marine and petroleum geology - usal · expedition-01 (nghp-01) site 17 (10 45.19120n, 93 6.73650e)...

lable at ScienceDirect

Marine and Petroleum Geology xxx (2014) 1e12

Contents lists avai

Marine and Petroleum Geology

journal homepage: www.elsevier .com/locate/marpetgeo

Research paper

A late MioceneeEarly Pliocene biogenic silica crash in the AndamanSea and Bay of Bengal

T. Cawthern a, *, 1, J.E. Johnson a, L. Giosan b, J.A. Flores c, K. Rose d, E. Solomon e

a Department of Earth Sciences, University of New Hampshire, Durham, NH 03824, USAb Department of Geology and Geophysics, Woods Hole Oceanographic Institute, Woods Hole, MA 02543, USAc Grupo de Geociencias Oceanicas, Universidad de Salamanca, Salamanca, 37008, Spaind National Energy Technology Laboratory, U.S. Department of Energy, Albany, OR 97321, USAe School of Oceanography, University of Washington, Seattle, WA 98195-7940, USA

a r t i c l e i n f o

Article history:Received 2 December 2013Received in revised form22 July 2014Accepted 26 July 2014Available online xxx

Keywords:Indian OceanBay of BengalIndonesian ThroughflowNGHP-01Biosilica

* Corresponding author. Department of Geographyden Ave., Salisbury University, Salisbury, MD 21801, U

E-mail address: [email protected] (T. Caw1 Now at Department of Geography and Geoscienc

isbury, MD 21801, USA.

http://dx.doi.org/10.1016/j.marpetgeo.2014.07.0260264-8172/© 2014 Elsevier Ltd. All rights reserved.

Please cite this article in press as: Cawthern,Marine and Petroleum Geology (2014), http

a b s t r a c t

Variations in the mass accumulation rate of biogenic silica (BSi) in continental margin sediments can beused to reconstruct relative changes in productivity through time in these settings. In the northern Bay ofBengal a lack of long sedimentary records has historically precluded this type of reconstruction. Theacquisition of 21 new long sedimentary records during the 2006 Indian National Gas Hydrate Program(NGHP) Expedition-01 has made it possible for the first time to reconstruct paleoproductivity in thisimportant region of the world that is dominated by intense changes in the geological and biologicalfluxes largely driven by tectonic and climate related mechanisms. In the research presented here, fluc-tuations in the mass accumulation rate of biogenic opal during the past ~9.4 Myrs are reconstructedusing continental margin sediment cores from the Andaman Sea (Site NGHP-01-17A) and the northernBay of Bengal (Site NGHP-01-19). Within these records, a biogenic silica crash is recorded at ~6 Ma and isconsistent with previous geotectonic, geochemical and paleontological studies of the southern IndianOcean and Pacific Ocean that suggest connectivity, and thus exchange of nutrient-rich water masses,between the eastern tropical Indian Ocean and western tropical Pacific Ocean was diminished as a resultof the tectonic restriction of the northerly sector of the Indonesian Throughflow (ITF). The biogenic silicacrash at Sites 17 and 19 is consistent with a decrease in surface water productivity that may have beendriven by the reduction of nutrient-rich Pacific waters delivered to the Andaman Sea and Bay of Bengalvia the northerly route of the ITF. Following the BSi crash at ~6 Ma, subsequent recovery of the BSi massaccumulation rates at Sites 17 and 19 occurred and was perhaps renewed by an enhanced supply ofnutrient-rich freshwater from the nearby Irrawaddy and Mahanadi Rivers, which could have occurredduring a documented increase in the intensity of the Indian monsoon at ~5 Ma. Although recovery isnoted at both core locations, biogenic silica mass accumulation rates did not fully recover in theAndaman Sea. This could be explained by the restricted nature of the Andaman basin and its more distallocation from a major source of nutrient-rich freshwater.

© 2014 Elsevier Ltd. All rights reserved.

1. Introduction

Several DSDP and ODP expeditions in the Pacific and north-ecentral Indian Oceans have documented significant increases inthe mass accumulation rates (MAR) of biogenic sediments during

and Geosciences, 1101 Cam-SA. Tel.: þ1 410 677 6514.thern).es, Salisbury University, Sal-

T., et al., A late MioceneeEarly://dx.doi.org/10.1016/j.marpe

the late MioceneePliocene (e.g., Leinen, 1979; Woodruff, 1985;Farrell et al., 1995; Dickens and Owen, 1996). These changes markimportant shifts in planktonic and benthic productivity, and areespecially well developed in oceanographic divergence zones(Peterson et al., 1992; Berger et al., 1993; Dickens and Owen, 1994).The time-coincident nature of biogenic sediment increases in thePacific and Indian Oceans resulted in the development of the“biogenic bloom” hypothesis that suggests enhanced nutrient up-welling to surface waters in divergence zones led to an increase inthe downward flux of biogenic sediments and an expansion of theoxygen minimum zone, enhancing preservation (e.g., Dickens andOwen, 1999). Sediment deposited between 9.0 and 3.5 Ma in

Pliocene biogenic silica crash in the Andaman Sea and Bay of Bengal,tgeo.2014.07.026

Delta:1_-

Delta:1_given name

Delta:1_surname

Delta:1_given name

Delta:1_surname

Delta:1_given name

mailto:[email protected]

www.sciencedirect.com/science/journal/02648172

http://www.elsevier.com/locate/marpetgeo

http://dx.doi.org/10.1016/j.marpetgeo.2014.07.026



T. Cawthern et al. / Marine and Petroleum Geology xxx (2014) 1e122

these ocean basins contain planktonic and benthic fauna that areindicative of high productivity and organic content and redox sig-natures indicative of low dissolved O2 (Dickens and Owen, 1999).

Two potential mechanisms have been proposed to explain theorigin of the late MioceneePliocene biogenic bloom: 1) elevatednutrient supply to the ocean basins via increased continentalweathering or 2) a redistribution of nutrients within the oceansresulting from a change in the strength of deep water formation(Dickens and Owen,1999) or an opening and closing of passages forocean circulation (e.g., Keller and Barron,1983; Cortese et al., 2004).Although both scenarios are plausible, a better understanding ofthe late MioceneePliocene biogenic bloom may be achieved as thespatial distribution of its occurrence and its timing are betterconstrained globally.

Although plausible explanations for the late MioceneePliocenebiogenic bloom have been proposed and discussed, the subse-quent cessation of the bloom has received limited discussion(Keller and Barron, 1983; Farrell et al., 1995; Dickens and Owen,1999). In this paper, we document the biogenic silica mass accu-mulation rate (MAR) of sediments located in the Andamanaccretionary wedge at National Gas Hydrate Program of IndiaExpedition-01 (NGHP-01) Site 17 (10� 45.19120N, 93� 6.73650E)and in the northwestern Bay of Bengal at NGHP-01 Site 19 (18�

58.65680N, 85� 39.52020E) and discuss the mechanisms that mayhave caused an observed abrupt decrease or “crash” in BSi at~6 Ma at both sites.

Figure 1. Regional map of the NE Indian Ocean region showing core locations from the 200region (yellow dots), the location of major rivers entering the NE Indian Ocean and Andamanlines are deep water masses, coarse dashed lines are intermediate waters, and fine dashed linlegend, the reader is referred to the web version of this article.)

Please cite this article in press as: Cawthern, T., et al., A late MioceneeEarlyMarine and Petroleum Geology (2014), http://dx.doi.org/10.1016/j.marpe

2. Geologic setting

The seafloor sedimentary record in the Bay of Bengal in thenorthern Indian Ocean is dominated by the Bengal and Nicobarsubmarine fans, which have received terrigenous sediments fromthe erosion of the Himalayan Mountains via transport through theGangeseBrahmaputra drainage systems during the last 15 millionyears (Curray et al., 2003; Bastia et al., 2010, Fig.1). The extent of theBengal Fan is limited to the west by the continental shelf and slopeof peninsular India, which formed as a passive margin during thebreak-up of Gondwana (Powell et al., 1988; Curray et al., 2003;Bastia et al., 2010) and has received terrigenous sediments fromthe modern Mahanadi, Krishna, Godavari and Kaveri Rivers andpredecessor rivers (Fig. 1). Site 19 is located on the slope, above themodern-day depocenter of the Bengal Fan, on the easternmargin ofpeninsular India, just south of the Mahanadi River discharge. To theeast, across the Bay of Bengal, much of the Bengal Fan has beensubducted or accreted beneath the Sunda Subduction Zone (Bowleset al., 1978; Curray et al., 2003). As the aseismic Ninety East Ridgecontinues to approach the subduction zone it also inhibits BengalFan sedimentation, isolating the more westerly and previouslydeposited fan sediments, now known as the Nicobar Fan, from themain Bengal Fan lobe (Bowles et al., 1978; Curray et al., 2003).Farther to the east, the semi-enclosed Andaman Sea is located in aback-arc basin formed by the oblique subduction of the Indian Platebeneath the Sunda Plate (Rodolfo, 1969; Karig et al., 1980; Pal et al.,

6 NGHP-01 Expedition (red dots and stars) and existing ODP and DSDP core sites in theSea, generalized ocean circulation, and regional volcanic centers (pink dots). Solid bluees are shallow water masses. (For interpretation of the references to color in this figure


T. Cawthern et al. / Marine and Petroleum Geology xxx (2014) 1e12 3

2003; Raju et al., 2007). Backarc extension and the concomitantformation of the Andaman Sea began between 11 and 4 Ma(Rodolfo, 1969; Karig et al., 1980; Pal et al., 2003; Raju et al., 2007).Since ~4 Ma, active backarc spreading in the Andaman Sea hasresulted in total extension of ~460 km (Raju et al., 2004; Khan andChakraborty, 2005). Within the Andaman Sea, terrigenous sedi-ments from the Irrawaddy River are deposited on the shelf andslope, north of the Andaman back-arc spreading center (Rodolfo,1975). Transportation of fine-grained sediments are restricted tothe nearshore region by monsoonal surface currents (Rodolfo,1975; Babu et al., 2010; Tripathy et al., 2011). Site 17 is located inthe Andaman accretionary wedge, east of Little Andaman Island.Both Sites 17 and 19 are located at depths above the reach ofsediment gravity flows associated with the Bengal Fan and thuspreserve nearly continuous records of pelagic/hemipelagic sedi-mentation that are ideal for paleoceanographic study.

3. Site and core stratigraphy

This study focuses on the stratigraphic records recovered fromtwo marine sedimentary records from the northeast Indian Ocean.Site 17 was drilled in the Andaman accretionary wedge, east ofLittle Andaman Island in the Andaman Sea (Fig. 1). This site islocated in 1344 m of water, far from major sources of terrestrialclastic influx, and preserves predominantly carbonate oozes withintervals rich in biogenic silica, within fine-grained (clay-sized)terrigenous material, punctuated by volcanic ashes (Collett et al.,2008; Rose et al., 2014). Site 19 was drilled in the Mahanadi Ba-sin, located along the eastern passive continental margin of Indiaproximal to the mouth of the Mahanadi River. Site 19 is located in1422 m of water and is the most heterolithic core from NGHPExpedition-01 (Collett et al., 2008). The hemipelagic record at Site19 consists of varying amounts of silt and clay sized minerals,volcanic glass, pyrite, authigenic carbonate and aragonite, plantdebris, nannofossils, biogenic silica and foraminifera (Collett et al.,2008; Phillips et al., 2014).

Shipboard smear slide and coarse fraction analyses of the sedi-ments at Site 17 indicate a shift in the marine microfossils occurredsince the lateMiocene (Collett et al., 2008). Specifically, biosiliceousand carbonate marine fossils dominate the lower half of the core,whereas carbonate-bearing marine fossils dominate the upper halfof the core (Fig. 2). The record recovered from Site 19 also followsthe same pattern in biogenic components (Fig. 2).

The ages of the records at Sites 17 and 19were determined usinga detailed nannofossil age model (Fig. 3). Calcareous nannofossilevents were identified at Sites 17 and 19 spanning an interval fromthe Pleistocene to themiddle Miocene (Flores et al., 2014), based onstandard biozones of Martini (1971) and Okada and Bukry (1980).The biochronology was developed following Raffi and Flores (1995)and Raffi et al. (2006) using astronomically calibrated ages frommagnetic reversals and reference isotope stratigraphies.

4. Oceanographic setting

In the Indian Ocean, intermediate and deep water massesoriginate from the Atlantic and Southern Oceans (Fig. 1). Interme-diate waters (~1200e3800 m) are comprised of a mixture of Ant-arctic Bottom Water (AABW) and North Atlantic Deep Water(NADW) (also known as Circumpolar DeepWater, CPDW) (Vincent,1974; Kawagata et al., 2006). Deep waters (>3800 m) are formedfrom northward-flowing cold, well-oxidized Antarctic BottomWater (AABW) (Peterson, 1984; Kawagata et al., 2006). Within theBay of Bengal, the upper ~1000 m of the water column is bathed inNorth Indian High Salinity Intermediate Water (NIHSIW) and aseasonally-influenced low salinity surface water mass. The NIHSIW


enters the Bay of Bengal from the Arabian Sea as the MonsoonCurrent, but is ultimately derived from the Red Sea and Persian Gulf(Fig.1; Murthy et al., 1992; Naqvi et al., 1994; Kawagata et al., 2006).The NIHSIW and low salinity surface water masses in the Bay ofBengal are highly dependent on the intensity of the summer andwinter monsoons (Varkey et al., 1996). Differential heating of theocean during the winter monsoon, compared to the Indian sub-continent results in the establishment of a high-pressure systemover the land that leads to an intensification of the northerly winds,and causes decreased precipitation over the land and an increase insea surface temperature in the Bay of Bengal (Varkey et al., 1996). Incontrast, the summer monsoon results in the intensification ofsoutherly winds that forces moisture-laden air originated from theBay of Bengal over the HimalayaeTibetan mountain range, andcauses increased rainfall and freshwater discharge into the Bay ofBengal (Varkey et al., 1996). The major implication of the Asianmonsoon and its intensity through time is the effect it has on theseasonal winds, precipitation, and run-off patterns, which affectthe volume and composition of run-off entering the marginal seas.Increased precipitation and river runoff during the summermonsoon restricts vertical mixing and upwelling of deep nutrientsto the photic zone proximal to the coast (e.g., Madhupratap et al.,2003; Prassana Kumar et al., 2002) in the Bay of Bengal andAndaman Sea. Another important influence of the monsoons is theseasonal variations they induce in temperature and density in theBay of Bengal that results in very different circulation patterns. Inparticular, during the winter monsoon, surface water is transportedclockwise in the Bay of Bengal, whereas during the summermonsoon, the direction of transport is reversed (Varkey et al., 1996).Upwelling-induced productivity in the Bay of Bengal and AndamanSea regions is primarily located away from the continental margins,where the wedge of freshwater influx is thinner and wind stress ishigh, resulting in the formation of localized mesoscale cycloniceddies (Wyrtki, 1961; Buranapratheprat et al., 2010; Rizal et al.,2012).

Deep, intermediate, and surface waters in the Andaman Sea arederived from freshwater runoff and the intermixing of watermasses that originate from the Indian and Pacific Oceans (Ibrahimand Yanagi, 2006; Amiruddin et al., 2011). The AndamaneNicobarIslands (located just west of Site 17) isolate most of the AndamanSea deepwaters from the Bay of Bengal deepwaters in the northernsector of the Andaman Basin, but these two water masses intermixin the central Andaman Sea through the Preparis and Ten DegreeChannel, and in the south through the Great Channel (Fig. 1; Varkeyet al., 1996; Dutta et al., 2007). In contrast, intermediate and surfacewaters from the Bay of Bengal freely intermix with those of theAndaman Sea. Surface waters in the Andaman Sea are stronglyaffected by seasonal variations in the monsoons, particularly as aresult of increased discharge of freshwater from the Irrawaddy, andBrahmaputra and Ganges Rivers (Fig. 1; Wyrtki, 1961; Rizal et al.,2012).

In addition to the seasonally-induced changes to the surfacewater within the Andaman Sea, studies from other ocean basins(i.e., Pacific) suggest that temporal variations in the connectivity ofIndian and Pacific Ocean surface waters via the IndonesianThroughflow (ITF) has affected the productivity of carbonate (i.e.,foraminifera) and biosiliceous plankton (i.e., diatoms and radio-larians) (Keller, 1985; Kennett et al., 1985; Romine and Lombari,1985). The flow of Pacific Ocean waters westward into the IndianOcean through the Indonesian archipelago andmarginal seaways isthe result of a pressure gradient that is established between the twoocean basins due to a difference in sea-surface elevation betweenthe two ocean basins. Today, the ITF transports relatively cool, freshnorthern Pacific surface water into the Indian Ocean (Gordon andFine, 1996; Cane and Molnar, 2001; Sengupta et al., 2013) through


Figure 3. A) Age models (see Flores et al., 2014) at NGHP-01 Sites 17 and 19, showing the age and occurrence of unconformities within these records and the average sedimentationrates above and below the unconformities. B) Sedimentation rates at Sites 17 and 19 between discrete age control points.

Figure 2. Generalized core stratigraphy at Sites 17 and 19 (Collett et al., 2008).


Please cite this article in press as: Cawthern, T., et al., A late MioceneeEarly Pliocene biogenic silica crash in the Andaman Sea and Bay of Bengal,Marine and Petroleum Geology (2014), http://dx.doi.org/10.1016/j.marpetgeo.2014.07.026

Figure 4. Time-integrated silica dissolution example for a sample from the northeastIndian Ocean. The first two points target biologic SiO2. The last four points representthe slower dissolution of mineral SiO2. The slope-break in the data occurs because ofthe different dissolution rates of amorphous and mineral SiO2 components. The y-intercept of the linear regression through the last four data points is the concentrationof BSi in this sample.


numerous channels located between islands of the Indonesian ar-chipelago and the surroundingmarginal seaways, such as the SouthChina Sea (Gordon et al., 1996, 2012; Fang et al., 2005, 2009; Tozukaet al., 2007). However, because of the narrow nature of the pas-sageways that connect these two oceans and active tectonismthroughout the region, uplift and/or subsidence of the seafloor(Nishimura and Suparka, 1997; Martin and Scher, 2006; Rai andSingh, 2012), as well as eustatic sea-level falls during the mid-lateMiocene and present (Haq et al., 1987) may have periodicallyrestricted deep water mass flow between these basins or re-routedflow paths (e.g., Gordon and Fine, 1996; Fang et al., 2010; Susantoet al., 2010, 2012), potentially limiting the exchange of nutrientsbetween the western Pacific and eastern Indian Oceans (Gourlanet al., 2007; Jiang et al., 2007). Geotectonic (e.g., Nishimura andSuparka, 1997), paleontological (e.g., Kennett et al., 1985), andgeochemical (e.g., Martin and Scher, 2006) datasets from thesouthern Indian and western Pacific Oceans have revealed theprogressive restriction of the ITF from early to late Miocene(~23.7 Mae5.3 Ma). Deep water flow through the ITF is believed tohave been more or less closed and/or re-routed by ~6 Ma (Kennettet al., 1985; Hall, 1996, 2002; Nishimura and Suparka, 1997). Morerecent studies show that deep water mass exchange through theITF was periodically re-established after ~6 Ma, with completeclosure or re-routing of the deep waters being more recent, around~4e3Ma (late Pliocene) (e.g., Cane andMolnar, 2001; Gourlan et al.,2007).

The World Ocean Atlas 09 indicates that the modern distribu-tion of dissolved nutrients for the Pacific Ocean is higher than thatof the Indian ocean (Garcia et al., 2010). More specifically, the Pa-cific Ocean contains higher concentrations of silicate, phosphate,and nitrate with depth as well as for surface water masses(Buranapratheprat et al., 2010; Garcia et al., 2010). Concentrationsof these nutrients are highest in the northern Pacific Ocean, whereITF waters originate in the modern ocean (Cane and Molnar, 2001).Prior to ~4e3 Ma, the source of ITF waters was from the warm,comparatively less nutrient-rich South Pacific (Cane and Molnar,2001).

5. Methods

The mass accumulation of biogenic silica (BSi) in marine sedi-ments is a function of the surface water plankton productivity inthe ocean and often correlates with the flux of marine organiccarbon and calcium carbonate (from plankton) to the seafloor (e.g.,Calvert, 1966; Lyle et al., 1988). Shipboard smear-slide and coarse-fraction analyses of the sediments at Sites 17 and 19 revealed adecrease in the relative abundance of biosiliceous marine organ-isms from the base to the top of the cores in both records (Collettet al., 2008) and it is this change that is the focus of this study.Preservation of biosiliceous sediments within these records is goodthroughout the timeframe studied based on the quality of silicatests examined in smear slides (Collett et al., 2008). Bulk sedimentsamples fromNGHP-01 Sites 17 and 19were initially dried and thenpowdered using a mortar and pestle. Powdered samples wereplaced in individually labeled sample vials and stored in a desic-cator until biogenic silica analyses were made. These same sampleswere also analyzed for CHN elemental abundance as well as bulkCorg isotopic analysis (Johnson et al., 2014).

Biogenic silica (BSi) was determined on 68 bulk sedimentsamples from Site 17 and 56 samples from Site 19 by extracting BSifrom ~25 mg of sample material using a modified version of thewet alkaline technique of DeMaster (1981) and Strickland andParsons (1972), which are outlined in the Stanford University Sta-ble Isotope Laboratory Standard Operating Procedure Manual(Mucciarone, 2003). This procedure uses a 0.1 M NaOH solution at a


constant temperature of 85 �C to leach silica from the sample(Mucciarone, 2003). Sequential sampling of each sample was per-formed at 30 min and at each hour from hours 1e5 in order toestablish a robust BSi leaching curve. After the allotted samplinginterval was complete, a molybdate reagent and a reducing solutionwere added to turn each sample a hue of blue that corresponds tothe concentration of silica (deeper blue for greater silica content).Standards were prepared to cover the range of sample concentra-tions. Standards and samples were then analyzed on a Milton RoySpectronic 501/601 spectrophotometer at a wavelength of 812 nm.Absorbance at this wavelength varies from 3.000 to 0.000 withanalytical precision of ±0.001. The absorbance value of each samplewas measured 3e4 times, or until two consecutive readings indi-cated the absorbance of the sample had stabilized. The average ofthese two analyses was then plotted relative to the aliquot sam-pling interval (time). Each individual sample thus has a corre-sponding plot containing six absorbance analyses, one at 30 minand five for each of the successive hours (1e5) plotted relative totime (Fig. 4). In order to determine the total weight percent BSi, aleast-squares regression analysis was performed on the increase inBSi extracted versus time. Early extractions (hours ~0e2) target thedissolution of amorphous silica (BSi or volcanic glass), whereaslater ones target Si from minerals (i.e., clay, feldspars, quartz). It isnoteworthy that all of the samples analyzed for BSi in this studywere from non-volcanic ash bearing intervals in both the Site 17and Site 19 records. Therefore, the y-intercept of the extrapolatedbest-fit trendline from the least-squares regression is representa-tive of the biogenic silica concentration for an individual sample(Fig. 4). The best-fit linear regressions for all samples equaled orexceeded R2¼ 0.9 (Fig. 4). In order to convert the BSi absorbance foreach sample to a concentration, these values were compared to thestandard curve of absorbance versus concentration. The followingsilica standard concentrations were prepared by diluting a1000 ppm SiO2 (~35,606.2 mM Si) standard solution (Ricca Chemi-cal) with Milli-Q water (18.2 MU * cm at 25 �C): 0 mM Si; 125 mM Si;463 mM Si; 908 mM Si; 1193 mM Si. The standard curve of absorption(R2 > 0.999) was run at the start and end of each sampling sessionto account for potential instrument drift. High (1193 mMSi) and low(125 mM Si) standards, and a sample of Milli-Q water were alsoanalyzed after each time-specific (i.e., 30 min, 1 h, 2 h, etc.) suite of


Table 1Biogenic silica MAR at NGHP-01 sites 17 and 19.

Site 17 17 17 17 17 17 17 17 17 17 17 17 17

Sample 1H2 40-42 2H4 40-42 3H7 40-42 5H2 38-40 6H3 40-42 7H7 40-42 8H7 40-42 10H6 40-42 12H1 38-40 13X2 40-42 14X3 40-42 16X2 40-42 16X6 40-42

Depth (mbsf) 2.40 9.40 22.90 34.88 45.84 61.40 70.40 88.40 99.88 110.90 121.90 139.60 145.29BSi (wt %) 0.20 0.15 0.16 0.15 0.22 0.16 0.18 0.19 0.26 0.14 0.19 0.21 0.19Age (Ma) 0.03 0.12 0.29 0.44 0.58 0.78 0.89 1.17 1.40 1.70 2.45 2.52 2.54BSi MAR

(g/cm2 * kyr)0.01 0.01 0.01 0.01 0.02 0.01 0.01 0.01 0.01 0.00 0.06 0.06 0.06

Site 17 17 17 17 17 17 17 17 17 17 17 17 17

Sample 18X2 41-43 18X4 41-43 18X5 41-43 19X2 40-42 19X4 41-43 19X6 40-42 21X2 40-42 22X1 38-40 22X6 40-42 24X4 13-15 25X5 40-42 27X1 42-44 28X1 40-42


(g/cm2 * kyr)0.05 0.07 0.07 0.00 0.00 0.00 0.00 0.02 0.02 0.02 0.04 0.06 0.13

Site 17 17 17 17 17 17 17 17 17 17 17 17 17



(g/cm2 * kyr)0.10 0.11 0.10 0.15 0.16 0.17 0.16 0.20 0.33 0.32 0.26 0.15 0.18

Site 17 17 17 17 17 17 17 17 17 17 17 17 17



(g/cm2 * kyr)0.19 0.14 0.13 0.20 0.15 0.16 0.22 0.16 0.24 0.17 0.19 0.23 0.63

Site 17 17 17 17 17 17 17 17 17 17 17 17 17



(g/cm2 * kyr)0.49 0.46 0.66 0.55 0.56 0.56 0.61 0.58 0.86 0.74 0.70 0.61 0.19

Site 17 17 17 19 19 19 19 19 19 19 19 19 19

Sample 80X5 36-38 82X3 38-40 83X1 42-44 1H1 40-42 2H1 40-42 3H3 40-42 3H6 40-42 4H6 38-40 6H1 43-45 7H5 40-42 7H7 40-42 8H1 40-42 8H2 40-42


(g/cm2 * kyr)0.12 0.13 0.20 0.02 0.01 0.03 0.07 0.07 0.06 0.07 0.07 0.00 0.00

Site 19 19 19 19 19 19 19 19 19 19 19 19 19

Sample 8H4 40-42 8H5 40-42 8H7 42-44 8H8 33-35 9H1 40-42 9H2 40-42 10H2 40-42 11H1 94-96 11H3 41-43 12X3 40-42 13X3 40-42 14X3 39-41 18X2 40-42

Depth (mbsf) 69.70 71.20 74.22 75.13 76.20 77.70 87.20 95.74 97.82 100.73 110.05 120.79 130.90BSi (wt %) 0.26 0.30 0.21 0.24 0.24 0.24 0.25 0.23 0.16 0.29 1.22 1.24 1.03

T.Cawthern

etal./

Marine

andPetroleum

Geology

xxx(2014)

1e12

6Pleasecite

thisarticle

inpress

as:Cawthern,T.,etal.,A

lateMiocene

eEarly

Pliocenebiogenic

silicacrash

inthe

Andam

anSea

andBay

ofBengal,Marine

andPetroleum

Geology

(2014),http://dx.doi.org/10.1016/j.marpetgeo.2014.07.026

Age

(Ma)

1.21

1.40

1.80

1.92

2.51

2.64

3.47

3.92

3.96

4.01

4.17

4.37

4.55

BSi

MAR

(g/cm

2*ky

r)0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.01

0.01

0.01

0.06

0.06

0.05

Site

1919

1919

1919

1919

1919

1919

19

Sample

18X640

-42

20X440

-42

21X540

-42

22X440

-42

23X343

-45

24X140

-42

24X440

-42

24X640

-42

28X240

-42

28X440

-42

28X740

-42

29X240

-42

29X540

-42

Dep

th(m

bsf)

136.88

153.20

164.29

172.40

180.53

187.20

191.70

194.70

197.13

199.83

203.22

206.82

211.32

BSi

(wt%)

0.34

0.59

0.87

0.33

0.30

0.26

0.26

0.48

0.42

0.34

0.50

0.85

0.83

Age

(Ma)

4.66

4.95

5.15

5.30

5.44

5.57

5.88

6.08

6.25

6.44

6.67

6.87

7.05

BSi

MAR

(g/cm

2*ky

r)0.02

0.03

0.04

0.02

0.02

0.00

0.00

0.01

0.01

0.01

0.01

0.02

0.02

Site

1919

1919

1919

1919

1919

1919

19

Sample

30X240

-42

30X540

-42

31X240

-42

31X540

-42

32X240

-42

32X540

-42

33X140

-42

33X340

-42

33X535

-37

34X139

-41

34X236

-38

34X439

-41

35X640

-42

Dep

th(m

bsf)

216.15

219.55

226.81

231.31

236.05

240.54

244.80

246.77

249.72

254.39

254.93

257.96

270.86

BSi

(wt%)

1.68

1.24

1.20

1.40

1.59

2.02

0.87

1.59

0.49

1.30

1.15

1.30

1.30

Age

(Ma)

7.24

7.38

7.67

7.85

8.04

8.22

8.39

8.46

8.58

8.77

8.79

8.91

9.43

BSi

MAR

(g/cm

2*ky

r)0.04

0.03

0.03

0.03

0.04

0.05

0.02

0.04

0.01

0.03

0.03

0.03

0.03

Site

1919

1919

1919

19

Sample

36X140

-42

36X239

-41

36X442

-44

36X637

-39

37X440

-42

38X140

-42

38X540

-42

Dep

th(m

bsf)

273.70

275.19

277.97

280.66

286.98

292.80

298.80

BSi

(wt%)

1.48

1.64

1.16

1.40

1.12

0.76

0.48

Age

(Ma)

9.54

9.60

9.71

9.82

10.07

10.30

10.54

BSi

MAR

(g/cm

2*ky

r)0.04

0.04

0.03

0.03

0.03

0.02

0.01



samples was analyzed in order to evaluate instrument drift and/orinaccuracies.

Biogenic silica weight percentage was determined on bulksamples from Sites 17 and 19 at fairly regularly spaced intervalsthroughout the cores. Weight percent BSi for each measuredsample was later converted into an instantaneous mass accumu-lation rate for discrete sampling intervals throughout the cores.Mass accumulation rates were calculated by multiplying the linearsedimentation rate (LSR) from the calcareous nannofossil agemodel (Fig. 3) by the dry bulk density (DBD), determined fromphysical properties measurements (Collett et al., 2008), and theweight percent BSi.

6. Results

The cores at Sites 17 and 19 were sampled at relatively constant~100e300 kyr intervals from the top of each core to the base. Therelatively coarse resolution of the sampling intervals permits thediscussion of only general trends in the data at both core locationsover million year time scales (Table 1).

The results from Site 17 are divided into three discrete intervals:1) a section of relatively high BSi accumulation from ~9.46 Ma to~5 Ma; 2) two segments of relatively low biogenic silica accumu-lation, one from ~5 Ma to ~4 Ma and another from ~1.7 Ma to thepresent; and 3) a period of increased BSi accumulation thatoccurred ~2.5 Ma (Fig. 5). The first interval exhibits variable rates ofBSi accumulation that range between ~0.7 g/cm2 * kyr at ~8.5 Maand ~0.002 g/cm2 * kyr at ~5 Ma. BSi initially increased within thisinterval from ~0.12 g/cm2 * kyr at ~9.2 Ma to the highest measuredvalue of ~0.7 g/cm2 * kyr at ~8.5 Ma. Mass accumulation ratessubsequently decreased from ~8.2 Ma and reached values of~0.14 g/cm2 * kyr by ~7.2 Ma before they increased again to ~0.3 g/cm2 * kyr by ~6.7 Ma. Despite this variability, the mass accumula-tion rates of BSi throughout this interval are the highest recordedMARs in the core. The second highest observed mass accumulationrate of BSi at Site 17 occurred from ~2.6 Ma to ~1.7 Ma with ratesthat attained a peak of ~0.07 g/cm2 * kyr at ~2.5Ma. This pulse in BSiaccumulation, however, is truncated between ~3.8 Mae2.6 Mabecause of a hiatus (most likely an unconformity) in the nannofossilbiostratigraphy (Flores et al., 2014). It is therefore difficult to sayexactly when the relative increase in BSi accumulation began,following the period of low BSi accumulation between ~5 Ma and~4 Ma, but the overall trend is that of gradual increase in BSi MARbetween ~5 and 4Ma (Fig. 5). The accumulation of BSi has remainedlow in the Andaman Sea region at Site 17 since ~1.7 Ma. In general,the MAR of biogenic silica at Site 17 decreased from ~9.4 Ma to~5 Ma and never returned to the pre-crash BSi levels observedduring the late Miocene.

The data from Site 19 are somewhat more variable than thosefrom Site 17 and the BSi MAR is an order of magnitude lower overallat Site 19 (Figs. 4 and 5, Table 1). The record of biogenic silicaaccumulation at Site 19 is divided into four discrete intervals: 1) asection of moderately high BSi accumulation that ranged between~0.01 g/cm2 * kyr at ~10.5 Ma and ~0.006 g/cm2 * kyr at ~6.3 Mawith a maximum rate of accumulation of ~0.05 g/cm2 * kyr at~8.22 Ma; 2) a second period of moderately high BSi accumulationthat ranged between ~0.004 g/cm2 * kyr at ~5.5 Ma and ~0.003 g/cm2 * kyr at ~3.8 Ma with a maximum rate of accumulation of~0.06 g/cm2 * kyr at ~4.2 Ma; 3) a third section of maximum BSiaccumulation that ranged between ~0.001 g/cm2 * kyr at ~0.8 Maand ~0.02 g/cm2 * kyr at present and consists of the highest rates ofaccumulation of ~0.07 g/cm2 * kyr at ~0.4 Ma; 4) three periods oflow BSi accumulation: one between ~6.3 Ma and ~5.5 Ma, a secondbetween ~3.65 Ma and 2.44 Ma, and a third from ~1.96 to 0.8 Ma(Fig. 5). A hiatus (most likely an unconformity) in the nannofossil


Figure 6. Porewater Si concentrations (gray area fill) and BSi MAR (yellow area fill) atSite 17 with respect to age within the core and ages of known occurrences of volcanicashes (red dots). Porewater Si concentrations (Rose et al., 2014) appear to be primarilyderived from amorphous volcanic glass phases, particularly in Subunits Ia and Ib,rather than the dissolution of biogenic silica tests, as evidenced by the covariance ofashes and high porewater [Si]. (For interpretation of the references to color in thisfigure legend, the reader is referred to the web version of this article.)

Figure 5. Biogenic silica mass accumulation rates at Sites 17 and 19 with respect to agein the core. BSi MAR is higher at Site 17 in the Andaman Sea compared to Site 19 in theMahanadi Basin.


biostratigraphy from ~3.8 to 3.65 Ma and ~2.44e1.96 Ma prohibitinterpretations within these intervals (Flores et al., 2014). In gen-eral, the MAR of biogenic silica at Site 19 decreased from ~10.5 Mato ~5.5Ma, with productivity recoveringmore rapidly following theBSi crash at ~6 Ma compared to Site 17. We consider the dramaticdecrease in BSi MAR at Site 17 from ~9.4 to 5 Ma and at Site 19 from~10.5 to 5.5 Ma to be time-coincident in the Andaman Sea and Bayof Bengal, and correlative to the reduction of BSi MAR reported inthe South China Sea (Wang et al., 2004), California margin (Farrellet al., 1995; Janecek, 2000), and elsewhere in the Pacific (Corteseet al., 2004) and Indian Oceans (Dickens and Owen, 1999).

7. Discussion

Quantitative analysis of percent biogenic silica has been thefocus of few studies in the Indian Ocean (Bohrmann and Ehrmann,1991; Hempel and Bohrmann, 1990; Schmitz, 1987). These studiesprovide evidence of two distinct phases of biogenic opal produc-tivity in the Indian Ocean, one from the Eocene-Oligocene(~35.8 Ma) and a second from late middle to late Miocene(~11 Ma) (Baldauf et al., 1992). To date, however, no studies of BSi


exist that have focused on long timescale productivity changes inthe northern Bay of Bengal or Andaman Sea regions.

The use of biogenic silica as a proxy for productivity in the watercolumn can be difficult to interpret, particularly in regions wherevolcanic glass is pervasive. Fortunately, the time-integrated disso-lution steps used in the modified wet alkaline technique ofDeMaster (1981) and Strickland and Parsons (1972) accounts forthe different rates of dissolution of amorphous and mineral phasesand therefore facilitates the identification of impure volcanic glassversus BSi signatures. Moreover, the discretionary selection ofsamples from intervals within the core devoid of volcanic ashes actsas a first step in ensuring the analysis for BSi does not contain asignificant signature from volcanic glasses. The purity of BSi sig-natures from Site 17 is not only revealed from the slope breaks inthe time-integrated dissolutions datasets (Fig. 4), but also in a plotof the downhole porewater Si concentration (Fig. 6). The concen-tration of silica in pore fluids from Site 17 varies such that the coremay be divided into three intervals: 1) an interval of high porewaterSi from ~9.5 Mae5.7 Ma, 2) a section of low porewater Si from~5.7Mae2.6Ma, and 3) an interval of moderately high porewater Siconcentration from ~2.6 Mae0.5 Ma. In the record at Site 17, theporewater Si concentration (Fig. 6) is highest in the core near thebottom (~9.5e5.7 Ma), where the mass accumulation rate of BSi ishigh and there are also abundant volcanic ashes. In the middle ofthe core (~5.7e2.6 Ma), where no continuous volcanic ash layersare present, although disseminated ashes are observed in lowabundances and the BSi MAR is low, the porewater Si concentra-tions are low. In contrast to both of the deeper portions of the core,the top of the core (the last ~2.6 Myrs) shows a bulge in porewaterSi that corresponds with the occurrence of volcanic glass, and anearly constant and low occurrence of BSi (Fig. 6). These resultsimply that the sediment porewater silica at Site 17 may be largelyderived from dissolution of volcanic glass phases. If the measure-ments of BSi in the uppermost part of the record at Site 17



(~2.6e0 Ma) had incorporated volcanic glass, the BSi profile wouldmore closely match the porewater profile in this interval. Instead,the porewater profile covaries with the abundance of volcanic glassin the record, particularly in Subunits Ia and Ib, where biogenicsilica is less abundant. Porewater Si in Subunit Ic is most likelycomprised of an admixture of both BSi and amorphous silica fromvolcanogenic sources. Coupled with the strength of separation inthe slope-breaks from the time-integrated BSi dissolutions (Fig. 4),the BSi data reported here do not include a significant signaturefrom volcanic glass in the Site 17 record. Because Site 19 is locatedfarther from the Sunda volcanic arc and only one ash occurrencewas documented in the core (Collett et al., 2008), the BSi mea-surements here are not affected by amorphous volcanic glasses.

The sedimentary records from Sites 17 and 19 are distinct fromone another as a result of their different depositional settings. Site17 preserves a predominantly marine pelagic record that containsvariable amounts of airfall volcanic ashes and fine grained (clay-sized) clastic materials (Collett et al., 2008). In contrast, Site 19 ispredominantly a hemipelagic record that consists of variable con-centrations of marine carbonate and biosiliceous fossils, terrestrialplant remains, and clastic detritus that most likely originated fromthe Mahanadi River (Collett et al., 2008; Phillips et al., 2014).Despite these differences, shipboard smear-slide and coarse-fraction analyses of these records revealed a general pattern ofdecreased biosiliceous fossils from the bottom to the top of thecores. These observations are now quantified through wet chemicalanalytical techniques and reveal the timing of a significant crash inthe BSi MAR at both Sites 17 and 19. In general, the biogenic silicacrash at Site 17 occurred between ~6.7 and 5 Ma, whereas the crashat Site 19 occurred between ~7.2 and 5.5 Ma. The crash in both corelocations is time-correlative. The duration of low BSi MAR thatfollowed the crash at Site 17 is difficult to estimate due to thepresence of an unconformity that lasted from ~3.8 to 2.6 Ma, but amoderate pulse in productivity is apparent that began prior to~3.8 Ma and continued after ~2.6 Ma. In contrast, BSi MARrebounded at Site 19 after a period of low productivity that lastedapproximately 0.7 Myr. Despite the different environments ofdeposition between these two core locations, biogenic silica pro-ductivity noticeably decreased between ~7.2 and 5 Ma at both corelocations. This crash is correlative to the reduction of BSi reportedin the South China Sea (Wang et al., 2004), California margin(Farrell et al., 1995; Janecek, 2004), and elsewhere in the Pacific(Cortese et al., 2004) and Indian Oceans (Dickens and Owen, 1999).

Numerous studies (e.g., Kennett et al., 1985; Martin and Scher,2006; Nishimura and Suparka, 1997) of the southern IndianOcean and western Pacific Ocean have correlated changes ingeotectonic, paleontological and geochemical datasets to the pro-gressive restriction of the Indonesian Throughflow (ITF) over longtimescales. The timescale (~23.7 Mae5.3 Ma) of these changesmakes it difficult to rectify orbital to sub-orbital scale changes, butmay reflect longer timescale changes in sedimentation rates andocean productivity due to large changes in the Indian monsoonsystem or oceanographic change. These datasets suggest that theflow of deep, nutrient-rich Pacific Ocean waters through the ITFbecame restricted by ~6 Ma (Hall, 1996, 2002; Kennett et al., 1985;Nishimura and Suparka, 1997; Martin and Scher, 2006), but wasperhaps periodically re-established thereafter to give rise to morerecent (late Pliocene ~4e3 Ma) estimates of complete deep waterrestriction (e.g., Cane and Molnar, 2001; Gourlan et al., 2007). Thetiming of the primary biogenic silica crash at Sites 17 and 19(~7.2e5 Ma) in the Andaman Sea and northern Bay of Bengal, aswell as opal MARs fromODP Site 1143 in the South China Sea (Wanget al., 2004) provide evidence of a dramatic change in paleo-productivity coincident in time with estimates for the timing of therestriction of the ITF at ~6 Ma.


Upwelling-induced productivity for the last ~40 kyr in the Bay ofBengal and Arabian Sea has been related to variations in thestrength of the summer monsoon (e.g., Schubert et al., 1998;Sijinkumar et al., 2010). When the summer monsoon is strong,increased rainfall over the Indian subcontinent provides a highsupply of H4SiO4-rich runoff to the Indian offshore region andenhanced southwesterly-directed winds that could potentiallypromote upwelling (Filippelli, 1997; Dettman et al., 2001; Gregg,2002; Gupta and Thomas, 2003; Hathorne and James, 2006; Cliftet al., 2008). Upwelling-induced productivity, however, has beenshown to be severely limited due to freshwater runoff during thesummer monsoon, particularly along the continental shelf andslope, where a wedge of low salinity, low nutrient surface waterstratifies the water column and inhibits productivity (PrassanaKumar et al., 2002; Madhupratap et al., 2003; Bolton et al., 2013).Over the past 110 kyr, biological productivity at Site 19 decreasedduring periods of enhanced summer monsoon intensity; however,BSi MAR increased relative to CaCO3 MAR during these periods ofintensified summer monsoon (Phillips et al., 2014), suggesting BSiMAR may be decoupled from overall productivity and related tosilicate availability in the northern Bay of Bengal. In the semi-restricted Andaman Sea, upwelling-induced productivity occursas a result of mesoscale cyclonic eddies generated due to seasonalfluctuations in the monsoonal winds (Wyrtki, 1961;Buranapratheprat et al., 2010; Rizal et al., 2012). On million yeartimescales, BSi MAR data from both Sites 17 and 19 record the samedecrease in BSi MAR from ~7.2 to 5Ma. Differences in BSi MAR since~5 Ma between the two sites are most likely attributed to theproximity of point sources of nutrients such as the Mahanadi,Irrawaddy, and GangeseBrahmaputra Rivers or localized sites ofupwelling that may have provided a means of sustaining the mostrecent pulses in BSi productivity at Sites 17 and 19, particularlyduring phases of reduced monsoon intensity.

The abrupt decrease in productivity of biogenic silica at Sites 17and 19 from ~7.2 to 5 Ma is consistent with geochemical, paleon-tological, and geotectonic datasets that support the gradual re-striction of the gateway linking deep Pacific Ocean water masses tothe Indian Ocean via the Indonesian Throughway from early to lateMiocene (~23.7 Mae5.3 Ma). Today, the Andaman Sea region haslimited deep water hydrographic communication to the SouthChina and Java Seas through the Strait of Malacca (Keller andRichards, 1967; Rizal et al., 2012). However, the flow into and outof this region may have been enhanced prior to ~6 Ma and couldhave provided enough vital nutrients to the Site 17 region and eveninto the Bay of Bengal to sustain high rates of BSi surface produc-tivity in both regions. As communication between the AndamanSea and the ITF gradually decreased, themass accumulation rates ofBSi also decreased. By ~6.7e5 Ma, the influx of nutrients that sus-tained BSi productivity may have been limited resulting in theobserved crash in the BSi MAR. A short pulse in BSi productivity atSite 17 between ~4 and 2.5 Ma could reflect riverine sources ofnutrients from the nearby Irrawaddy and GangeseBrahmaputraRivers and/or favorable localized upwelling. The pattern of BSi MARat Site 17 are nearly identical to those observed at ODP Site 1143 inthe South China Sea (Wang et al., 2004) and supports previousstudies that report diminished ITF transport by ~6 Ma (e.g., Hall,1996, 2002; Kennett et al., 1985; Nishimura and Suparka, 1997;Martin and Scher, 2006). Because this time interval spans the un-conformity from ~3.8 to 2.6 Ma at Site 17, however, it is difficult toassess the true signature of any recovery of BSi in the Andaman Sea.The cessation of this secondary pulse in BSi productivity at Site 17 isconsistent with a late Pliocene (~4e3 Ma) estimate for the finalrestriction of deep water mass flow through the ITF from Cane andMolnar (2001) and Gourlan et al. (2007). Alternatively, enhancedsupply of nutrient-rich freshwater from the nearby Irrawaddy and



Mahanadi Rivers in response to the documented increase in theintensity of the Indian monsoon at ~5 Ma (Zhisheng et al., 2001;Curray et al., 2003; Gupta and Thomas, 2003) may have initiatedthe increase in productivity after this time.

Tracking the flow of Pacific Ocean water masses as they flowthrough the ITF and into the Indian Ocean has been the focus ofnumerous studies (e.g., Lee et al., 2002; Valsala and Ikdea, 2007;Sengupta et al., 2013). In general, once entering the Indian Ocean,ITF waters flow westward across the Indian Ocean and uponreaching the western boundary, re-route into three branches: 1) aneastward-flowing branch located just south of the equator, 2) aneastward-flowing branch located just north of the equator, 3) anorthward-flowing branch that ultimately upwells in the Somaliregion and spreads as surface waters across the northern IndianOcean (Valsala and Ikdea, 2007). The third branch of the ITF withinthe Indian Ocean may have had a profound effect on surface-waterproductivity in the Site 19 region. In particular, the restriction ofand concomitant re-routing of ITF waters around Australia at ~6Mamay explain the similar timing of the BSi MAR crash at Site 19(~7.2e6.3 Ma) to the BSi MAR crash that occurred in the AndamanSea and elsewhere in the southern Indian and Pacific Oceans, aswell as marginal seaways such as the South China Sea. Because theITF is much more modulated in the vastly larger Indian Oceancompared to the semi-restricted Andaman Sea, the effects of areduction in ITF may have been less pronounced in the northernIndian Ocean. Nevertheless, the crash in BSi MAR at Sites 17 and 19is coincident and most likely reflects a major change in the deliveryof nutrient-rich deep Pacific Ocean waters to the Bay of Bengal andAndaman Sea. The relatively rapid recovery of BSi MAR at Site 19was likely primarily controlled by surface-water runoff andregionally derived nutrients from the Indian subcontinent (e.g.,Farrell et al., 1995; Filippelli, 1997; Dettman et al., 2001; Gupta andThomas, 2003; Hathorne and James, 2006; Clift et al., 2008), whichwould have increased with the documented strengthening of themonsoon after ~5 Ma (Zhisheng et al., 2001; Curray et al., 2003;Gupta and Thomas, 2003).

8. Conclusions

Quantitative analysis of biogenic silica was performed on 68bulk sediment samples from NGHP-01 Site 17 and 56 samples fromNGHP-01 Site 19 in order to better understand an abrupt crash inbiogenic silica at both of these sites. The primary BSi crash at bothof these northeast Indian Ocean sites occurred between ~7.2 and5 Ma, which is in good agreement with previously reported studiesof the geotectonics, paleontological and geochemistry of marinesediments from the southern Indian Ocean and western PacificOcean, which document the closure of the Indonesian Throughflowat ~6 Ma. In contrast to BSi mass accumulation rates observed insediments from the northern Bay of Bengal at Site 19, which appearto be influenced heavily by freshwater runoff that emanated fromthe Mahanadi River, those of the Andaman Sea region appear tohave beenmuchmore sensitive to changes in the ITF. Biogenic silicawithin the Andaman Sea region never fully recovered to lateMiocene levels following the closure of a northern lobe of ITF be-tween ~6.7 and 5 Ma, whereas productivity at Site 19 was rejuve-nated coincident with the proposed strengthening of the monsoonafter ~5 Ma.

Acknowledgments

The authors wish to thank those that contributed to the successof the National Gas Hydrate Program Expedition 01 (NGHP-01).NGHP-01 was planned and managed through collaboration be-tween the Directorate General of Hydrocarbons (DGH) under the


Ministry of Petroleum and Natural Gas (India), the U.S. GeologicalSurvey (Contract # 07CRSA0708) (USGS), and the Consortium forScientific Methane Hydrate Investigations (CSMHI) led by Over-seas Drilling Limited (ODL) and FUGRO McClelland Marine Geo-sciences. The platform for the drilling operation was the researchdrill ship JOIDES Resolution, operated by ODL. Much of the drilling/coring equipment used was provided by the Integrated OceanDrilling Program (IODP) through a loan agreement with the USNational Science Foundation. The financial support for NGHP-01,from the Oil Industry Development Board, Oil and Natural GasCorporation Ltd., GAIL (India) Ltd., and the Oil India Ltd., isgratefully acknowledged. This research was supported by theNational Gas Hydrates Program of India, the Ed Picou FellowshipGrant for Graduate Studies in Earth Sciences from the Gulf CoastSection of the Society for Sedimentary Geology, a University ofNew Hampshire Department of Earth Sciences Research Awardand Natural Resources and Earth Systems Science ProgramResearch Award. We thank the entire NGHP-01 Expedition ship-board science party and crew for the characterization of thestratigraphy and for the collection of the cores in 2006. The BSimeasurements at UNH would not have been possible without thesupport of Linda Kalnejais and the generosity of Laurie Westoverat the University of New Hampshire.

References

Amiruddin, A.M., Ibrahim, Z.Z., Ismail, S.A., 2011. Water mass characteristics in theStrait of Malacca using ocean data view. Res. J. Environ. Sci. 5, 49e58.

Babu, C.P., Pattan, J.N., Dutta, K., Basavaiah, N., Ravi Prasad, G.V., Ray, D.K., Govil, P.,2010. Shift in detrital sedimentation in the eastern Bay of Bengal during the lateQuaternary. J. Earth Syst. Sci. 119, 285e295.

Baldauf, J.G., Barron, J.A., Ehrmann, W.U., Hempel, P., Murray, D., 1992. Biosiliceoussedimentation patterns for the Indian Ocean during the last 45 million years,American Geophysical Union. Geophys. Monogr. 70, 335e349.

Bastia, R., Das, S., Radhakrishna, M., 2010. Pre- and post-collisional depositionalhistory in the upper and middle Bengal fan and evaluation of deepwaterreservoir potential along the northeast continental margin of India. Mar. Pet.Geol. 27, 2051e2061.

Berger, W.H., Leckie, R.M., Janecek, T.R., Stax, R., Takayama, T., 1993. Neogene car-bonate sedimentation on Ontong Java Plateau: highlights and open questions.Proc. Ocean. Drill. Program, Sci. Results 130, 711e744.

Bohrmann, G., Ehrmann, W.U., 1991. Analysis of sedimentary facies using bulkmineralogic characteristics in Cretaceous to Quaternary sediments from theKerguelen Plateau: sites 737, 738 and 744. In: Barron, J.A., Larsen, B., et al. (Eds.),Proceedings of the Ocean Drilling Program, Scientific Results, vol. 119. OceanDrilling Program, College Station, TX, pp. 211e223.

Bolton, C.T., Change, L., Clemens, S.C., Kodama, K., Ikehara, M., Medina-Elizalde, M.,Paterson, G.A., Roberts, A.P., Rohling, E.J., Yamamoto, Y., Zhao, X., 2013.A 500,000 year record of Indian summer monsoon dynamics recorded byeastern equatorial Indian Ocean upper water-column structure. Quat. Sci. Rev.77, 167e180.

Bowles, F., Ruddiman, W., Jahn, W., 1978. Acoustic stratigraphy, structure, anddepositional history of the Nicobar Fan, Eastern Indian Ocean. Mar. Geol. 26,269e288.

Buranapratheprat, A., Laongmanee, P., Sukramongkol, N., Prommas, R.,Promjinda, S., Yanagi, T., 2010. Upwelling induced by meso-scale cyclonic eddiesin the Andaman Sea. Coast. Mar. Sci. 34, 68e73.

Calvert, 1966. Accumulation of diatomaceous silica in the sediment of the Gulf ofCalifornia. Geol. Soc. Am. Bull. 77, 569e572.

Cane, M.A., Molnar, P., 2001. Closing of the Indonesian Seaway as a precursor to eastAfrican aridification around 3-4 million years ago. Nature 411, 157e162.

Clift, P.D., Hodges, K.V., Heslop, D., Hannigan, R., Van Long, H., Calves, G., 2008.Correlation of Himalayan exhumation rates and Asian monsoon intensity. Nat.Geosci. 1, 875e880.

Collett, T.S., Riedel, M., Cochran, J., Boswell, R., Presley, J., Kumar, P., Sathe, A.,Sethi, A., Lall, M., Sibal, V., the NGHP Expedition-01 Scientists, 2008. NationalGas Hydrate Program Expedition 01 Initial Reports.

Cortese, G., Gersonde, R., Hillenbrand, C.-D., Kuhn, G., 2004. Opal sedimentationshifts in the world ocean over the last 15 Myr. Earth Planet. Sci. Lett. 224,509e527.

Curray, J.R., Emmel, F.J., Moore, D.G., 2003. The Bengal Fan: morphology, geometry,stratigraphy, history and processes. Mar. Pet. Geol. 19, 1191e1223.

DeMaster, D.J., 1981. The supply and accumulation of silica in the marine environ-ment. Geochim. Cosmochim. Acta 45, 1715e1732.

Dettman, D.L., Kohn, M.J., Quade, J., Ryerson, F.J., Ojha, T.P., Hamidullah, S., 2001.Seasonal stable isotope evidence for a strong Asian monsoon throughout thepast 10.7 m.y. Geology 29, 31e34.


http://refhub.elsevier.com/S0264-8172(14)00251-7/sref2




































































Dickens, G.R., Owen, R.M., 1994. Late MioceneeEarly Pliocene manganese redirec-tion in the central Indian Ocean: expansion of the intermediate water oxygenminimum zone. Paleoceanography 9, 169e181.

Dickens, G.R., Owen, R.M., 1996. Sediment geochemical evidence for an Early-middle Gilbert (Early pliocene) productivity peak in the North Pacific red ClayProvince. Mar. Micropaleontol. 27, 107e120.

Dickens, G.R., Owen, R.M., 1999. The Latest MioceneeEarly pliocene biogenicbloom: a revised Indian Ocean perspective. Mar. Geol. 161, 75e91.

Dutta, K., Bhushan, R., Somayajulu, B.L.K., 2007. Rapid vertical mixing rates in deepwaters of the Andaman Basin. Sci. Total Environ. 384, 401e408.

Fang, G., Susanto, R.D., Wirasantosa, S., Qiao, F., Supangat, A., Fan, B., Wei, Z.,Sulistiyo, B., Li, S., 2010. Volume, heat, and freshwater transports from teh SouthChina Sea to Indonesian seas in the boreal winter of 2007e2008. J. Geophys.Res. Oceans 115, 1e11.

Fang, G., Wang, Y., Wei, Z., Fang, Y., Qiao, F., Hu, X., 2009. Interocean circulation andheat and freshwater budgets of the South China Sea based on a numericalmodel. Dyn. Atmos. Oceans 47, 55e72.

Fang, G., Susanto, R.D., Soesilo, I., Zheng, Q., Qiao, F., Wei, Z., 2005. Notes on theupper-layer interocean circulation of the South China Sea. Adv. Atmos. Sci.COAA Spec. Issue 22, 946e954.

Farrell, J.W., Raffi, I., Janecek, T.R., Murray, D.W., Levitan, M., Dadey, K.A., Emeis, K.-C., Lyle, M., Flores, J.-A., Hovan, S., 1995. Late Neogene sedimentation patterns inthe eastern equatorial Pacific Ocean. In: Pisias, N.G., Mayer, L.A., Janecek, T.R.,Palmer-Julson, A., van Andel, T.H. (Eds.), Proceedings of the Ocean DrillingProgram, Scientific Results, vol. 138. Ocean Drilling Program, College Station, TX,pp. 717e756.

Filippelli, G.M., 1997. Intensification of the Asian monsoon and a chemical weath-ering event in the late MioceneeEarly Pliocene: implications for late Neogeneclimate change. Geology 25, 27e30.

Flores, J.-A., Johnson, J.E., Mejia-Molina, A.E., Alvarez, M.C., Sierro, F.J., Singh, S.D.,Mahanti, S., Giosan, L., 2014. Sedimentation rates from calcareous nannofossiland planktonic foraminifera biostratigraphy in the Andaman Sea, northern Bayof Bengal, and eastern Arabian Sea. J. Mar. Pet. Geol.

Garcia, H.E., Locarnini, R.A., Boyer, T.P., Antonov, J.I., Zweng, M.M., Baranova, O.K.,Johnson, D.R., 2010. World ocean atlas 2009. In: Levitus, S. (Ed.), Nutrients(Phosphate, Nitrate, Silicate), NOAA Atlas NESDIS 71, vol. 4. U.S. GovernmentPrinting Office, Washington, D.C., p. 398.

Gordon, A.L., Fine, R.A., 1996. Pathways of water between the Pacific and IndianOceans in the Indonesian Seas. Lett. Nat. 379, 146e149.

Gordon, A.L., Huber, B.A., Metzger, E.J., Susanto, R.D., Hurlburt, H.E., Adi, T.R., 2012.South China Sea Throughflow impact on the Indonesian Throughflow. Geophys.Res. Lett. 39, 1e7.

Gourlan, A.T., Meynadier, L., Allegre, C.J., 2007. Tectonically driven changes in theIndian Ocean circulation over the last 25 Ma: Neodymium isotope evidence.Earth Planet. Sci. Lett. 267, 353e364.

Gregg, W.W., 2002. Tracking the Sea WIFs record with a coupled physical/biogeo-chemical/radiative model of the global oceans. Deep Sea Res. II 49, 81e105.

Gupta, A.K., Thomas, E., 2003. Initiation of Northern Hemisphere glaciation andstrengthening of the northeast Indian monsoon: ocean Drilling Program Site758, eastern equatorial Indian Ocean. Geology 31, 47e50.

Hall, R., 1996. Reconstructing Cenozoic SE Asia. In: Hall, R., Blundell, D. (Eds.),Tectonic Evolution of Southeast Asia, Geological Society of London SpecialPublication, vol. 106, pp. 153e184.

Hall, R., 2002. Cenozoic geological and plate tectonic evolution of SE Asia and theSW Pacific: computer-based reconstructions, model and animations. J. AsianEarth Sci. 20, 353e434.

Haq, B.U., Hardenbol, J., Vail, P.R., 1987. Chronology of fluctuating sea levels since theTriassic. Science 235, 1156e1167.

Hathorne, E.D., James, R.H., 2006. Temporal record of lithium in seawater: a tracerfor silicate weathering? Earth Planet. Sci. Lett. 246, 393e406.

Hempel, P., Bohrmann, G., 1990. Carbonate free sediment components and aspectsof silica diagenesis at Sites 707, 709 and 711. In: Duncan, R.A., Backman, J.,Peterson, L., et al. (Eds.), Proceedings of the Ocean Drilling Program, ScientificResults, vol. 115. Ocean Drilling Program, College Station, TX, pp. 677e698.

Ibrahim, Z.Z., Yanagi, T., 2006. The influence of the Andaman Sea and the SouthChina Sea on water masses in the Malacca Strait. Mer 43, 33e42.

Janecek, T.R., 2000. Data report: late Neogene biogenic opal data for Leg 167 sites onthe California margin. In: Lyle, M., Koizumi, I., Richter, C., Moore Jr., T.C. (Eds.),Proceedings of the Ocean Drilling Program, Scientific Results, vol. 167. OceanDrilling Program, College Station, TX, pp. 213e214.

Jiang, S., Wise, S.W., Wang, Y., 2007. Cause of the Middle/Late Miocene carbonatecrash: dissolution or low productivity?. In: Teagle, D.A.H., Wilson, D.S.,Acton, G.D., Vanko, D.A. (Eds.), Proceedings of the Ocean Drilling Program,Scientific Results, vol. 206, pp. 1e24.

Johnson, J.E., Phillips, S.C., Torres, M.E., Pinero, E., Rose, K.K., Giosan, L., 2014. In-fluence of total organic carbon deposition on the inventory of gas hydrate in theIndian continental margins. J. Mar. Pet. Geol.

Karig, D.E., Lawrence, M., Moore, G., Curray, J., 1980. Structural framework of thefore-arc basin, NW Sumatra. J. Geol. Soc. Lond. 137, 77e91.

Kawagata, S., Hayward, B.W., Gupta, A.K., 2006. Benthic foraminiferal extinctionslinked to late PlioceneePleistocene deep-sea circulation changes in thenorthern Indian Ocean (ODP Sites 722 and 758). Mar. Micropaleontol. 58,219e242.


Keller, G.H., Richards, A.F., 1967. Sediments of the Malacca Strait, southeast Asia.J. Sediment. Petrol. 37, 102e127.

Keller, G.H., Barron, J.A., 1983. Paleoceanographic implications of Miocene deep-seahiatuses. Geol. Soc. Am. Bull. 94, 590e613.

Keller, G., 1985. Depth stratification of planktonic foraminifers in the Mioceneocean. In: Kennett, J.P. (Ed.), The Miocene Ocean, Geological Society of AmericaMemoir, vol. 163, pp. 177e195.

Kennett, J.P., Keller, G., Srinivasan, M.S., 1985. Miocene planktonic foraminiferalbiogeography and paleoceanography development of the Indo-Pacific region.In: Kennett, J.P. (Ed.), The Miocene Ocean: Paleoceanography and Biogeography,Geological Society of America Memoirs, vol. 163, pp. 197e236.

Khan, P.K., Chakraborty, P., 2005. Two-phase opening of Andaman Sea: a newseismotectonic insight. Earth Planet. Sci. Lett. 229, 259e271.

Lee, T., Fukumori, I., Menemenlis, D., Xing, Z., Fu, L.-L., 2002. Effects of the Indo-nesian Throughflow on the Pacific and Indian Oceans. J. Phys. Oceanogr. 32,1404e1429.

Leinen, 1979. Biogenic silica accumulation in the central equatorial Pacific and itsimplications for Cenozoic paleoceanography. Geol. Soc. Am. Bull. 90, 1310e1376.

Lyle, M., Murray, D., Finney, B., Dymond, J., Robbins, J., Brooksforce, K., 1988. Therecord of Late Pleistocene biogenic sedimentation in the eastern tropical PacificOcean. Paleoceanography 3, 39e59.

Madhupratap, M., Gauns, M., Ramaiah, N., Kumar, S.P., Muraleedharan, P.M., deSousa, S.N., Sardessai, S., Muraleedharan, U., 2003. Biogeochemistry of the Bayof Bengal: physical, chemical and primary productivity characteristics of thecentral and western Bay of Bengal during the summer monsoon 2001. Deep SeaRes. Part II 50, 881e896.

Martin, E.E., Scher, H., 2006. A Nd isotopic study of southern sourced waters andIndonesian Throughflow at intermediate depths in the Cenozoic Indian Ocean.Geochem. Geophys. Geosyst. 7, 1e14.

Martini, E., 1971. Standard tertiary and Quaternary calcareous nannoplanktonzonation. In: Farinacci, A. (Ed.), Proceedings of the 2nd International Conferenceon Planktonic Microfossils, vol. 2, pp. 739e785.

Mucciarone, D., 2003. Biogenic silica analysis procedure. In: Stanford UniversityStable Isotope Laboratory Standard Operating Procedure Manual. StanfordUniversity Department of Geological and Environmental Sciences, Stanford, CA,pp. 93e105.

Murthy, V.S.N., Sarma, Y.V.B., Rao, D.P., Murthy, C.S., 1992. Water characteristics,mixing and circulation in the Bay of Bengal during southwest monsoon. J. Mar.Res. 50, 207e228.

Naqvi, W.A., Charles, C.D., Fairbanks, R.G., 1994. Carbon and oxygen records ofbenthic foraminifera from the Northeast Indian Ocean: implications on glacial-interglacial atmospheric CO2 changes. Earth Planet. Sci. Lett. 121, 99e110.

Nishimura, S., Suparka, S., 1997. Tectonic approach to the Neogene evolution ofPacificeIndian Ocean seaways. Tectonophysics 281, 1e16.

Okada, H., Bukry, D., 1980. Supplementary modification and introduction of codenumbers to the low-latitude coccolith biostratigraphic zonation (Bukry, 1973,1975). Mar. Micropaleontol. 5, 349e379.

Pal, T., Chakraborty, P., Gupta, T., Singh, C., 2003. Geodynamic evolution of the outer-arcdforearc belt in the Andaman Islands, the central part of the Burma-Javasubduction complex. Geol. Mag. 140, 289e307.

Phillips, S.C., Johnson, J.E., Giosan, L., Rose, K., 2014. Monsoon-influenced variationin productivity and lithogenic sediment flux since 110 ka in the offshoreMahanadi Basin, northern Bay of Bengal. J. Mar. Pet. Geol.

Peterson, L.C., 1984. Recent abyssal benthic foraminiferal biofacies of the easternequatorial Indian Ocean. Mar. Micropaleontol. 8, 479e519.

Peterson, L.C., Murray, D.W., Ehrmann, W.U., Hempel, P., 1992. Cenozoic carbonateaccumulation and compensation depth changes in the Indian Ocean. In:Duncan, R.A., Rea, D.K., Kidd, R.B., Von Rad, U., Weissel, J.K. (Eds.), Synthesis ofResults from Scientific Drilling in the Indian Ocean, American GeophysicalUnion Geophysical Monograph, vol. 70, pp. 311e333.

Powell, C.M., Roots, S.R., Veevers, J.J., 1988. Pre-breakup continental extension inEast Gondwanaland and the early opening of the eastern Indian Ocean. Tec-tonophysics 155, 261e283.

Prassana Kumar, S., Muraleedharan, P.M., Prasad, T.G., Gauns, M., Ramaiah, N., deSouza, S.N., Sardesai, S., Madhupratap, M., 2002. Why is the Bay of Bengal lessproductive during summer monsoon compared to the Arabian Sea? Geophys.Res. Lett. 29 http://dx.doi.org/10.1029/2002GL016013.

Raffi, I., Backman, J., Fornaciari, E., Palike, H., Rio, D., Lourens, L., Hilgen, F., 2006.A review of calcareous nannofossil astrobiochronology encompassing the past25 million years. Quat. Sci. Rev. 25, 3113e3137.

Raffi, I., Flores, J.-A., 1995. Pleistocene through Miocene calcareous nannofossilsfrom eastern equatorial Pacific Ocean (Leg 138). In: Pisias, N.G., Mayer, L.A.,Janecek, T.R., Palmer-Julson, A., van Andel, T.H. (Eds.), Proceedings of the OceanDrilling Program, Scientific Results, vol. 138, pp. 233e286.

Rai, A.K., Singh, V.B., 2012. Response of eastern Indian Ocean (ODP Site 762B)benthic foraminiferal assemblages to the closure of the Indonesian seaway.Oceanologia 54, 449e472.

Raju, K.A., Murty, G., Amarnath, D., Kumar, M., 2007. The west Andaman fault and itsinfluence on the aftershock pattern of the recent megathrust earthquakes in theAndamaneSumatra region. Geophys. Res. Lett. 34, 1e5.

Raju, K.A., Ramprasad, T., Rao, P., Rao, B., Varghese, J., 2004. New insights into thetectonic evolution of the Andaman basin, northeast Indian Ocean. Earth Planet.Sci. Lett. 221, 145e162.
























































































































http://refhub.elsevier.com/S0264-8172(14)00251-7/sref37r































































http://refhub.elsevier.com/S0264-8172(14)00251-7/sref38rty
















http://dx.doi.org/10.1029/2002GL016013
























Rizal, S., Damm, P., Wahid, M.A., Sundermann, J., Ilhamsyah, Y., Iskandar, T.,Muhamma, 2012. General circulation in the Malacca Strait and Andaman Sea: anumerical model study. Am. J. Environ. Sci. 8, 479e488.

Rodolfo, K., 1969. Bathymetry and marine geology of the Andaman Basin, andtectonic implications for Southeast Asia. Geol. Soc. Am. Bull. 80, 1203e1230.

Rodolfo, K., 1975. The Irrawady Delta: tertiary setting and modern offshore sedi-mentation. Deltas: Models Explor. 339e356.

Romine, K., Lombari, G., 1985. Evolution of Pacific circulation in the Miocene:radiolarian evidence from DSDP Site 289. In: Kennett, J.P. (Ed.), The MioceneOcean: Paleoceanography and Biogeography, Geological Society of AmericaMemoirs, vol. 163, pp. 273e290.

Rose, K.K., Johnson, J.E., Torres, M.E., Hong, W., Giosan, L., Solomon, E.A., Kastner, M.,Cawthern, T., Long, P.E., Schaef, H.T., 2014. Anomalous porosity preservation andpreferntial accumulation of gas hydrate in the Andaman Accretionary Wedge,NGHP-01 Site 17A. J. Mar. Pet. Geol.

Schmitz, B., 1987. Barium, equatorial high productivity and the northward wan-dering of the Indian Continent. Paleoceanography 2, 63e77.

Schubert, D.J., Villanueva, J., Calvert, S.E., Cowie, G.L., von Rad, U., Schulz, H.,Berner, U., Erlenkeuser, H., 1998. Stable phytoplankton community structure inthe Arabian Sea over the past 200,000 years. Nature 394, 563e566.

Sengupta, S., Parekh, A., Chakraborty, S., Ravi Kumar, K., Bose, T., 2013. Verticalvariation of oxygen isotopes in the Bay of Bengal and its relationships withwater masses. J. Geophys. Res. Oceans 118 (12), 6411e6424.

Sijinkumar, A.V., Nath, B.N., Gupta, M.V.S., 2010. Late Quaternary record of pteropodpreservation from the Andaman Sea. Mar. Geol. 275, 221e229.

Strickland, J.D.H., Parsons, T.R., 1972. A Practical Handbook of Seawater Analysis,second ed. In: Bulletin 167 Fisheries Research Board of Canada, Ottawa.

Susanto, R.D., Fang, G., Soesilo, I., Zheng, Q., Qiao, F., Wei, Z., Sulistyo, B., 2010. SITE:South China Sea-Indonesian Seas transport/exchange. EOS Trans. 30, 261e263.


Susanto, R.D., Ffield, A., Gordon, A.L., Adi, T.R., 2012. Variability of IndonesianThroughflow within Makassar Strait, 2004-2009. J. Geophys. Res. 117, 1e16.

Tripathy, G.R., Singh, S.K., Bhushan, R., Ramaswamy, V., 2011. Sr-Nd isotopecomposition of the Bay of Bengal sediments. Impact Clim. Eros. Himal. 45,175e186.

Tozuka, T., Qu, T., Yamagata, T., 2007. Dramatic impacts of the South China Sea onthe Indonesian Throughflow. Geophys. Res. Lett. 34, 1e5.

Valsala, V.K., Ikdea, M., 2007. Pathways and effects of the Indonesian Throughflowwater in the Indian Ocean using particle trajectory and tracers in an OGCM.J. Clim. 20, 2994e3017.

Varkey, M.J., Murty, V.S.N., Suryanarayana, A., 1996. Physical oceanography of theBay of Bengal and Andaman Sea. Oceanogr. Mar. biol.: An Annu. Rev. 34, 1e70.

Vincent, E., 1974. Cenozoic planktonic biostratigraphy and paleoceanography of thetropical western Indian Ocean. Initial Rep. Deep Sea Drill. Proj. 24, 1111e1150.

Wang, R., Li, J., Li, B., 2004. Data report: Late Miocene-Quaternary biogenic opalaccumulation at ODP Site 1143, southern South China Sea. In: Prell, W.L.,Wang, P., Blum, P., Rea, D.K., Clemens, S.C. (Eds.), Proceedings of the OceanDrilling Program, Scientific Results, vol. 184. Ocean Drilling Program, CollegeStation, TX, pp. 1e12.

Woodruff, F., 1985. Changes in Miocene deep-sea benthic foraminiferal distributionin the Pacific Ocean: relationship to paleoceanography. Geol. Soc. Am. Bull. 143,131e171.

Wyrtki, K., 1961. Physical Oceanography of the Southeast Asian Waters, first ed.University of California, California, p. 195.

Zhisheng, A., Kutzbach, J.E., Prell, W.L., Porter, S.C., 2001. Evolution of Asian mon-soons and phased uplift of the HimalayaeTibetan plateau since Late Miocenetimes. Lett. Nat. 411, 62e66.




























































http://refhub.elsevier.com/S0264-8172(14)00251-7/sref39lko

















marine and petroleum geology - usal · expedition-01 (nghp-01) site 17 (10 45.19120n, 93 6.73650e)...

Documents