overpressured RFT respectively. Red and Green line is the illitic and smectitic line from Dutta (2002).

139 Possibilities include cementation by carbonate or silica, however if the cause of the excursion had been carbonate, an excursion on the resistivity log would also be expected and there is none.

Figure 4.32 shows crossplots of smoothed sonic transit time against smoothed density in the studied mudstone intervals for Bravo-X, Bravo-Y and Bravo-L. Dutta's (2002) reference lines for smectite rich and illite rich mudstones are also shown. At shallow depth, the sonic-density path is a linear trend in all three wells which runs parallel with the illitic line. The sonic-density path in Bravo-Y is slightly above the illitic line (Figure 4.32b). This relatively high density may be due to the siderite content and illite dominated clay mineralogy. Changes in the siderite content may also explain the upward step in density from 2.47 g/cm3 _{to 2.54 g/cm}3 _{seen on the crossplot for Bravo-X (Figure 4.32a) but cannot}

be confirmed in the absence of mineralogical data. A similar upward step can be seen on the crossplot of Bravo-L (Figure 4.32c) however its interpretation is unclear since it is located below the top of overpressure and might be the result of unloading. There is a clear trend of increasing sonic transit time with (initially) constant density associated with the top of overpressure in Bravo-X and Bravo-Y. The same trend is masked in Bravo-L due to the presence of a high velocity layer (Figure 4.32c). This trend is typical of severe unloading due to rapid increase in pore pressure. Fluctuations and various switchbacks on the crossplot trends are attributed to lithological variation.

There are clear reversals in the density logs too, however in all three wells these are below the top of present-day overpressure (Figures 4.31 and 4.32). Since the bulk mineralogy results show consistent mineralogy with no apparent depth variation, density reversals are most easily explained as retention of higher porosity from early overpressure generated by disequilibrium compaction. Another possible interpretation of density reversal would be opening of microcracks associated with incipient hydrofracturing (e.g. Ramdhan and Goulty, 2011). However, opening of microcracks is only expected in the zone of hard overpressure where the pore pressure is approximately equal to the lithostatic stress. Also, the opening of microcracks on its own would not be able to explain the observed reduction in density, equivalent to a porosity increase of approximately 5% (Goulty et al., 2012)(Figure 4.33).

140 In overall, the following interpretation can be drawn based on the wireline logs. Early overpressure was generated by disequilibrium compaction and subsequent development of strong overpressure is due to unloading.

According to Madon (2007) the primary cause of overpressure in the Central Malay Basin is disequilibrium compaction. Similar conclusion has been reached by Hoesni (2004, 2015) who also suggested that overpressure generation by disequilibrium compaction was enhanced by clay diagenesis and associated chemical compaction. However sharp pressure ramps into high overpressure –such as those observed in these three wells— are generally considered to be the result either of gas generation or load transfer or lateral/vertical pressure transfer (Swarbrick et al., 2002). Clay diagenesis and load transfer are not viable pressure generating mechanisms in the Formation E mudstones based on the observation that illitization is already in its final stages at a depth of ~1400 m below seafloor. Furthermore, the present-day top of overpressure is at considerably greater depth and shows no correlation with clay mineralogy. The temperature is high enough (>120 o_{C) at the top of overpressure for gas generation to be a viable pressure generating}

mechanism. In addition, gas may be migrating upwards through conduits from deeper sources. This is further supported by the existence of a large gas cloud observed on seismic sections adjacent to the studied wells (Figure 4.6). Although clay diagenesis did not contribute directly to effective stress reduction, chemical compaction likely resulted in anomalously low mudstone porosities (see also 4.7.4). This might have contributed indirectly to overpressure generation by creating a more effective seal.

In addition to this, at the present day, Formation D is ~100 m shallower in Bravo-L and Bravo-X where the top of overpressure is also ~100 m shallower than in Bravo-Y (Figure 4.8), suggesting that there may be a local stratigraphic control on the depth to the top of overpressure.