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doi:10.14379/iodp.sp.353.2014

Background

Motivation for drilling in the Bay of Bengal

Pliocene–Pleistocene

A threefold motivation exists for targeting the precipitation/salinity signal recorded in sediments at the Expedition 353 drilling locations (Fig. F1). First, the Bay of Bengal/Andaman Sea and surrounding catchments are within the Earth’s strongest hydrological regime, impacting billions of people (Fig. F2); a solid understanding of the physics behind monsoonal climate change is of significant societal relevance (Nicholls et al., 2007). The net annual surface water exchange (precipitation plus runoff minus evaporation) within the Bay of Bengal and Andaman Sea during the summer monsoon is 184 × 1010 m3, dominating the winter signal of –32 × 1010 m3 for an annual average of 152 × 1010 m3 (Varkey et al., 1996). The effects of this budget are clearly evident in the surface salinity climatology (Fig. F3) (Antonov et al., 2010), indicating a well-defined strong signal that can be used to monitor changes in monsoonal precipitation via chemical, physical, and isotopic indicators for changes in precipitation, salinity, and terrestrial erosion/runoff. The strength of this runoff signal is sufficient to mute (via stratification) what would otherwise result in strong summer season productivity in response to wind-driven upwelling along the eastern Indian margin, similar to that seen in the Arabian Sea (Guptha et al., 1997; Kumar et al., 2002). Compared to other monsoon regions, the Bay of Bengal is optimal for isolating and recording the summer monsoon precipitation signal (Fig. F4).

Second, recent studies have called into question the extent to which basin-scale monsoon winds and continental precipitation are coupled over a range of timescales and space scales (Clemens et al., 2010; Clemens and Prell, 2007; Liu et al., 2006; Molnar, 2005; Ruddiman, 2006; Wang et al., 2008; Ziegler et al., 2010). Nearly all proxy records indicate strong coupling between summer monsoon winds and precipitation across the Indo-Asian monsoon subsystems at the millennial scale (Altabet et al., 2002; Cai et al., 2006; Clemens, 2005; Schulz et al., 1998; Sun et al., 2011; Wang et al., 2001); this tight coupling is likely attributed to the strong role of the winter westerlies in linking high- and low-latitude climate change. However, the all-important physical mechanisms behind these links are not fully understood; this was a primary goal of recent Integrated Ocean Drilling Program Expedition 346. Progress is also being made in understanding winter monsoon and summer monsoon linkages at the millennial timescale. For example, recent work offshore Goa, western India, shows synchronous breakdown in summer and winter monsoon airflow over the Arabian Sea during Heinrich events (Singh et al., 2011), which is in contrast to the East Asian Monsoon system that shows an asynchronous relationship between summer and winter monsoon strength at the millennial scale (Yancheva et al., 2007). Consensus does not yet exist on the extent of the coupling or the ultimate forcing of monsoon winds and precipitation at the orbital and longer timescales (An et al., 2011; Caley et al., 2011a, 2011b, 2011c; Cheng et al., 2009; Clemens and Prell, 2007; Clemens et al., 1996, 1991, 2008; Clift et al., 2008; Ruddiman, 2006; Wang et al., 2008; Ziegler et al., 2010). Some argue for a close coupling between changes in Indian and East Asian summer monsoon winds and precipitation across the entire region spanning the Arabian Sea (Ocean Drilling Program [ODP] Leg 117), the South China Sea (ODP Leg 184), and terrestrial records from the Loess Plateau. In this case, changes in the strength of summer monsoon circulation across these regions are thought to be sensitive to Northern Hemisphere sensible heating (insolation), the timing of energy release from the Southern Hemisphere Indian Ocean, and the timing of global ice volume minima (An et al., 2011; Caley et al., 2011b, 2011c; Clemens and Prell, 2003, 2007; Clemens et al., 1996, 2008). In contrast, others interpret the timing of summer monsoon circulation, on the basis of speleothem records from southeast China, as forced entirely and directly by external insolation with little or no influence from internal boundary conditions such as ice volume or Southern Hemisphere ocean–atmosphere latent heat exchange (Cheng et al., 2009; Ruddiman, 2006; Wang et al., 2008). Caballero-Gill et al. (2012) demonstrate that these contrasting interpretations are not attributable to differences in terrestrial and marine chronologies. Therefore, this lack of consensus points either to a strong deficit in our understanding of monsoon sensitivity to changes in the most basic of boundary conditions including insolation, ocean/atmosphere energy exchange, ice volume, and atmospheric greenhouse gas concentrations or to the misinterpretation of the influence of seasonality on proxy records (e.g., Fig. F4).

Lack of consensus also extends to the tectonic scale, where the timing of monsoon intensification to modern strength is also debated. Some proxy records suggest initial intensification occurred at ~7–8 Ma (e.g., Kroon et al., 1991; Prell et al., 1992), whereas others suggest a considerably earlier intensification, perhaps as early as ~22 Ma (Clift et al., 2008; Guo et al., 2002; Sun and Wang, 2005). Emergence and expansion of arid-adapted C4 flora in South Asia argues for reduced precipitation since ~8 Ma (e.g., Cerling et al., 1997; Huang et al., 2007; Quade and Cerling, 1995), whereas proxies dedicated to reconstructing seasonality suggest, instead, little variability in the monsoon over the last 10 m.y. (Dettman et al., 2001). Clift and Plumb (2008), Molnar et al. (2010), and the report from the Detailed Planning Group “Asian Monsoon and Cenozoic Tectonic History” (www.iodp.org/​doc_download/​2336-mmdpgreport) provide comprehensive overviews of these issues. More recently, Rodriguez et al. (2014) suggested that the ~8 Ma intensification inferred on the basis of increased Globigerina bulloides concentrations is an artifact of increased preservation related to uplift of the Owen Ridge at this time, resulting in enhanced preservation.

Third, recent work suggests that interpretation of the oxygen minimum zone (OMZ) signal in the northern Arabian Sea (Leg 117) may be complicated by changing oxygen content of southern-source intermediate waters (Anand et al., 2008; Caley et al., 2011c; Schmittner et al., 2007; Ziegler et al., 2010). This presents a potential complication in the interpretation of the OMZ signal as a direct response to atmospheric circulation in the core region of summer monsoon winds (i.e., oxygen drawdown in response to decay of upwelling-produced organic carbon). A multibasin, multiproxy approach is required to resolve these outstanding issues. Precipitation, salinity, and runoff indicators are not influenced by the chemistry of externally sourced intermediate and deepwater masses, offering the potential to disentangle the influences of these factors in interpreting monsoon proxy records.

The iMonsoon drilling effort, targeting the core of the monsoon precipitation signal in the Bay of Bengal (Figs. F2, F4), will directly address these outstanding issues regarding large-scale monsoonal circulation through a meridional transect approach targeting the northeast Indian margin, the Andaman Sea, and the southern Bay of Bengal (Fig. F1A–F1C). Specifically, new material from this critical region will allow us to assess (1) the relative sensitivity of the monsoon to external insolation forcing and internal climate boundary conditions, including the export of latent heat from the southern hemisphere, the extent of global ice volume, and greenhouse gas concentrations; (2) the timing and conditions under which monsoonal circulation initiated and evolved; and (3) the extent to which Indian and East Asian monsoon winds and precipitation are coupled and at what temporal and geographic scales. Finally, a detailed record of monsoon evolution is also central to testing models that link the climatic evolution of South Asia to the tectonic development of the Himalaya and rising of the Tibetan Plateau. If climatically modulated surface processes really do control the structural evolution, then well-dated climate records are needed to correlate with the increasingly well constrained ages of faulting and exhumation on shore.

Resolving these outstanding questions using the geological record is critical to providing verification targets for climate models, especially given the critical point that the vast majority of current Atmosphere-Ocean General Circulation Models (AOGCMs) used in the Intergovernmental Panel on Climate Change (IPCC) reports do not accurately simulate the spatial or intraseasonal variability of monsoon precipitation (Randall et al., 2007).

Deep time

Benthic foraminiferal diversity and assemblage composition, in conjunction with geochemical proxies indicate a stepwise increase in primary production and carbon export and an expansion of the intermediate water OMZ in the northeastern Indian Ocean since the late Oligocene (Gupta et al., 2013). The main increases in productivity started at 14 (Gupta et al., 2013) or 10 (Gupta et al., 2004) Ma and reached levels associated with present Indian monsoon conditions around 2.3 Ma. This late Miocene to Pliocene “biogenic bloom” (Farrell et al., 1995) implies important changes in nutrient cycling in the Indian Ocean and probably on a global scale, which in particular affected the silica and phosphate cycles (Dickens and Owen, 1999). The productivity increase between 10 and 8 Ma in the eastern equatorial Indian Ocean (onset of the biogenic bloom) may have been linked either to global cooling and the expansion of Antarctic ice sheets leading to a major change in deep ocean circulation and nutrient cycling, to the initiation of the Indian monsoons, or a combination of both (Gupta et al., 2004). Deep-sea benthic foraminiferal diversity in the Indian Ocean further decreased between 8 and 6 Ma and is associated with a negative δ13C shift at 3.2–2.3 and 1.6–0.9 Ma, coinciding with an increased abundance of species indicative of enhanced organic carbon flux (Singh and Gupta, 2005). Since ~2.8 Ma, roughly coeval with the onset of Northern Hemisphere glaciation, benthic foraminiferal species that are well adapted to strong seasonal fluctuations in carbon flux dominate the assemblages. This has been related to increased duration and strength of the northeast (winter) monsoon, which is accompanied by relatively low primary production in the eastern equatorial Indian Ocean (Gupta and Thomas, 2003). iMonsoon will provide new Neogene intermediate water benthic foraminiferal assemblage and δ18O and δ13C records along a meridional transect to better understand the relative contribution of monsoon-related productivity–carbon flux changes and changes in intermediate water circulation linked to high-latitude climatic events such as fluctuations in the extension of the East Antarctic Ice Sheet.

The meridional transect of iMonsoon provides a unique opportunity to investigate the nature and timing of variations in deepwater radiogenic isotope composition in response to restriction of the deepwater connection between the Pacific and Indian Oceans through the Indonesian Gateway since the mid-Miocene and to evaluate the influence of enhanced Himalayan weathering since the late Oligocene. The broad passage between the Indian and Pacific Oceans during the Paleogene must have enabled significant surface and intermediate water exchange and the possibility of deepwater flow from the Indian Ocean to the Pacific Ocean (Thomas et al., 2003). The progressive closure of the Indonesian Gateway because of the northward movement of Australia (Hall, 2002; Hall et al., 2011; Kuhnt et al., 2004) induced changes in deep and intermediate water circulation through the Indonesian Gateway, which may have resulted in a significant change in the eastern Indian Ocean Nd isotope composition during the middle Miocene (Frank et al., 2006; Gourlan et al., 2008). A second shift in eastern Indian Ocean Nd isotopes may have been related to a shift in the source area of the Indonesian Throughflow toward the North Pacific around 3.5 Ma (Cane and Molnar, 2001; Gourlan et al., 2008).

Geological setting

The eastern continental margin of India is the result of the separation of India and the Australia/Antarctica portion of Gondwanaland during the Early Cretaceous at ~130 Ma (Scotese et al., 1988; Powell et al., 1988). All the major Indian rivers draining into the Bay of Bengal are thought to be associated with graben features resulting from the rifting of India from Antarctica as well as subsequent Indian plate motion. The Mahanadi Graben appears to have a continuation in Prydz Bay, Antarctica, known as the Lambert Graben (Federov et al., 1982). The 2000 m isobaths of the northeast Indian continental margin and the Lambert Graben of East Antarctica (Prydz Bay) are closely matched, supporting the inferred alignment of India and Antarctica prior to rifting (Subrahmanyam et al., 2008). The Mananadi Basin includes both onshore and offshore regions of northeast India. The offshore portion extends from ~170 to 240 km off the coast. Sediments are dominated by input from the Mahanadi River drainage basin, which occupies 141,589 km2 of catchment (extending from ~80° to 85°E and ~19° to 24°N) and currently supplies ~7.1 × 109 kg of sediments per year to the offshore basin (Fig. F5) (Subrahmanyam et al., 2008). Analysis of Bay of Bengal surface sediments indicates that the foraminiferal lysocline, the depth delimiting well preserved from noticeably dissolved assemblages, shoals significantly from south to north. The foraminiferal lysocline rises from 3800 to 3300 m between 0° and 7°N (about Site N90E-2C) then systematically shoals to ~2000 m at ~20°N (Indian margin sites) (Cullen and Prell, 1984).

The Ninetyeast Ridge (NER) is an aseismic volcanic ridge spanning from ~31°S to ~10°N, where it is buried beneath Bengal Fan sediments. The NER is thought to have formed by age-progressive hotspot volcanism from plume sources currently beneath the Kerguelen Plateau (Royer et al., 1991; Sager et al., 2010). The ridge top rises to a height of ~3.5 km above the surrounding abyssal plain with depths as shallow as ~2000 meters below sea level (mbsl). Site N90E-2C is located at ~5°N at 2963 mbsl. This location provides for good preservation of carbonate microfossils, given that the foraminiferal lysocline in this region is close to 3300 m (Cullen and Prell, 1984).

The Andaman Sea is situated between the Andaman Islands and the Malaya Peninsula (Fig. F1A). The Andaman-Sumatra island arc system results from the oblique subduction of the Indo-Australian plate beneath the Eurasian plate (Singh et al., 2013). Stretching and rifting of the overriding plate in the early Miocene (~25 Ma) has resulted in two distinct plates (Sunda and Burma) separated by an active spreading center (Curray, 1991, 2005) located in the deepest portion of the Andaman Sea. An accretionary wedge complex scraped off the subducting slab lies west of the spreading center, forming a series of shallower basins associated with back-thrust faulting within the accreted sediments (Fig. F6). The Andaman Sea drilling sites are within the Nicobar-Andaman Basin, bounded on either side by the Diligent and Eastern Margin Faults. Terrigenous sediment supply to the Andaman Sea is dominantly from the Irrawaddy and Salween Rivers (Colin et al., 1999, 2006). Analysis of Andaman Sea surface sediments indicates that foraminifers are abundant and well preserved shallower than ~1800 mbsl (>100,000 individuals/gram) and decrease to <100 individuals/gram deeper than 3000 mbsl (Frerichs, 1971).

Atmospheric and oceanographic circulation

The Indian summer monsoon is characterized by low atmospheric pressure over the Indo-Asian continent (Indo-Asian Low) relative to high atmospheric pressure over the southern subtropical Indian Ocean (Mascarene High). The resulting pressure gradient leads to large-scale displacement of the Intertropical Convergence Zone (ITCZ) and the cross-equatorial flow of low-level winds carrying moisture that is ultimately released over South Asia, the Bay of Bengal, and southeast China (Hastenrath and Greischar, 1993; Liu et al., 1994; Loschnigg and Webster, 2000; Webster, 1987a, 1987b, 1994; Webster et al., 1998). Modern meteorological observations and moisture transport budgets (Fig. F7) quantitatively show that the Southern Hemisphere Indian Ocean is the dominant source of moisture (latent heat) to the Indian and East Asian summer monsoons during June, July, and August (JJA) (Bosilovich and Schubert, 2002; Ding and Chan, 2005; Ding et al., 2004; Emile-Geay et al., 2003; Liu and Tang, 2004, 2005; Park et al., 2007; Simmonds et al., 1999; Wajsowicz and Schopf, 2001; Xie and Arkin, 1997; Zhu and Newell, 1998). The Arabian Sea is a very minor moisture source (evaporation > precipitation), whereas the Bay of Bengal/Andaman Sea, India, the South China Sea, and southeast China are all moisture sinks (precipitation > evaporation).

A total of 12 major rivers (Fig. F5, F8) feed the Bay of Bengal/Andaman Sea (Ganga, Brahmaputra, Meghna, Damodar, Mahanadi, Godavari, Krishna, Irrawaddy, Salween, Penner, Kavery, and Mahaweli Rivers), discharging in total 943 × 109 m3 of water during the summer monsoon months (JJA) (Varkey et al., 1996). Annual rainfall within and surrounding the Bay of Bengal is dominated by precipitation during the summer monsoon months (JJA) with the exception of the Madras Basin in the southernmost peninsular India, where rainfall peaks in November (Fig. F9). The dominance of the summer (JJA) precipitation signal is reflected in the Bay of Bengal surface salinity patterns (Fig. F3), which reach their lowest values in August and September, spanning salinities of 20–34 over both seasonal (summer–winter) and spatial (north–south) dimensions.

Primary surface ocean currents (Schott and McCreary, 2001; Schott et al., 2009) reflect the seasonal wind forcing in both the eastern Arabian Sea and the Bay of Bengal (Fig. F10). The West Indian Coastal Current (WICC) flows south during the summer monsoon, connecting with the Southwest Monsoon Current (SMC) that carries high-salinity waters eastward around the tip of India and Sri Lanka into the southern Bay of Bengal at a rate of 8.4 Sverdrup (Sv; 106 m3/s). This influx of high-salinity water is reflected in the July, August, and September salinity patterns of the southern Bay of Bengal (Fig. F3) and is successfully modeled as a passive tracer in mixed-layer ocean models (Jensen, 2001, 2003). Southwest summer monsoon winds in the Bay of Bengal also drive the northward-flowing East Indian Coastal Current (EICC). During the winter monsoon, northeast winds drive all these surface currents in the opposite directions, transporting 11 Sv of water toward the eastern Arabian Sea.

The proposed drilling plan is designed to take advantage of these strongly seasonal patterns to reconstruct changes in summer monsoon circulation by reconstructing the meridional precipitation/salinity gradients as well as erosion and runoff from proximal drainage basins. An array of drilling locations is proposed, including the Mahanadi Basin, off the northeast Indian margin (BB sites), the Andaman Sea (AA sites), and the southernmost Bay of Bengal (Site N90E-2C) (Fig. F1A–F1C). Salinity on the Indian margin, northwest Bay of Bengal, reaches a minimum of ~22 in September (Fig. F3); this is a lagged response to JJA rainfall over the Bay of Bengal and the surrounding drainage basins. Salinity at this location reaches a maximum of ~34 during the spring months. The Andaman Sea sites, situated between the modern 32 and 33 isohalines, monitor drainage from the Irrawaddy and Salween Rivers. Site N90E-2C (ODP Site 758) is closely pinned to the 34 isohaline year round, anchoring the southern end of the modern salinity gradient at near open-ocean values. Although this site does not currently experience significant seasonal salinity variability, it does record large-scale changes in precipitation and runoff at the millennial, orbital, glacial–interglacial, and tectonic scales as discussed below. The full meridional transect (spanning the Indian margin, Andaman Sea, and northern NER) has a modern salinity range of 12, equivalent to an ~2‰ surface water δ18O signal.

Changes in the meridional salinity gradient will provide robust means of tracking changes in summer monsoon precipitation. Furthermore, the signal should not be influenced to a large degree by temperature gradients, which are small compared to the salinity gradients. Maximum temperature seasonality is ~3°C on the northwest Indian margin (equivalent to <0.8‰ surface water δ18O signal) and considerably less at the other sites (Fig. F10) (Locarnini et al., 2010). This is largely because of the cooling influence of cloud cover and precipitation during the summer, limiting the amount of sensible surface heating.

Terrestrial runoff products are also of great utility in assessing linkages between monsoon circulation, chemical weathering, and transport at timescales from millennial to tectonic. These topics are recognized by the community as being of considerable importance (Clift and Plumb, 2008; Wang et al., 2005). Changes in monsoon strength are well documented at ~23, 15, 8–7, and 2.75 Ma (Clift and Plumb, 2008). The iMonsoon targets will allow measurement of the consequent impact on weathering rates and transport of particulate materials to the ocean basins in a variety of settings both proximal and distal relative to river inputs.

Water masses and circulation

Comprehensive descriptions of eastern Indian Ocean regional oceanography are provided in Wyrtki (1971), Mantyla and Reid (1995), Tomczak and Godfrey (2003), and Schott et al. (2009), from which we briefly summarize descriptions of water masses and circulation patterns relevant for Bay of Bengal drilling in the depth range between 1100 and 3000 m targeted for Expedition 353.

Indian Deep Water (IDW) occupies the depth range from 3800 to ~1500 m within the equatorial and northern Indian Ocean (Fig. F11). IDW in the eastern Indian Ocean is characterized by high salinities reaching salinity maxima of 34.8 in the southwestern Indian Ocean and 34.75 in the southeastern Indian Ocean, where the IDW upper limit rises to 500 m (Tomczak and Godfrey, 2003). IDW temperature, salinity, and oxygen properties in the high-salinity core are virtually identical with those of North Atlantic Deep Water (NADW) in the Atlantic sector of the Southern Ocean, indicating that IDW is mainly of NADW origin and not originally formed in the Southern Ocean, as is the Antarctic Bottom Water (AABW) that occupies the Indian Ocean deeper than 3800 m (Tomczak and Godfrey, 2003). The flow of IDW is northward and concentrated along western boundaries of the African margin and NER as indicated by the World Ocean Circulation Experiment (WOCE) I08I09 oxygen, silicate, and temperature profiles (Fig. F12). IDW further penetrates northward into the Northern Hemisphere, is modified by mixing with thermocline water from above and upwelling of AABW from below, and spreads into the Arabian Sea and the Bay of Bengal.

Two water masses occupy the thermocline of the Indian Ocean: Indian Central Water (ICW) and Indonesian Throughflow Water (ITW) or Australasian Mediterranean Water (Fig. F13). ICW originates from downwelling in the subtropical convergence south of 30°S and ITW is derived from North Pacific Intermediate Water, strongly modified during its passage through the Indonesian archipelago. There is no formation of thermocline water in the Bay of Bengal, and its thermocline water masses to 1500 m water depth are derived from ICW and ITW. Transfer of ICW to the northern Indian Ocean is accompanied by a rapid decrease in oxygen content, indicating aging along the path. The lowest oxygen values occur in the Bay of Bengal, which contains the oldest ICW. The strong oxygen decrease in the northern Indian Ocean can be explained by restriction of the transfer of ICW to the southwest monsoon season, resulting in a small annual net transfer rate. ITW also contributes to the renewal of thermocline water in the northern Indian Ocean, resulting in significant freshening of the ICW along its path into the Bay of Bengal. Further freshening is observed in the Bay of Bengal near 90°E, resulting from ITW advection directly from its outflow area into the tropical eastern Indian Ocean. The variability and evolution of thermocline circulation in the Bay of Bengal are strongly dependent on monsoonal forcing; however, the present extremely low oxygen levels indicate a very low renewal rate for the thermocline waters of the Bay of Bengal.

The uppermost 100 m of the eastern Indian Ocean in the Bay of Bengal consists of a low-salinity water mass derived from river runoff from India and Indochina, the Bay of Bengal Water (BBW), with surface salinity strongly fluctuating with seasons but remaining below 33 throughout the year. The lower boundary to the ICW is characterized by a strong halocline. Its southward extension is highest during October–December when it reaches the area along the western Indian coast and is lowest during April–June before the summer monsoon leads to a new expansion of the BBW surface water mass.

Site survey data

Supporting site survey data for Expedition 353 are archived at the IODP Site Survey Data Bank.

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