Pore water geochemistry in shallow sediments from the northeastern continental slope of the South China sea

Pore water geochemistry in shallow sediments from the northeastern continental slope of the South China sea

Accepted Manuscript Pore water geochemistry in shallow sediments from the northeastern continental slope of the South China sea Hong Ye, Tao Yang, Guo...

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Accepted Manuscript Pore water geochemistry in shallow sediments from the northeastern continental slope of the South China sea Hong Ye, Tao Yang, Guorong Zhu, Shaoyong Jiang, Lushan Wu PII:

S0264-8172(16)30067-8

DOI:

10.1016/j.marpetgeo.2016.03.010

Reference:

JMPG 2497

To appear in:

Marine and Petroleum Geology

Received Date: 30 October 2015 Revised Date:

28 February 2016

Accepted Date: 7 March 2016

Please cite this article as: Ye, H., Yang, T., Zhu, G., Jiang, S., Wu, L., Pore water geochemistry in shallow sediments from the northeastern continental slope of the South China sea, Marine and Petroleum Geology (2016), doi: 10.1016/j.marpetgeo.2016.03.010. This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting proof before it is published in its final form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.

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Pore water geochemistry in shallow sediments from the northeastern continental slope of the South China Sea Hong Yea,b , Tao Yanga,b,∗, Guorong Zhua , Shaoyong Jianga,c,∗∗, Lushan Wud,1 a State

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Key Laboratory for Mineral Deposits Research, School of Earth Sciences and Engineering, Nanjing University, Nanjing 210093, China b Collaborative Innovation Center of South China Sea Studies, Nanjing University, Nanjing 210093, China c State Key Laboratory of Geological Processes and Mineral Resources, Faculty of Earth Resources, Collaborative Innovation Center for Exploration of Strategic Mineral Resources, China University of Geosciences, Wuhan 430074, China d School of Marine Geosciences, China University of Geosciences, Beijing 100083, China e Guangzhou Marine Geological Survey, Guangzhou 510075, China

Abstract

The northeastern continental slope of the South China Sea is one of the promising areas for gas hydrates, which is characterized by its variable topography and complex tectonics. In this paper, two gravity piston cores were sampled from

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the shallow sediments in the vicinity of the region where various morphologies of gas hydrate was recovered during the second gas hydrate expedition of

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Guangzhou Marine Geological Survey. Conventionally, the identification of gas hydrate was primarily dependent on the bottom simulating reflector (BSR). Apart from BSR, the anomalies in pore water geochemistry, which are highly

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related to various microbially-mediated processes and features of fluid from deep hydrocarbon reservoir, are also helpful for the gas hydrate exploration. In this study, upward methane flux may be responsible for distinct characteristics of

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pore waters observed at the two sample sites. A reactive transport model was accordingly established to compare contrastive fluid features and to estimate the upward methane flux. Since fluid migration is typically associated with complex tectonic settings of the study area, it is proposed that heterogeneous distribution ∗ Corresponding

author author Email addresses: [email protected] (Tao Yang), [email protected] (Shaoyong Jiang) ∗∗ Corresponding

Preprint submitted to Marine and Petroleum Geology

March 8, 2016

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of gas hydrates along the continental margin should be well considered.

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Keywords: South China Sea, pore water, geochemistry, gas hydrate, methane flux, model

1. Introduction

Methane is significant to global carbon cycle, which can occur dissolved in

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pore water, as free gas, or as gas hydrate within sediment deposits along con-

tinental margins (Dickens, 2003). A considerable amount of methane is stored as gas hydrate, which is an ice-like crystalline compound formed by guest gas

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(mostly methane) within water cages under low temperature and high pressure conditions (Sloan and Koh, 2007). The extensive distribution of sediments on continental slopes thus favors the formation and accumulation of gas hydrates (Paull et al., 1994). Consequently, gas hydrates serve as a potential energy resource (Kvenvolden, 1993) and have been investigated in the offshore areas (e.g., Matsumoto, 2000; Torres et al., 2002; Paull et al., 2005; Suess, 2005). The

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amount of gas hydrate below the seafloor can be indirectly estimated. A widely used technique is the bottom simulating reflectors (BSRs) observed from reflec-

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tion seismic profiles, instantaneous amplitude and phase profiles (Hyndman and Spence, 1992; Bangs et al., 1993). It detects the interface between gas hydrate and free gas, namely, the base of hydrate stability zone (Miller et al., 1991),

1993).

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though a BSR is not necessarily associated with gas hydrate (Kvenvolden et al.,

Alternatively, numerous studies have suggested that pore water geochemistry

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of shallow sediments is linked to the presence of deep buried gas hydrate (e.g., Borowski et al., 1996; Niewöhner et al., 1998; Dickens, 2001; Snyder et al., 2007; Bhatnagar et al., 2008). This is because of an upward flux of methane that fuels anaerobic oxidation of methane (AOM) across a thin zone where downward sulfate is reduced by methane with 1:1 stoichiometry (Barnes and Goldberg, 1976; Reeburgh, 1976). The overall reaction is often written as: − − CH4 + SO2− 4 → HCO3 + HS + H2 O

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(1)

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In spite of the unclear reaction pathways, AOM generally exhausts both sulfate

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and methane in the sulfate-methane transition (SMT) zone of shallow sediments (Iversen and Jørgensen, 1985). As a result, marine sediment is a major sink of methane (Reeburgh, 2007). Assuming that the sediment is in a steady state

condition and AOM dominants sulfate consuming reactions within the SMT, the sulfate flux should balance the upward methane flux (Borowski et al., 1996).

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Anomalies in sulfate concentration gradients, as well as of other related ions can be judged in depth profiles, although this approach remains controversial (e.g.,

Berelson et al., 2005). An alternative way of sulfate consumption is organic

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matter degradation, according to

− − + 2 Corg + SO2− 4 → 2 HCO3 + HS + H .

(2)

In addition, multiple sources may be involved in competitive reactions, which should be examined carefully when quantifying species fluxes and reaction rates of pore water (Dickens and Snyder, 2009; Burdige and Komada, 2011). To solve the issue, reactive transport models are often used, especially when applied to

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environments of various flux fluids (e.g., Chatterjee et al., 2011; Hong et al.,

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2014).

Here we present pore water data of two sediment cores obtained from a ridge on the northeastern continental slope of the South China Sea, where BSRs are

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distributed widely but discontinuously (Liu et al., 2006). The distinct geochemical characteristics of different sites lead us to further discussion. By quantifying methane fluxes and AOM rates with the model, we made a comparison between

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the study sites with other previous research areas of gas hydrates, and thus discuss the reasons for different geochemical results.

2. Geological setting The South China Sea, tectonically controlled by interactions of Eurasian

Plate, Pacific Plate and Indo-Australian Plate, is one of the largest marginal basins in the western Pacific Ocean, which has experienced a complex tectonic

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evolution since the Mesozoic era (Taylor and Hayes, 2013; Shi and Li, 2012;

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Wang et al., 2014). The northwestern margin of the South China Sea is a typical passive margin while the northeastern part is an active convergent margin

(McDonnell et al., 2000; Wu et al., 2005). A series of sedimentary basins (e.g.,

the Qiongdongnan Basin, the Pearl River Mouth Basin, the Taixinan Basin,

etc), covered by sediments as thick as 10 km, are well developed (Fig. 1a) (Lüd-

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mann and Wong, 1999; Li et al., 2013b; Zhang et al., 2015), where abundant

petroleum resources have ever been discovered (Hao et al., 1996; Suess, 2005; Wu et al., 2009; Yu et al., 2009; Gong et al., 2011; Zhu et al., 2012).

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Investigation of gas hydrates in the northern South China Sea has been carried out since 1999 (Wu and Liu, 2011), and evidences of gas hydrate occurrences were also reported, including cold vents in the Dongsha Area and disseminated hydrate samples in the Shenhu Area (e.g., Suess, 2005; Han et al., 2008). More recently, a second gas hydrate expedition of Guangzhou Marine Geological Survey has drilled 13 sites and recovered a large amount of gas hydrates of various morphologies (Zhang et al., 2015).

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The study area is situated at the southwestern part of the Taixinan Basin (Fig. 1a), which is in transition from a passive margin to the west, to an ac-

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tive subduction zone to the east (McDonnell et al., 2000). It is characterized by variable topography, complex tectonics. Plenty of scarps and canyons are

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distributed along the continental slope (Yan et al., 2008; Li et al., 2013a). Sediments have been accumulating since early Mesozoic (Wu et al., 2005) and the sedimentation rate falls in one order of magnitude in a unit of cm/kyr (Wang

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et al., 2000; Zhang et al., 2015). Source rocks of high maturity of organic matter are favorable for gas hydrate formation (Chen and Pei, 1993; Yu et al., 2009). The sedimentary processes in Quaternary are controlled by two distinct groups of fault systems, NW-NWW and NNE-NEE striking respectively (Zhang et al., 2015). Faulting activity since late Cretaceous has influenced the tectonic environment which is beneficial to concentrated hydrates (McDonnell et al., 2000; Wu et al., 2007; Fuh et al., 2009). Previous studies have documented BSR distribution and geochemical anoma4

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lies in the study area (e.g., Suess, 2005; Wu et al., 2005; Liu et al., 2006; Li et al.,

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2015; Lin et al., 2016). It is considered as one of the most promising area in the northern South China Sea for gas hydrate formation and accumulation. To further explore the gas hydrate resource in this area, two ridge sites that differ

in BSR distribution, were selected. Site A is located close to the topographic

crest of the ridge where BSR is not detected, while Site B is on the western

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flank of the ridge where BSR is well developed (Fig. 1b). Table 1 shows the details of the sampling locations.

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3. Materials and methods 3.1. Sampling and analytical methods

Cored sediments were taken during R/V Haiyang-4 (Ocean IV) cruise conducted by Guangzhou Marine Geological Survey in 2012. The core lengths are 6.67 m (Site A) and 7.67 m (Site B), respectively (Table 1). The sediments in both of the two cores are lithologically characterized by greenish silt to grayish-

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green silty clay. The sediments of Site B yields a strong odor of hydrogen sulfide at the bottom.

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After extrusion from the core liner, the surface of each whole-round core samples was carefully cleaned. Pore waters samples were then collected onboard at room temperature by a vacuum extraction device from 3-cm-long core section

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at intervals of 20 cm (Cheng et al., 2012). The amounts of pore water samples are 32 (Site A) and 38 (Site B), respectively. These samples were collected

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in Corning 2-ml cryogenic vials and preserved at ∼4 ◦ C in accordance with analytical requirements. For major and trace element analyses of pore water, 2 % HNO3 was added to

prevent carbonate precipitation and microbial reaction, while for the analyses of dissolved inorganic carbon (DIC) and stable carbon isotope (δ 13 C), no HNO3

was added. For chromatography analysis, each core section was subsampled by collecting a 6-ml sediment plug with a cut-off syringe. Each sediment plug was placed into a glass serum vial and sealed with a Butyl rubber stopper and

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aluminum cap. To ensure a gas-tight seal on the stopper, 3-ml methane-free

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deionized water was quickly added and the vials placed upside-down in a −20 ◦

C freezer. All the offshore analyses were performed at State Key Laboratory

for Mineral Deposits Research, Nanjing University.

+ + 2+ The anion (e.g., Cl– , SO2– , Ca2+ ) were 4 ) and cation (e.g., K , Na , Mg

measured using the standard method of ion chromatography (Metrohm 790-1,

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Metrosep A Supp 4-250/Metrosep C 2-150). The relative standard deviation was less than 3 %.

Trace elements were measured by inductively coupled plasma mass spec-

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trometry (Finnigan Element II). Before measurement, samples were prepared by diluting in 2 % HNO3 for Ba2+ , or aqua ammonia for I– with 10 ppb of Rh as an internal standard. The analytical precisions were estimated to be < 5 % and < 2 % for Ba2+ and I– , respectively.

We adopted the new method (Assayag et al., 2006; Yang et al., 2008) to obtain simultaneously DIC concentration and its δ 13 C-DIC composition using a continuous flow mass spectrometer (Finnigan Deltaplus XP). A small amount

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of pore water sample (0.5 ml) was treated with pure H3 PO4 in a glass vial at 25◦ C. The CO2 produced was stripped with He and transferred into the mass

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spectrometer through which δ 13 C values were measured (Yang et al., 2008). δ 13 C were reported using the conventional delta notation in per mil (%) against

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Vienna Pee Dee Belemnite (V-PDB) international standard. The analytical precision of this method was estimated to be < 5 %. Methane concentrations were measured onboard using gas chromatograph

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method (Agilent 7890A). 10-ml of methane-free water was added to each sample vial to replace 10-ml headspace needed for the chromatograph injection. Samples were analyzed by headspace extraction. The precision for methane was < 5 %.

3.2. Numerical model A reactive transport model based on aquatic mass balance was established to quantify the geochemical processes in pore waters, especially diagenetic re6

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actions coupled with fluid migration (e.g., Berner, 1980; Boudreau, 1997). In

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this study, the finite volume method was used for discretization, and multiple nonlinear equations of different chemical species were iterated with the operator splitting method (Strang, 1968; Karlsen et al., 2001; Versteeg and Malalasekera,

2007). The model was explained in Appendix A, and its basic parameters were listed in Table 2.

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Due to lack of data of the study area, we configured the model by analogy

with adjacent sites whose parameters were available in literatures (e.g., ODP Leg 184), although we admitted that the geological conditions actually varied

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from region to region. Porosity (φ) is calculated from Site 1144, ODP Leg 184 (Wang et al., 2000) using an exponential fitting. Tortuosity (θ) is dependent on the porosity as

θ2 = 1 − ln φ2 .

(3)

The geothermal gradient in the study area is about 50 ◦ C km−1 (Li et al., 2015), which is not sensitively reflected with the model depth. Therefore, constant seawater diffusion coefficients are assumed. The model includes sulfate (SO2– 4 ),

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dissolved methane (CH4(aq) ), bicarbonate (HCO–3 ), calcium (Ca2+ ) and magne-

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sium (Mg2+ ). In shallow sediments where pH is between 7.1∼8.1, alkalinity is often treated as carbonate alkalinity by ignoring HS– and other minor species (Chatterjee et al., 2011; Sauvage et al., 2014). For the same reason, carbonate

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alkalinity is dominated by HCO–3 . The parameter to be determined is the velocity of upward fluid flow (u, Table 2), which represents advective flux of mass from bottom of the model domain. We would verify whether it is the fact that

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Site A is diffusion-dominated while Site B primarily differs from it in elevated u.

4. Results

The concentrations of selected major and trace ions as well as alkalinity

and the δ 13 C value of DIC are illustrated in Fig. 2, and full data are included

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in Appendix B. The depth is reported in the unit of meter below the seafloor (mbsf).

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As depicted in Fig. 2, the concentration distributions of Cl– , Na+ and K+

are basically constant throughout the core depths within the range of seawater salinity at both sites. However, distinct characteristics of pore water are observed from other depth profiles.

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For Site A, gradual decline in Mg2+ and Ca2+ concentrations is almost linear

with the downward depth. Ba2+ keeps its trace concentration except that a few peak values appear in the middle of the core as high as about two- to fourfold

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of that near the seafloor. The concentration of SO2– 4 decreases by 35% from top to the bottom of the core, expressing a linear depth profile. In contrast, I– concentration rises up to eight times the value at 0.27 mbsf with the increasing depth. Increase in alkalinity is a similar trend, but with a small multiplier. δ 13 C value of DIC is negative, suggesting that pore water is generally depleted in δ 13 C. Moreover, the δ 13 C value becomes more negative with respect to depth. However CH4 concentration is at a relatively lower level when compared to that

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of Site B.

Although, both of Site A and Site B generally show similar trend in depth

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profiles, the linear concentration gradient of Site B is larger than that of Site A, as illustrated in Fig. 2d, e, h, i and j. A concave-up SO2– 4 profile is worth

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noting, which shares a linear part above approximately 2 mbsf with that of Site A but deviates from its remaining linear part. The concentration of SO2– 4 of Site B hereafter decreases rapidly and approaches the zero value at the base of the

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core, where the CH4 concentration is detected in mM magnitude (Fig. 2g). It is interesting that there is an inflection in the Ba2+ profile, which is indicative of Ba2+ diffusing from greater depth. The concave-up SO2– 4 profile at Site B may be attributed to near-surface

bioactivities or transient fluid redistribution in pore water (Hensen et al., 2003; Coffin et al., 2008). Although the trend is nonlinear, we could obtain the SMT depth by extrapolating the linear part of the SO2– 4 profile (Fig. 2g). The inferred SMT depths are ∼20 mbsf for Site A and ∼9 mbsf for Site B, respectively. 8

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The results of pore water geochemistry suggests that shallow pore water

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anomalies are consistent with the BSR distribution. Site B, which is within the BSR area (Fig. reffig:geosetb), exhibits steeper gradients, greater curvatures and inferred shallower SMT depth than Site A, where there is no BSR sign. Ac-

cordingly, interpretation of pore water anomalies is an alternative proxy during

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gas hydrate exploration.

5. Discussion

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5.1. Salinity and halogen ions

As a sensitive geochemical proxy, pore water is used to identify various diagenetic processes in the marine sediments (Borowski et al., 1996; Yang et al., 2013). Pore water anomalies are widely observed, and are considered to be related to gas hydrate occurrence (e.g., Ussler III and Paull, 1995; Egeberg and Dickens, 1999; Hesse, 2003; Wallmann et al., 2006; Wu et al., 2013). The concentration of Cl– in seawater is generally proportional to the salin-

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ity. Pore water Cl– is conservative because it is rarely involved in diagenetic reactions. However, it could respond to changes of thermodynamic processes

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such as gas hydrate formation/dissociation (Hesse and Harrison, 1981), fluid migration through faults (Hooper, 1991), or global climate variation (Adkins et al., 2002). In the first case, gas hydrate removes salts during its formation,

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which results in a high salinity anomaly. Conversely, if the temperature and pressure could not sustain the stability condition for gas hydrate (e.g., during core recovery), it will dissociate and release fresh water. Thus excursion from

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the Cl– concentration baseline occurs (Hesse and Harrison, 1981; Egeberg and Dickens, 1999). Chloride anomalies are reported at the gas hydrate areas e.g., the Blake Ridge and the Cascadia margin (Egeberg and Dickens, 1999; Torres et al., 2004). Chloride concentrations at Site A and B are fluctuating around the seawater value (∼560 mM), despite of greater dispersion at the base of the core. There is no strong evidence to support the presence of gas hydrate (if

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any) within the sediment of drilling depth, and/or the shallow sediments are not affected by fluids of salinity anomaly (if any).

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Concentration gradient anomalies are found in I– profiles. Iodine plays an

important role in organic compounds. It is one of the most biophilic halogens,

enriched in ocean phytoplankton and algae (Elderfield and Truesdale, 1980). Usually, I– enrichment in pore water is caused by mineralization of organic mat-

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ter (Muramatsu and Wedepohl, 1998). As a result, oceanic I– content depends

on the amount of buried organic compounds and the rate of organic matter degradation (Martin et al., 1993; Tomaru et al., 2009). Both of the two sites

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exhibit marked increasing I– concentrations in the depth profiles (Fig. 2h). We calculated their fluxes with the formula of

F = −φDs

and

Ds =

∂C ∂z

Dsw θ2

(4)

(5)

where Dsw denotes the diffusion coefficient in the seawater obtained from ther-

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modynamic equations of state (e.g., Boudreau, 1997); Ds is the diffusion coefficient corrected with sediment parameters such as porosity (φ) and tortuosity

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(θ); and ∂C/∂z is the gradient of concentration (C) with depth (z). The dependency of θ is given in Equation (3). We estimate the gradients using linear

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regression fitting, and assume that average porosity of shallow sediment is 0.75 (Yang et al., 2010). The diffusion coefficients Ds,A (I− ) and Ds,B (I− ) are 347 cm2 yr−1 and 324 cm2 yr−1 for Site A and Site B, respectively. Therefore, I–

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flux of Site A is 4.6 × 10−3 µmol cm−2 yr−1 . It is less than a half of the value of Site B, 10.3×10−3 µmol cm−2 yr−1 . Both of these values are comparable to that at Site 997 of the Blake Ridge, (7.2 × 10−3 µmol cm−2 yr−1 , data from Egeberg and Dickens (1999)), Site 1145 of the Hydrate Ridge (7.6×10−3 µmol cm−2 yr−1 , data from Lu et al. (2008)), or Tokai of Nankai Trough (7.8 × 10−3 µmol cm−2 yr−1 , data from Muramatsu et al. (2007)). There could be two controversial

explanations for the resulting high I– flux at Site B. One is the release of a large quantity of I– due to organic matter degradation in the deep buried sediments 10

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(i.e. ex-situ), which may also serve as a hydrocarbon source (Tomaru et al.,

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2009). Then upward fluid takes the dissolved species to shallow sediments. Alternatively, microbial activity in surface sediments is high enough to decompose the organic matter and thus produce in-situ I– (Kennedy and Elderfield, 1987; Martin et al., 1993). However, the core length is beyond the depth of I– -IO–3

cycling and the steep concentration gradient indicates a source deep beneath

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the core. On that account, we believe that the reason of ex-situ I– -bearing fluid migration is plausible.

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5.2. AOM and the SO2– 4 flux

Sulfate is supplied from the seawater and sinks within the marine sediments (Reeburgh, 1976). In shallow sediments of the continental slope, diagenetic reactions are taking place according to the sequence of free energy yielding for organic matter oxidation (Froelich et al., 1979; Berner, 1980). Typically, SO2– 4 is reduced by organic matter in the so-called sulfate reduction zone and its concentration declines gradually with depth (Henrichs and Reeburgh, 1987;

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Pohlman et al., 2008; Arndt et al., 2013). However, if CH4 exists, a microbiallymediated efficient reaction, i.e. AOM, is functioning in the SMT just at the base

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of the sulfate reduction zone, where SO2– 4 and CH4 are consumed at significantly high rate. (Reeburgh, 1976; Henrichs and Reeburgh, 1987; Boetius et al., 2000). This process results in steep and linear concentration gradients of both reactants

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near the SMT. Buried organic matter will further degrade into small molecules such as H2 , CO2 or CH4 under anaerobic environments where SO2– 4 is exhausted

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(Claypool and Kvenvolden, 1983; Regnier et al., 2011). Fermentation of organic matter, namely, methanogenesis produces CH4 , which is an important gas source to hydrate formation (Whiticar et al., 1986). Methane in deep sediments will migrate towards the seafloor with upward fluid flow, driving the shoaling of the SMT (Dickens, 2001; Chatterjee et al., 2011). The SMT depth is therefore linked to the intensity of AOM, which is conducted by microbial consortium (Hoehler et al., 1994; Boetius et al., 2000) and fueled by the upward CH4 flux, as well as the downward SO2– 4 flux (Dickens and Snyder, 2009). The cross-plot 11

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of the concentrations of I– and SO2– 4 yields correlation coefficients of 0.82 for

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Site A and 0.97 for Site B, respectively (Fig. 3). Others sites along the northern continental slope of the South China Sea exhibit similar correlation (Yang et al., – 2013, see Fig. 3). It is evident that SO2– 4 consumption, as well as I enrichment,

is probably relevant to degradation of organic matter.

According to Reaction (1), fluxes into the SMT of two reactants should be

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equal if AOM is the dominant reaction within the SMT, and if no SO2– 4 or CH4

could penetrate the SMT (Borowski et al., 1996; Niewöhner et al., 1998). With the linear part of the depth profile used to SMT inference, we obtained SO2– 4

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2− gradients of both sites. The diffusion coefficients Ds,A (SO2− 4 ) and Ds,B (SO4 )

are 175 cm2 yr−1 and 162 cm2 yr−1 for Site A and Site B, respectively. By applying Equation (4) and (5), calculated fluxes of SO2– 4 into the SMT are consistent with data of neighboring sites (Table 3). Site A (14.7 mmol m−2 yr) is comparable to Site C14 of the Xisha Uplift (Luo et al., 2013), higher than Site T1 of the Qiongdongnan Basin (data from Wu et al., 2011). The flux is within the range of the Shenhu Area (Wu et al., 2013). Site B is characterized

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−2 by its relatively higher SO2– yr, which falls in the range 4 flux of 41.5 mmol m

between 81.0 mmol m−2 yr (Yung-An Ridge of the Taixinan Basin, Lim et al.,

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2011) and 31.6 mmol m−2 yr (the Dongsha Area, data from Cao and Lei, 2012). It is worth noting that an accretionary prism is location to the east of the study

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area (Fig. 1a) and SO2– 4 fluxes reported at the active margin are usually much higher (e.g., Treude et al., 2005, Table 3). Statistically, variation in flux is, however, often large in the same area, especially when increasing the number of

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comparative samples. For example, Wu et al. (2013) examined 53 sites in the −2 Shenhu Area, where SO2– yr. 4 fluxes vary from 2.0 to 26.9 mmol m

The concentration of Ba2+ markedly rises up to about four times of the

seawater value at the base of the core of Site B, where it is close to the depth of SO2– 4 depletion. In the continental margin settings, Ba can precipitate as authigenic barite which has a very low solubility in water (Torres et al., 1996). Sediment burial drives barite downward to SO2– 4 -depleted pore water where it dissolves (Brumsack and Gieskes, 1983; Dickens, 2001). Elevated Ba2+ con12

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centration implies upward dissolved Ba2+ diffusing from just below the SMT

concentration is a supportive evidence of the SMT depth. 5.3. Alkalinity and δ 13 C-depleted DIC

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depth (Dickens, 2001; Snyder et al., 2007). As a result, a sharp increase in Ba2+

Alkalinity is a chemical measurement of the capacity of pore water to neu-

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tralize acids. It is highly affected by microbial activity in pore water (Berner et al., 1970), especially AOM (Nauhaus et al., 2005). Organic carbon is oxidized during microbial degradation (Arndt et al., 2013), turning into inorganic carbon,

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mostly DIC. The products of AOM also include DIC, which brings considerably additional alkalinity into pore water at the SMT (Boetius et al., 2000; Snyder et al., 2007). The total effect of these processes on pore water is a noticeable increase in alkalinity gradient between the seafloor and the SMT depth, as observed in this study (Fig. 2i). Consequently, there is often a maximum in the alkalinity profile within the SMT depth, reflecting intensive microbial activities (Malinverno and Pohlman, 2011; Chatterjee et al., 2011).

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High alkalinity is beneficial to precipitation of authigenic minerals such as authigenic carbonates and pyrites (Boetius et al., 2000; Jørgensen et al., 2004;

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Snyder et al., 2007). Authigenic carbonates accumulate in the solid phase by removing DIC, and Ca2+ or Mg2+ , which are discovered along the continental margins (e.g., Paull et al., 1995; Bohrmann et al., 1998; Chen et al., 2005; Treude

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et al., 2005), including the study area (Han et al., 2008; Tong et al., 2013, marked in Fig. 1b). Such carbonates reveal the correlation between authigenic mineral

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diagenesis and hydrate-bearing sediments, as is evidenced by tubular carbonates (Bohrmann et al., 1998; Greinert et al., 2001). Seep carbonates found in the vicinity of Site A are characteristic of the large fraction of high magnesium calcites and aragonites of tubular or chimney morphology (Han et al., 2008). This might support the decline trend in concentrations of Ca2+ and Mg2+ in this

study area. Moreover, the morphological evidence indicates possible conduits for fluid transport. An average δ 13 C value of -45.7 % of seep carbonates was discussed in pre13

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vious work (Tong et al., 2013), implying that ambient pore water DIC during

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mineralization is much depleted in δ 13 C. It is in agreement with the δ 13 C value of DIC from pore water of Site B. It is convenient to determine carbon sources with various δ 13 C values. Organic matter contains carbon typically ranging from

-30 to -20 %(Holland and Turekian, 2010), comparing to that of thermogenic methane (>-50 %) (Whiticar, 1999). Fractionation of 13 C during methanogen-

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esis is relatively greater (ε=∼60 %, Botz et al., 1996). Therefore, it produces δ 13 C-depleted CH4 and δ 13 C-enriched DIC. In contrast to the elevated alkalinity with respect to depth, the corresponding δ 13 C-DIC ratio goes negative

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downcore (Fig. 2j). Site B has a wide range of δ 13 C values from the seawater value (∼0) to highly negative -42 % near the inferred SMT. A steep gradient of δ 13 C-DIC in the shallow pore water is associated with SMT shoaling or more negative δ 13 C-DIC from in-situ CH4 oxidation, which may be due to intense AOM (Kastner et al., 2008). The value of Site A varies from -5 to -25 %, which seems derived from organic matter degradation, rather than AOM. However, as Chatterjee et al. (2011) pointed out, this could occur by ascending fluid of

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δ 13 C-enriched DIC sourced from methanogenesis below the SMT. Thus the DIC concentration is complicated by coupling multiple sources and reactions (Chat-

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terjee et al., 2011; Hong et al., 2014). As illustrated in Fig. 4, the source of pore water DIC actually consists of four different parts: organic matter degradation

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(δ 13 C = -20 ∼ -30), methanogenesis (δ 13 C-enriched, may great than zero), CH4 through AOM, and external fluid. The last two parts again depends on ex-situ pore water environments.

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On the whole, the particularly negative δ 13 C value of Site B, together with other pore water anomalies, could support our inference of shallow SMT depth. While for Site A, it is more or less problematic. Fig. 5 provides a comprehensive comparison of two sites. 5.4. Upward CH4 flux It is mentioned above that the linear or concave-up SO2– 4 concentration gradients allow us to believe the existence of AOM, which is fueled by upward CH4 14

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flux. Therefore it is often related to the amount of deep petroleum reservoir

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(Kvenvolden, 1993; Tréhu et al., 2003). A simple estimate of upward CH4 flux into the SMT is calculating its gradient directly if possible, or alternatively using SO2– 4 flux from the seafloor by assuming a 1:1 stoichiometry due to AOM (Borowski et al., 1996). However, it is only valid in a close system, because

external flux from below could change species in shallow pore water (Dickens

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and Snyder, 2009; Burdige and Komada, 2011; Chatterjee et al., 2011). In this study, we determine CH4 flux with a reactive transport model (Appendix A). A linear gradient of SO2– 4 concentration reflects steady state of species redis-

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tribution (Borowski et al., 1996). Site A is accordingly considered as a reference site of no advection, from which the other case differs by increasing advective velocity (i.e., external CH4 advective flux) at the base of the sediment column (Fig. 6). Parameters are given in Table 2. Following discussion is based on the model result.

The resulting CH4 fluxes into the SMT are 14.6 mmol m−2 yr and 28.0 mmol m−2 yr at Site A and Site B, respectively, which are less than the result from

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linear SO2– 4 concentration gradients (Table 4). It is possibly attributed to a fraction of SO2– 4 reduced during organic matter degradation, and CH4 produced

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from methanogenesis is limited because of the burial depth. However, the result of SMT depth according to SO2– 4 concentration gradient is acceptable as a rough

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estimate, especially when shallow sediment cores do not penetrate the SMT, as in this study.

Contribution of AOM within the SMT is dominant, which is in accordance

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with our previous assumption. The depth-integrated AOM rate is consistent with other results (e.g., measured rates in Table 3), higher than model results of the Blake Ridge (10.5 mmol m2 yr−1 , Wallmann et al., 2006) but lower than that of the Hydrate Ridge (4500 mmol m2 yr−1 , Luff et al., 2005). Compared to Site A, the shallower SMT at Site B accompanied with a higher

reaction rate. If the sedimentation rate and organic carbon content are fixed, that is because more organic matter is buried below the SMT, where more insitu CH4 is produced. This biogenic CH4 , together with external flux charged by 15

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deep petroleum reservoir, migrates into the SMT (Snyder et al., 2007; Dickens

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and Snyder, 2009). Although other processes might contribute to carbon budget in marine sediments (e.g., Hong et al., 2013; Komada et al., 2015), our model demonstrates that external CH4 flux could significantly drive the shoaling of the SMT.

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5.5. Migration pathways of fluids

Previous seismic surveys suggest that patchy yet prominent BSRs are distributed in our study area (Liu et al., 2006; Li et al., 2015), and the thickness of

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gas hydrate stability zone is calculated accordingly (Li et al., 2013b). However, gas hydrate occurrence is far more complicated than what seismic profiles can predict. It also depends on thermodynamic stability conditions, as well as fluid migration, which is controlled by regional tectonic settings and local geological structures. The study area is situated in a complex tectonic setting where the passive continental margin is significantly influenced by the accretionary prism (McDonnell et al., 2000, see the highlighted small map in Fig. 1a). Faults and

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diapirs are well developed during geological time, presenting variable submarine topography with scarps and canyons (Li et al., 2013a; Zhang et al., 2015). It

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is reported that these faults cut through newer deposits, and even extend to the seafloor (Wu et al., 2009). They act as migration pathways for upward CH4 from deeply buried organic-rich sediments accumulated during Pliocene

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and Pleistocene (Fuh et al., 2009; Li et al., 2013a). The seafloor topography of Site A is relatively gentle. In contrast, the topography gradients around Site

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B are highly steep with a submarine canyon deeply incised to the west. If it is true, perhaps gas is trapped in the reservoir whose upper boundary is parallel to the gentle seafloor, but migrates to the slope along conduits of local geological structures (e.g., strata, faults or diapirs). Inferred from the extensively distributed BSRs, gas hydrate spreads across

two distinctive tectonic provinces (Liu et al., 2006). Tectonism in the study area is not so intensive as on the accretionary prism offshore southwest Taiwan. This favors the deposit of thick sediments over considerable geological time (McDon16

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nell et al., 2000). A plenty of fault-controlled morphologies are well developed,

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providing conditions to fluid migration. Although, it seems to be disadvantageous to accumulation of petroleum resources, gas hydrates are still able to form under proper conditions and be concentrated along those pathways. A variation

in CH4 fluxes might reflect lithologic heterogeneity of marine sediments (Chatterjee et al., 2014), which eventually modifies geochemical features of shallow

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pore waters. Consequently, it is indispensable to integrate multiple exploration methods.

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6. Conclusions

Through analysis and discussion of the pore water data sampled from the northeastern South China Sea, we made conclusions as follows: 1. Two cores were sampled from shallow sediments on a ridge of the northeastern South China Sea, where patchy BSRs were mapped. The results of pore water geochemistry suggests that shallow pore water anomalies are

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consistent with the BSR distribution. Therefore, pore water geochemistry is an important proxy for gas hydrate exploration.

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2. Pore water geochemistry reflects diagenetic reactions taking place in hydrate bearing sediments. The SMT depth inferred by SO2– 4 concentration gradient is acceptable as a rough estimate, especially when shallow sedi-

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ment cores do not penetrate the SMT. The sharp increase in Ba2+ concentration usually is a precursor of the SMT, which is in correspondence with

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13 the SO2– 4 profile. The ratio of δ C is another important proxy indicating

the CH4 sources, as well as microbially-mediated reactions.

3. Assuming that reactions in shallow sediments are mostly associated with CH4 supplies from deep gas reservoir, we established a reactive transport model. It yields reasonable estimates of upward CH4 fluxes, fitting to data results. Modeled AOM rates are comparable to other sites of continental margins. The model demonstrates a causative link between shallow pore water and deep gas amount. 17

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4. The study area belongs to a passive margin involved in complex tectonic

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settings. Variable fluxes might result from heterogeneous sediment lithology, which is largely related to the occurrence of gas hydrates. The heterogeneity in the study area needs further exploration, with combination of geochemical and geophysical methods.

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Acknowledgments

We thank the crew and scientists of R/V Haiyang-4 (Ocean IV) for taking core samples. We also thank Prof. Gerald (Jerry) Dickens, one anony-

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mous reviewer and Associate Editor Dr. Constantin Cranganu for their helpful comments that greatly improved the earlier version of this manuscript. This study was supported by the National 973 Project (Grant No. 2013 CB835000), the National Natural Science Foundation of China (Grant Nos. 41230102 and 40903002) and a National Key Scientific Project organized by the China Geological Survey (Grant No. GZH201200305-06-04).

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Appendix A

In our model, the one-dimensional form of dissolved species of pore water is

adopted as the basic equation, which is given by   X ∂(φC) ∂(φvC) ∂ ∂C + = (φDs ) + φα(C0 − C) + φ R ∂t ∂x ∂x ∂x 32

(6)

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where C is the concentration of a dissolved species in terms of mass per unit

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volume of pore water (C0 at the sediment water interface); φ is the porosity of the sediment; Ds is molecular diffusion coefficient in sediment, including the

effects of tortuosity; and the irrigation is introduced as a source/sink term with a depth-dependent irrigation coefficient α (Boudreau and Westrich, 1984).

The last term of the equation is named reaction term. For organic matter

1991) with the form of RG = k(t) · G

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and

SC

degradation, we adopted a reactive continuum model (Boudreau and Ruddick,

k(t) = ν · (a + t)−1

(7)

(8)

where a denotes the apparent initial age of the initial organic matter mixture and ν is the shape parameter of the organic compound distribution. Thus sulfate reduction (s.r.) with organic matter can be simplified to a Michaelis-Menten rate form as



CSO4 KSO4 ,half + CSO4



(9)

D

Rs.r. = RG ·

TE

and similarly methanogenesis (m.g.) is modeled as regular degradation using Rm.g. = RG

(10)

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Methanogenesis will not take effect until the concentration of sulfate is below a threshold value (Reeburgh, 1976). The AOM within the SMT is modeled as reviewed in Regnier et al. (2011)

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taking the form

Raom = Raom,max ·



CCH4 KCH4 ,half + CCH4



CSO4 KSO4 ,half + CSO4

 (11)

where KCH4 ,half or KSO4 ,half is the half saturation constant of corresponding species and Raom,max denotes the maximum reaction rate. A modified diagenetic equation suggested by Turchyn and DePaolo (2011) is employed when describing carbonate precipitation which could be written as

33

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that ∂C ∂2C + f M Cs Rcarb = ∂t ∂x2

(12)

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(1 + Kads )

where Cs is the concentration of the solid phase; Kads is the distribution coeffi-

cient of the ratio of Cs to Cf ; Rcarb herein denotes the rate of recrystallization; M = ρs (1 − φ)/ρf φ is the mass ratio function of the depth, and f is defined as

the volume fraction of carbonate in the sediment (Turchyn and DePaolo, 2011).

SC

The rate of recrystallization is based on the exponential fitting model (Richter and Liang, 1993)

Rcarb = a + (kcp − a) · exp(−bx)

(13)

are both tunable parameters.

Appendix B

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where kcp is kinetic constant of recrystallization rate at the seafloor, and a, b

Appendix B includes full data of analytical results of Site A (5) and Site B (6).

Seafloor temperature (◦ C)

Water depth (m)

Sampling method

Core length (m)

Amount of pore water samples

Site A

907.19

Gravity piston core

4

6.67

32

Site B

1606.51

Gravity piston core

2

7.67

38

TE

D

Site

Table 1: Information of sites where pore waters were sampled from the north-

EP

eastern South China Sea.

a Measured

values onboard. from water depth. c Porosity is a function of the depth z, which is φ = φ ∞ + (φ0 − φ∞ ) exp(−κφ z). d According to Site 1144, ODP 184 (Wang et al., 2000). e Wang et al. (2000) and Zhang et al. (2015). f Estimated in the model. g Assumed in the model. h According to Site 1144, ODP 184 (Wang et al., 2000). i Approximate to measured values offshore. j The diffusion coefficient is dependent on the salinity, temperature and pressure of the

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b Calculated

water, which can be calculated using equations of state (Boudreau, 1997). k Reviewed in Arndt et al. (2013). l Suggested in Wegener and Boetius (2009). m According to Turchyn and DePaolo (2011).

34

0

RI PT Location of the sample site of this study

-750 -1000 -1250 -1500 -1750

600

ODP Leg 184, Site 1144

Area of the 2nd gas hydrate expedition of Guangzhou Marine Geological Survey

Passive continental margin

Thrust front

118°E 114°E

116°E 20°N

18°N 112°E

rl Pea

Qiongdongnan Basin

n asi hB out erM Riv

Bathymetry

This study area (Fig. 1b)

Basin boundary

120°E

0

BijiananBasin

Sunda Block

in

s nBa

a Taixin

Hongkong

22°N

24°N

a

EP

Guangzhou

TE

China

D

Xiamen

South China Block

Ea ste Se rn Ba a sin

ary tion cre ism c A Pr

AC C

26°N

122°E

200 km

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n

Taiw a

Taipei

-2500

b

-2250

SC

-2000

Area of BSR distribution

-500

-250

5 km 0

Site B

Site A

-1600

0

20

-1

Water depth (m)

0

0 -16

0 -80

Seepage site

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Figure 1: Location of the study area. (a) The geological setting of the northeastern South China Sea. The study area (red star) is situated at the southwestern part of the Taixinan Basin, which is close to Site 1144 of ODP Leg 184 (red dot) and area of the second gas hydrate exploration of Guangzhou Marine Geological Survey (red box). Highlighted box around the study area is a tectonic map showing structural elements modified from McDonnell et al. (2000). (b) The two sample sites (round mark) are located on the same ridge, where methane-associated seepage is reported (triangle mark). Bottom simu35 lating reflector (BSR) distribution is mapped in the study area (shaded area). The bathymetric map is according to Suess (2005).

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550

0

1

1

2

2

3

3

4

4

5

5

6

6

7

a

8

0

200

20

Mg (mM)

40

Na (mM) 400

500

60

0

600

0

7

Na (Site A) Na (Site B)

b

5

Ca (mM) 10

15

20

0

4

4

5

5

EP

0

3

6

6

8

8

D TE

7

Mg (Site A) Mg (Site B)

AC C

d

7

e

c

Ca (Site A) Ca (Site B)

0

1

7

8

f

Figure 2: Analytical results of pore water chemistry. Red dots denote the concentrations or isotope composition at Site A and blue dots are referred to as those at Site B. (a) chloride; (b) iodide; (c) potassium and sodium; (d) sulfate and dissolved methane; (e) alkalinity; (f) δ 13 C component of dissolved inorganic carbon (DIC); (g) magnesium; (h) calcium; (i) barium.

36

16

K (Site A) K (Site B)

8

3

6

14

6

2

5

12

5

2

4

K (mM)

4

1

3

10

3

1

2

8

2

8

0

300

1

7

Cl (Site A) Cl (Site B)

1

Depth (mbsf)

600

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500

SC

Cl (mM)

450

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Depth (mbsf)

0

400

Ba (μM)

2

3

4

Ba (Site A) Ba (Site B)

5

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2

(Site (Site (Site (Site

40

A) B) A) B)

0

4

4

5

5

6

6

60

80

I (Site A) I (Site B)

7

g 0

500

0

5

1000

1500

CH4 (ppm)

2000

Alkalinity (mM) 10

15

2500

20

Alkalinity (Site A) Alkalinity (Site B)

1

TE

3

4

0

h

-50

-40

δ13C-DIC (‰ V-PDB) -30

-20

-10

0

10

2

3

4

5

EP

5

8

1

D

2

6

7

7

AC C

6

8

40

M AN U

3

0

I (μM)

20

2

3

8

0

1

7

Depth (mbsf)

30

RI PT

20

SO4 SO4 CH4 CH4

1

Depth (mbsf)

SO4 (mM)

10

SC

0

0

i

8

j

d13C-DIC (Site A) d13C-DIC (Site B)

Figure 2: Analytical results of pore water geochemistry. Red dots denote the concentrations or isotope values at Site A and blue dots are referred to those at Site B. (a) chloride; (b) sodium; (c) potassium; (d) magnesium; (e) calcium; (f) barium; (g) sulfate and dissolved methane; (h) iodide; (i) alkalinity; and (j) δ 13 C value of dissolved inorganic carbon (DIC).

37

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35

SO4 vs I (Site A) Linear fit (Site A) SO4 vs I (Site B) Linear fit (Site B) SO4 vs I (HS-A) Linear fit (HS-A) SO4 vs I (HS-B) Linear fit (HS-B) SO4 vs I (HQ-1PC) Linear fit (HQ-1PC)

SC

30

20

R2=0.8214

15

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SO4 (mM)

25

R2=0.9859

10 5

20

40

D

0

R2=0.9796

60

I (μM)

80

R2=0.9825

100

120

TE

0

R2=0.9661

– Figure 3: Cross-plot of sulfate (SO2– 4 ) vs iodide (I ). There is a good correla-

tion between sulfate and iodide concentrations. A comparison is made among

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different locations on the continental slope of the northern South China Sea. The correlation coefficients are labeled closed to the fitting lines, together with

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corresponding annotations. Sites HS-A and HS-B are situated in the Shenhu Area, southeast of the Pearl River Mouth Basin. Site HQ-1PC is located in the northwest slope of the Qiongdongnan Basin. The data are taken from Yang et al. (2010) and Yang et al. (2013).

38

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seawater sediment

SO4 reduction 1

3

2

M AN U

DIC (?)

SC

AOM

Methanogenesis

SMT

Corg (-20 ~ -30)

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SO4

4

ex-CH4 (?)

bio-CH4 (-110 ~ -60)

ex-DIC (?)

D

bio-CH4 (-110 ~ -60)

TE

thermo-CH4 (-50 ~ -20)

Figure 4: Cartoon of carbon cycle within shallow sediments. Organic matter (Corg ) contains carbon of moderate δ 13 C value. Above the SMT, microbially-

EP

mediated degradation with sulfate (SO2– 4 ) yields dissolved inorganic carbon (DIC) (Source 1). When organic matter is buried beneath the SMT, methanogenesis with large isotope fractionation could produce δ 13 C-enriched DIC

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(Source 2), as well as δ 13 C-depleted biogenic methane (bio-CH4 ). This in-situ bio-CH4 migrates upward and oxidized be (SO2– 4 ) within the SMT, which is named anaerobic oxidation of methane (AOM) (Reeburgh (2007)). The DIC derived from AOM contributes to the DIC pool (Source 3). Additionally, external fluid flow brings deep DIC (Source 4). Taking external methane (ex-CH4 ) into consideration, various isotope compositions in biogenic and thermogenic methane (thermo-CH4 ) complicate carbon cycle.

39

0

Ba (μM)

1

0

2

10

3

SO4 (mM) 20

4

30

5

0

40

0

Ba (μM)

1

2

10

1

3

4

5

6

SO4 (mM) 20

5

30

40

SO4 (mM) Ba (μM) δ13C-DIC (‰V-PDB) Alkalinity (mM)

0

-40

-30

-20

-10

δ13C (‰ V-PDB)

0

TE

-50

D

7

8

4

M AN U

Depth (mbsf)

2

3

SC

0

RI PT

ACCEPTED MANUSCRIPT

5

10

15

Alkalinity (mM)

10

-50

20

0

-30

-20

-10

δ13C (‰ V-PDB) 5

10

15

Alkalinity (mM)

0

10

20

b

EP

a

-40

Figure 5: Comparison of pore water anomalies in shallow sediments. The particularly negative δ 13 C value, sharp increase in barium (Ba2+ ) concentration,

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and steep concentration gradients in sulfate (SO2– 4 ) and alkalinity, could support

the inference of shallow SMT depth. (a) Site A. The SMT might be at great depth; (b) Site B. A shallow SMT is inferred just beyong the core length.

40

ACCEPTED MANUSCRIPT

Concentration (mM)

5

20

30

40

Concentration (mM)

50

60

5

SO4 Mg

10

10

20

30

40

50

60

SMT

10

M AN U

Depth (mbsf)

Ca

HCO3

0

0

RI PT

10

SC

0

0

15

15

SMT

20

Site A

a

25

CH4

b Site B

EP

25

TE

CH4

D

20

Figure 6: Results of the numerical model. Different dissolved species are denoted in different colors. (black: methane; green: sulfate; pink: bicarbonate;

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yellow: calcium; blue; magnesium.) Black arrows in the bottom represent relative amount of external methane flux. The SMT depth is marked with a gray dashed line. (a) Model of Site A. The SMT depth is estimated at about 18 mbsf. No advective flux of methane is supplied in this case. (b) Model of Site B. The SMT depth is estimated at about 9 mbsf. In this case, elevated advective flux of methane drives SMT shoaling.

41

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Description

T

Temperature at the seafloor

Unit ◦C

4a

Site B 2a

P

Pressure

atm

90.98b

160.32b

S

Salinity

PSU

34

L

Length of sediments Porosityc at the seafloor

m

φ0 φ∞

Porosity at great depth

κφ ω

Attenuation coefficient for porosity

u

Velocity of upward fluid flow

m yr−1 m yr−1

α0

Irrigation at the seafloor

yr−1

κα

Attenuation coefficient for irrigation

Kads

Adsorption coefficient

TOC CSO ,0 4

Total organic carbon 2– SO4 concentration

CCH ,0 4

CH4 concentration

mM

– HCO3 concentration

mM

CCa,0 CCa,s /CCa

Ca concentration

mM

CMg,0 CMg,s /CMg

Mg concentration

FCH ,ex 4

Diffusive CH4 flux at the bottom

mol m−2 yr−1

0.006g

ρs

Grain density

ρf sw DSO 4 sw DCH 4 sw DHCO 3 sw DCa sw DMg ia

Fluid density 2– SO4 diffusion coefficient

kg m−3 kg m−3

2500g 1030g

KSO ,half 4 KCH ,half 4 Raom,max

1.0h 29i

SC

wt% mM

0i

2.8i

M AN U

12i

mM

2000 48i

47i

1000

m2 yr−1

1.75e2j

1.62e-2j

CH4 diffusion coefficient

m2 yr−1

2.84e-2j

2.66e-2j

– HCO3 diffusion coefficient

m2 yr−1

1.85e-2j

1.70e-2j

Ca diffusion coefficient

Mg diffusion coefficient

m2 yr−1 m2 yr−1

1.30e-2j 1.21e-2j

1.20e-2j 1.13e-2j

Initial age of organic matter

yr

Decay coefficient of organic matter 2– Half saturation constant of SO4

2.0e4 0.16k

mM

0.5l

Half saturation constant of CH4

mM

2.0l

Max rate of anaerobic oxidation

mM yr−1 yr−1

0.18f 1e-6m

Kinetic constant of carbonate precipitation

TE

kcp

1.4e-3f

0.10g 1.5

D

ν

5e-4e 5e-4f

0.1

Ca distribution ratio

Mg distribution ratio

25 0.80d 0.55d 0.01

Sedimentation rate

CHCO ,0 3

Site A

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Symbol

Table 2: Parameters and boundary conditions used in the model.

EP

FSO4 (mol m−2 yr−1 )

FCH4 (mol m−2 yr−1 )

Raom n (mol m−2 yr−1 )

zSMT (mbsf)

0.0147

∼20

Taixinan Site B (This study)

0.0415

∼9

Taixinan (Lim et al. (2011); Cao and Lei (2012)o )

0.0450∼0.0810

Good Weather Ridge (Lin et al., 2006) Dongsha ZD3 (Cao and Lei, 2012o )

0.0018∼0.2190 0.0316

Shenhu (Yang et al., 2010; Wu et al., 2013)

0.0020∼0.0269

7.7∼87.9

Xisha C14 (Luo et al., 2013)

0.0144

14.3

Qiongdongnan (Wu et al. (2011)o ; Yang et al. (2013))

0.0078∼0.0595

∼7

Blake Ridge (Dickens, 2001)

0.0072∼0.0079

21.4∼22.8

Ulleung Basin (Kim et al., 2007o )

0.0158∼0.0393

3.17∼7.55

Black Sea (Jørgensen et al., 2001)

0.0401∼0.1022

2∼4

Gulf of Mexico (Coffin et al., 2008)

0.0204∼0.2491

Namibia (Niewöhner et al., 1998)

0.0215∼0.0615

0.0243∼0.0517

Skagerrak (Knab et al., 2008)

0.0839∼0.2263

0.0511∼0.0803

0.073∼0.106

0.5∼1.0

Chile (Treude et al., 2005)

0.0460∼0.0997

0.0254∼0.0467

0.001∼1.124

2.15∼3.65

AC C

Taixinan Site A (This study)

5.5∼6 0.00001∼0.087

8

0∼4.1

Table 3: Comparison of SO2– 4 fluxes of various sites.

42

1∼>6

4∼10

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FCH ,mod 4 (mmol m−2 yr−1 )

FSO ,mod 4 (mmol m−2 yr−1 )

FSO ,grad 4 (mmol m−2 yr−1 )

(mmol m−2 yr−1 )

Site A

6.0

Site Bp

27.4 (6+21.4)

ΣRAOM

10.3

14.6

14.7

13.2

27.6 (25.4+2.21)

28.0 (33.6-5.6)

41.5

27.0

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FCH ,ex 4 (mmol m−2 yr−1 )

Site

Table 4: Fluxes of SO4 and CH4, model estimation vs. gradient calculation.

Sample ID

Depth

Cl–

(mbsf)

(mM)

2– SO4 (mM)

Na+

K+

Mg2+

Ca2+

I–

Ba2+

δ 13 C

Alk.

(mM)

(mM)

(mM)

(mM)

(µM)

(µM)

(% V-PDB)

(mM)

CH4 (ppm)

0.3579

-6.69

2.62

20.39

2.3

0.4986

-5.53

3.64

0

558

30.3

454.5

11.3

48.3

11.7

A-2/32

0.27

569

29.5

460.0

11.5

48.1

12.0

A-3/32

0.47

561

28.5

453.3

11.9

47.6

A-4/32

0.72

566

28.4

457.8

11.2

48.6

A-5/32

1.07

561

27.9

461.8

11.2

49.2

A-6/32

1.27

567

27.6

453.3

11.6

47.7

A-7/32

1.47

562

27.4

457.0

11.4

48.6

A-8/32

1.67

568

26.9

455.4

11.4

48.3

A-9/32

1.87

566

26.6

A-10/32

2.07

565

26.3

A-11/32

2.27

567

25.8

A-12/32

2.47

569

25.4

A-13/32

2.67

569

25.2

A-14/32

2.87

568

24.9

A-15/32

3.07

569

A-16/32

3.27

563

A-17/32

3.47

577

A-18/32

3.67

568

A-19/32

3.87

566

23.4

A-20/32

4.07

564

22.7

A-21/32

4.27

564

A-22/32

4.47

576

A-23/32

4.67

575

A-24/32

4.87

582

A-25/32

5.07

A-26/32

5.27

A-27/32

5.47

A-28/32

5.67

A-29/32

5.87

A-30/32

6.07

A-32/32

3.7

0.4067

-8.77

4.28

12.4

4.1

0.3887

-8.51

4.34

11.4

5.2

0.4083

-8.81

4.84

11.4

5.8

0.5843

-10.36

5.11

11.6

6.1

0.4966

-17.35

5.41

11.9

7.3

0.4905

-12.02

5.92

M AN U

12.0

458.5

11.7

48.2

10.9

8.5

0.5319

-12.67

6.01

452.6

11.4

47.5

10.8

7.6

0.8664

-14.45

6.31

453.0

11.7

47.8

10.6

8.6

0.3690

-14.26

6.74

441.2

11.4

46.9

10.5

10.1

0.5863

-15.18

7.10

453.0

11.3

47.6

10.5

10.4

0.8892

-15.94

7.13

447.9

11.2

47.2

10.1

11.0

0.5375

-16.45

7.28

46.5

9.8

450.8

11.5

46.4

9.7

11.3

46.9

13.9

13.7

0.4690

45.5

7.8

13.1

0.6472

7.68 7.84

47.5

9.7

13.9

0.4758

-18.47

7.64

454.3

11.2

48.1

12.9

14.6

0.3909

-20.07

8.55

46.5

7.3

8.49

D

1.8883

9.6

14.5

0.3812

-19.30

8.60

49.3

9.1

17.2

0.9505

-20.26

8.39

46.4

9.0

18.0

0.5818

-21.17

9.03

21.5

11.4

43.9

8.1

17.6

0.5257

-21.54

9.04

447.7

11.3

44.1

8.1

17.7

0.5031

-22.02

9.16

21.1

450.1

11.2

43.8

7.8

17.9

0.5166

-21.57

9.26

20.6

445.3

11.2

44.5

4.5

18.1

0.5816

-21.90

9.52

20.4

445.7

11.2

44.0

7.6

18.9

0.5346

-23.03

9.53

19.8

452.5

11.0

44.6

7.4

18.3

0.5033

-24.24

9.92

22.4

462.2

11.6

22.0

457.7

11.3

567

448.4

578

21.3

578 553 572

TE

11.3

6.27

564

20.1

450.5

11.2

43.8

7.6

20.1

0.4355

-23.10

9.63

6.47

561

19.4

449.7

11.1

44.2

8.2

18.6

0.4902

-24.03

10.25

AC C

Table 5: Pore water geochemistry data for Site A.

n Only

measured values are provided. result is calculated in this study, using data of the author(s). p The values in the parentheses include diffusive flux plus advective flux. o The

43

11.76

11.77

11.62

7.75 -18.19

11.1

454.4

14.17

7.63

446.5

22.3

EP

A-31/32

23.9

SC

A-1/32

11.34

8.02

17.55

Depth

Cl–

(mbsf)

(mM)

2– SO4 (mM)

Na+

K+

Mg2+

Ca2+

I–

Ba2+

δ 13 C

Alk.

(mM)

(mM)

(mM)

(mM)

(µM)

(µM)

(% V-PDB)

(mM)

CH4 (ppm)

0.6498

0.67

3.03

7.58

0.5650

-10.01

3.25

0.4140

B-1/38

0

561

29.7

450.0

11.3

47.4

12.5

3.2

B-2/38

0.27

560

29.2

461.0

12.1

46.6

11.1

4.6

B-3/38

0.47

562

28.9

453.3

11.5

47.2

11.5

6.1

B-4/38

0.67

563

28.6

452.0

11.6

47.1

B-5/38

0.83

559

28.4

466.8

11.9

46.2

B-6/38

1.07

560

28.2

475.1

12.1

48.4

B-7/38

1.27

558

27.6

444.6

11.1

46.0

B-8/38

1.47

563

27.1

447.5

11.5

45.9

B-9/38

1.67

560

26.7

447.3

11.6

45.7

B-10/38

1.87

567

26.2

444.7

11.5

45.4

B-11/38

2.07

561

26.0

467.5

11.9

45.2

B-12/38

2.27

563

25.7

461.4

11.7

B-13/38

2.47

567

25.2

457.9

B-14/38

2.67

566

24.5

B-15/38

2.87

560

23.8

B-16/38

3.07

564

23.4

B-17/38

3.27

564

23.0

B-18/38

3.47

566

22.0

B-19/38

3.67

560

21.6

B-20/38

3.87

566

20.8

B-21/38

4.07

559

20.2

M AN U

Sample ID

RI PT

ACCEPTED MANUSCRIPT

B-22/38

4.27

562

19.4

B-23/38

4.47

561

B-24/38

4.67

563

B-25/38

4.87

560

B-26/38

5.07

B-27/38

5.27

B-28/38

5.47

B-29/38

5.67

B-30/38

5.87

B-31/38

6.07

3.57

7.1

0.3993

-8.08

3.83

11.2

6.5

0.3756

-10.00

3.79

12.2

7.9

0.3697

-11.81

3.99

11.0

9.5

0.3649

-11.88

4.34

10.9

10.6

0.3546

-15.87

4.68

11.3

11.2

0.3148

-15.29

4.89

10.2

13.2

0.3803

-15.32

5.43

4.9

14.8

0.3210

-17.26

5.72

45.4

10.6

16.3

0.9386

-21.53

6.02

11.5

45.0

10.2

18.2

0.3005

-19.53

6.37

443.1

11.2

43.7

10.1

19.0

0.3782

-20.34

6.67

441.0

11.1

44.8

9.4

21.4

0.2735

-21.84

6.89

450.1

11.1

44.7

9.4

20.9

0.3423

-22.41

7.27

436.7

11.2

43.1

9.4

20.0

0.2945

-22.39

7.35

448.4

11.3

43.9

9.2

25.2

0.3581

-23.76

7.97

453.3

11.2

44.0

9.3

23.9

0.4444

-23.83

8.11

441.8

10.9

42.6

8.5

27.2

0.3630

-24.55

8.54

445.3

10.7

43.4

8.2

27.1

0.2998

-19.12

8.93

446.9

11.2

43.5

8.7

29.0

0.6590

-25.70

9.41

18.7

448.6

10.9

42.9

8.2

31.3

0.5518

-27.64

9.67

18.0

470.3

11.5

42.3

7.6

33.9

0.5542

-28.79

9.99

17.2

438.6

10.8

41.5

6.8

33.4

0.5200

-30.09

10.44

562

16.3

447.0

10.7

41.4

6.9

34.8

0.3734

-30.43

10.89

551

15.2

446.2

11.0

39.9

6.0

38.6

0.5037

-31.08

11.73

553

14.0

438.8

10.9

38.9

5.3

38.5

0.4827

-32.82

11.71

561

13.3

472.9

11.2

40.7

5.7

37.2

0.5872

-34.14

12.62

565

12.1

470.8

11.0

40.8

4.9

38.5

0.5444

-36.03

13.07

565

11.6

TE

D

SC

-5.67

11.7

11.6

40.6

4.8

39.9

0.6080

-36.49

13.36

553

10.4

469.5

11.7

39.8

3.8

41.4

0.8039

-36.45

13.42

6.47

552

9.4

468.6

12.7

38.9

4.2

44.8

0.7549

-37.77

14.03

6.67

561

8.3

464.7

11.2

39.4

3.4

47.9

1.0003

-39.29

14.62

6.87

551

7.3

451.2

10.8

37.7

3.5

45.9

1.0909

-40.92

15.37

7.07

560

6.4

459.6

11.3

37.9

3.1

48.6

1.4747

-41.32

16.08

B-37/38

7.27

561

5.4

463.1

10.8

37.8

3.0

48.6

1.8516

-41.35

16.53

B-38/38

7.47

568

4.4

462.6

11.0

37.8

2.5

57.1

2.3123

-42.29

17.69

B-33/38 B-34/38 B-35/38

AC C

B-36/38

EP

467.5

6.27

B-32/38

Table 6: Pore water geochemistry data for Site B.

44

7.60

15.08

10.58

11.06

9.25

21.79

16.35

20.11

22.56

2120.00

1789.00