Assessing biogeochemical controls on porewater dissolved inorganic carbon cycling in the gas hydrate-bearing sediments of the Makran accretionary wedge, Northeastern Arabian Sea off Pakistan

Abstract

Quantitatively assessing the porewater dissolved inorganic carbon (DIC) cycling in methane-enriched marine sediments is crucial to understanding the contributions of different carbon sources to the global marine carbon pool. In this study, Makran accretionary wedge was divided into Zone 1 (high methane flux area) and Zone 2 (background area). Porewater geochemical compositions (Cl^–, SO₄^2–, NH₄⁺, Mg²⁺, Ca²⁺, Ba²⁺, DIC and δ¹³C-DIC) and a reaction-transport model were used to determine the DIC source and calculate the DIC flux through carbonate precipitation and releasing into overlying seawater in sediments. Zone 1 is characterized by the shallower depth of sulfate-methane transition (SMT), where most of porewater sulfate was consumed by anaerobic oxidation of methane (AOM). In contrast, a relatively low flux of methane diffusion in Zone 2 results in a deeper SMT depth and shallow sulfate is predominantly consumed by organoclastic sulfate reduction (OSR). Based on the porewater geochemical profiles and δ¹³C mass balance, the proportions of porewater DIC originating from methane were calculated as 51% in Zone 1 and nearly 0% in Zone 2. An increase of porewater DIC concentration leads to authigenic carbonate precipitation. Solid total inorganic carbon (TIC), X-ray diffractometry (XRD) and scanning electron microscopy (SEM) analysis display that carbonate content increases with depth and aragonite appears at or below the depths of SMT. Meanwhile, the flux of DIC released from sediments calculated by the reaction-transport model is 51.3 ~ 90.4 mmol/m²·yr in Zone 1, which is significantly higher than that in Zone 2 (22.4 mmol/m²·yr). This study demonstrates that AOM serves as the dominant biogeochemical process regulating the porewater DIC cycle, which has an important impact on the authigenic carbonate burial and the seawater carbonate chemistry.

1 Introduction

Methane (CH₄) seepage refers to methane-rich fluid leakage from the seabed, which is widely observed in active and passive continental margin slopes. The fluids show apparent spatial-temporal variations (Haas et al., 2010; Feng et al., 2016; Feng et al., 2020; Liu et al., 2020b; Wei et al., 2021), and often be accompanied by gas hydrate accumulations in areas of gas hydrate stability and where methane concentrations reach saturation (Milkov, 2004; Coffin et al., 2014; Mau et al., 2017; Feng et al., 2018; Zhang et al., 2021a). Under the driving force of geological structure, pressure or temperature changes, the gas hydrate can dissociate and increase the intensity of gas seeps, thereby significantly affecting benthic biogeochemical processes that control the highly spatially heterogeneous seafloor ecosystems (Suess, 2014; Liu et al., 2020a; Liao et al., 2022).

Although the CH₄ carbon estimates of hydrate-bound gas in global marine sediments are thought to be highly uncertain, the current estimate is widely believed to range from 1000 to 2500 Gt (Kvenvolden, 1988; Lerche, 2000; Milkov, 2004; Kretschmer et al., 2015). This buried organic carbon reservoir greatly exceeds many other reservoirs such as marine dissolved materials (980 Gt), terrestrial biota (830 Gt), the atmosphere (3.6 Gt) and marine biota (3 Gt) in the global carbon cycle (Kvenvolden, 1988). The dissociation of gas hydrates will result in the consequent liberation of various carbonaceous compounds such as CH₄, CO₂, and HCO₃^– (Suess et al., 1999; Phrampus and Hornbach, 2012). Previous studies have shown that more than 90% of the CH₄ is consumed through anaerobic oxidation of methane (AOM) in marine sediments (e.g., Boetius et al., 2000; Haese et al., 2003; Chen et al., 2021; Wei et al., 2021; Li et al., 2023), which effectively prevents the diffusion of dissolved CH₄ into overlying seawater or even the atmosphere. The AOM is mainly coupled to the sulfate reduction and occurs at a distinct zone which is known as the sulfate-methane transition (SMT) in the sediments (Equation 1) (Reeburgh, 1976; Borowski et al., 1996; D'Hondt et al., 2002; Snyder et al., 2007).

The upward flux of methane is theoretically nearly equal to the downward diffusive sulfate flux within the SMT due to the stoichiometric ratio of 1:1 between SO₄^2- and CH₄ in the AOM (Niewöhner et al., 1998; Knittel and Boetius, 2009; Akam et al., 2020). However, the calculated downward diffusive sulfate flux often exceeds that of the upward flux of methane (e.g., Wu et al., 2016; Akam et al., 2020; Zha et al., 2022). A compilation by Egger et al. (2018) from 740 sites covering a wide range of oceanic conditions observed the global mean net flux ratio between SO₄²⁻ and CH₄ of 1.4:1. This observed flux imbalance can be attributed to the organoclastic sulfate reduction (OSR) through the oxidation of buried organic matter (Berner, 1980; Jørgensen et al., 2019; Chen et al., 2020; Zha et al., 2022), and the reaction is represented in very simple form as Equation 2:

Both the microbially-mediated AOM and OSR can generate dissolved inorganic carbon (DIC) and increase porewater alkalinity (Equations 1, 2), thus leading to the promotion of authigenic carbonate precipitation (Equation 3) (Chatterjee et al., 2011; Smith and Coffin, 2014; Chen et al., 2023).

It should be pointed out that other alkaline earth elements such as Mg, Sr, Ba, and Fe can substitute for Ca in varying amounts during the process of carbonate precipitation (Equation 3). In addition to being removed by carbonate precipitation, a portion of the porewater DIC will migrate toward the seafloor and finally enter the overlying seawater (e.g., Smith and Coffin, 2014; Graves et al., 2015; Feng et al., 2020). Therefore, quantitative assessment of the porewater DIC cycle in marine sediments is crucial for better understanding the global authigenic carbonate burial and seawater carbonate chemistry.

The DIC concentration and its isotopic composition (δ¹³C-DIC) in porewater serve as sensitive indicators of the alteration of carbon species in marine sediments (Chen et al., 2010; Feng et al., 2020; Zha et al., 2022). Previous studies of DIC concentration and its δ¹³C-DIC have quantitatively tracked various sources, sinks, and transformation terms using the end-member isotopic compositions of various carbon pools (e.g., Aller and Blair, 2004; Formolo et al., 2004; Winde et al., 2014; Wu et al., 2016). Due to the high spatial and temporal heterogeneity of methane turnover above the gas hydrate stability zone (GHSZ), it is necessary to conduct geochemical investigations of the dissolved species found in the porewater of closely associated marine sediments. These resulting geochemical data will help constrain the cycling and fate of DIC related to diagenetic processes in marine sediments with similar continental margin environments, and thus contribute to ascertaining global carbon flux in these settings.

Methane seepage and obvious gas emissions widely occur on the continental slope of the Makran accretionary wedge off Pakistan (von Rad et al., 2000; Delisle and Berner, 2002; Fischer et al., 2012; Wei et al., 2021). Existing works have mainly focused on the use of geophysical methods to identify the distribution and reserves of methane hydrate in this area (von Rad et al., 2000; Römer et al., 2012; Wei et al., 2021; Wu et al., 2021). The conversion of methane carbon to DIC by geochemical processes in the Makran accretionary wedge was rarely reported and in particular, the quantitative assessment of the DIC cycle was not involved. In 2018, China Geological Survey carried out a China - Pakistan joint marine scientific cruise on the Makran continental margin offshore Pakistan (Figure 1A). The hydroacoustic evidence for Bottom Simulating Reflections (BSRs) indicates significant methane-hydrate burial in the study area (e.g., Wei et al., 2021; Wu et al., 2021; Zhang et al., 2021b). In this study, we collected six sediment gravity cores and analyzed the porewater geochemical composition (e.g., Cl^–, SO₄^2–, NH₄⁺, Ca²⁺, Mg²⁺, Ba²⁺, DIC and δ¹³C-DIC), total organic carbon (TOC) and carbonate phase compositions. Our aims are to determine the source of DIC, the relative contributions of different geochemical processes (AOM vs. OSR) to the porewater DIC pool, and quantitatively assess the fate of DIC in the study area using the reaction-transport model.

Figure 1

2 Study location and methodology

2.1 Study area

The study area is located in the Makran accretionary wedge of the northern Indian Ocean, which has a complex thrust structure and the lowest subduction angle worldwide (about 3°) (Kopp et al., 2000; Wu et al., 2021) (Figure 1A). Shallow gas accumulates and moves along normal faults in abundant anticlinal hinges created by the highest subducted sediments, resulting in seepage to the seafloor (Römer et al., 2012; Himmler et al., 2015; Liao et al., 2022). Seismic data revealed widespread distributions of BSRs in the slope area across the Makran continental margin (Figure 1C), indicating the widespread occurrence of gas hydrates in this area (e.g., Grando and McClay, 2007; Zhang et al., 2020; Liao et al., 2022). Various types of faults, mud diapirs, gas chimneys, and associated diapiric structures provide conduits for methane seepage in the Makran (von Rad et al., 2000; Schlüter et al., 2002; Khan et al., 2020; Wei et al., 2021; Wu et al., 2021; Liao et al., 2022). Methane seepage has been discovered ubiquitously (Römer et al., 2012), accompanied by cold seep ecosystems (Fischer et al., 2012) and authigenic carbonates (Himmler et al., 2015).

2.2 Sample collection

From December 2018 to February 2019, a systematic multi-beam sonar survey for gas emissions was conducted with the R/V “Haiyangdizhi 10” during a scientific diving expedition of the Makran accretionary wedge off Pakistan (Wu et al., 2021). Based on active flare density and methane seepage intensity, six sediment gravity cores with lengths ranging from 355 to 835 cm were recovered in two zones: Zone 1 (sites G01, G16, G19, and G34) has a high methane flux and a water depth of 500-2500 m; Zone 2 is the background area near Zone 1, with a water depth greater than 3000 m. In terms of geological structure, Zone 1 is heavily transected by imbricate faults and accretionary structures, while the forearc area is largely intact in deeper Zone 2 (Figure 1B).

After being retrieved, the core was promptly transported to the onboard laboratory for the purpose of collecting methane and extracting porewater. Sediment samples for methane analysis (3 ml) were taken using a 10 ml cut syringe and immediately transferred into 10 ml headspace vials, which contained 3 ml of 1 mol/l NaOH solution. The vials were subsequently capped with a rubber septum and sealed with a crimper, stored at low temperatures for the determination of sediment CH₄ concentration. Porewater samples were collected at 60 cm intervals using Rhizon samplers according to the procedure presented by Liu et al. (2020a). The porewater samples were then immediately chemically treated and preserved for a variety of analysis purposes. The solid phase was divided into sub-samples at 10 cm intervals and stored at −20 °C for further analysis.

About 2 ml of porewater subsamples for cations (Ca²⁺ and Mg²⁺) and anions (SO₄^2– and Cl^–) analysis was acidified with 20 μl 65% HNO₃ (suprapure, Merck) to prevent precipitation and sulfur oxidation. Samples for porewater NH₄⁺ analysis were preserved in a 2 ml PE tube and frozen. Subsamples for measurements of DIC concentration and its isotopic composition were stored in 3 ml Labco anaerobic flasks without a headspace and kept at 4 °C in the research vessel’s cold room. Once on land, these samples were immediately placed in a refrigerated case and shipped to the onshore laboratories for further analysis within 4 weeks.

2.3 Analytical methods

2.3.1 Chemical and isotopic analysis of porewater

Dissolved Ca²⁺, Mg²⁺ and Ba²⁺ concentrations were diluted at 1:30 with 18-MΩ water and measured with inductively coupled plasma atomic emission spectroscopy (ICP-AES; IRIS Advantage, Thermo Jarrell Ash). Scandium (0.5 mg/L) was added to both standards and samples for correcting signal drift during instrument operation. Porewater samples were diluted with 2% subboiled HNO₃ along with reference materials (CASS-4) from the National Research Council of Canada (NRCC). The accuracy and precision were better than ± 5%.

Porewater samples for SO₄²⁻ and Cl⁻ concentrations were diluted at 1:250 using 18-MΩ water and measured using a Dionex DX-600 ion chromatograph equipped with a 4 mm AS-9HC column. Reagents from the National Research Center for Certified Reference Materials (Beijing, China) were used as standards. Measurement of duplicate samples indicates the analytical error is less than 2%.

The samples for ammonium (NH₄⁺-N) concentration were diluted at 1:5 and determined using flow injection spectrophotometry in wavelength 543 nm, as described by Hall and Aller (1992). NH₄⁺-N standard solution was prepared using the NH₄⁺-N reference material (BWZ6768-2016). All standard solutions were appropriately diluted to generate a series of standard concentrations. Duplicate analysis after every 10 samples was performed to validate the porewater analysis procedure, and the reproducibility is better than ± 5%.

The CH₄ concentration was measured using the method of Jørgensen et al. (2001). The vial containing the sediment sample was shaken thoroughly to allow the porewater methane between the aqueous and gaseous phases to equilibrate. A 0.2-ml headspace gas sample was injected into a gas chromatograph (Agilent 6820). For porewater dissolved CH₄ concentration, the calculated CH₄ concentration per sediment volume is corrected by assuming a constant sediment porosity of 0.75. The two duplicate samples were measured with a precision of ± 3.0%.

The DIC concentration and its δ¹³C-DIC were analyzed with the 0.3 ml porewater samples according to the methods described by Liu et al. (2020a). The sealed porewater samples were acidified in precleaned extainers with 0.2 ml of 2 mol/l H₃PO₄, and the resulting CO₂ was purged from the samples. The DIC concentration and the carbon isotopic composition of the CO₂ were measured using a Thermo Finnigan MAT 253 isotope ratio mass spectrometer coupled with a Thermo Finnigan Gasbench II. The DIC concentration was measured with a precision of ± 2%. The carbon isotopic ratios are reported in the standard δ notation concerning the Vienna PeeDee Belemnite (VPDB) standard and with both precision and accuracy of better than ± 0.15% (1σ).

2.3.2 Solid sediment analysis

Sediment total organic carbon (TOC) and its stable isotopic values (δ¹³C-TOC) were analyzed in duplicate, using an elemental analyzer (Vario EL III, Germany) connected to a continuous flow system (Isoprime 100), after removal of carbonate with HCl. The average standard deviations of the measurements were ± 0.2% for TOC and ± 0.2‰ for δ¹³C-TOC. Weight percentages of total carbon (TC) were determined in duplicate (without acidification) by combustion (at 1000°C) on the elemental analyzer, and total inorganic carbon (TIC) was determined by difference: TIC = TC − TOC.

The X-ray diffraction (XRD) was performed to characterize the mineralogy of the powders of interest. Selected samples (55 cm, 415 cm, and 475 cm at site G16; 55 cm, 415 cm, and 535 cm at site G34; 55 cm, 355 cm and 775 cm at site G58) of 1.5 g were micronized under ethanol with a McCrone micronizing mill and then oven-dried at 60°C. After drying, the micronized samples were thoroughly mixed in an agate mortar and pestle to ensure homogeneity, followed by analyses using an X’pert Pro X-ray powder diffractometer (PANalytical, Netherlands) with Cu Kα radiation at 40 kV and 40 mA. The measurements were carried out in the range of diffraction angles 2θ from 4° to 80° with a scanning step of 0.033° and rotation of the sample at a speed of 15 rpm. Qualitative and quantitative analyses of minerals were performed using the software Profex (version 5.0.2) (Doebelin and Kleeberg, 2015).

Scanning electron microscopy (SEM) (Tescan Mira 3) with energy-dispersive X-ray microanalysis (EDX) (Oxford Ultim Max 40) was used to determine the morphology and elemental composition of the sites G16 and G34 samples because of their high level of carbonate content based on the XRD analysis (Supplementary Table S1). The freeze-dried subsamples were fixed onto aluminum stubs with two-way adherent tabs and allowed to dry overnight. Subsequently, the samples were sputter-coated with gold for 30 s. The SEM was operated at 2 kV with a working distance of 10 mm to facilitate optimum image collection while minimizing charging and sample damage. For EDX analysis, an accelerating voltage of 20 kV was used to generate sufficient X-ray counts.

3 Results

3.1 The depth profiles of porewater geochemistry

Figure 2 shows the depth profiles of the Cl^–, SO₄^2–, CH₄ and NH₄⁺ concentrations in the porewater of Zone 1 and Zone 2. Porewater Cl^– concentration at all sampling sites remains almost constant and close to that in seawater. The SO₄^2– concentration decreases almost linearly with depth in Zone 1, and complete depletion occurs at around the depth of 400 cmbsf (centimeters below sea floor) ~ 500 cmbsf (SMT depth). By contrast, complete depletion of SO₄^2– concentration appears at around 1800 cmbsf at site G57 and 1100 cmbsf at site G58 in Zone 2. A noticeable increase in CH₄ concentration was observed around the SMT of the sites G01, G16, G19 and G34 in Zone 1, but CH₄ concentration is close to 0 and no change with depth at sites in Zone 2. The concentration of NH₄⁺ ranges from 209.50 μM to 3243.00 μM, and an obvious negative correlation between the NH₄⁺ and SO₄^2– in depth profiles occurs at all sites.

Figure 2

Depth profiles of the porewater Ca²⁺, Mg²⁺ and Ba²⁺ are shown in Figure 3. The porewater Ca²⁺ concentrations at sites in Zone 1 and Zone 2 vary from 1.38 – 7.96 mM and 4.40 – 8.50 mM, respectively. The porewater Mg²⁺ concentrations are in the range of 37.76 – 55.91 mM and 50.64 – 56.33 mM at sites in Zone 1 and Zone 2. The concentrations of Ca²⁺ and Mg²⁺ decrease with depth in Zone 1, and their decreasing trends are more pronounced than those in Zone 2. The porewater Ba²⁺ concentrations at the sites of Zone 1 vary from 0.33 to 4.76 μM, and display a noticeable increasing trend below the depth of complete depletion of porewater sulfate. At the sites of Zone 2, however, the Ba²⁺ concentrations remain almost constant in the depth profiles, with a range of 0.35 to 0.54 μM.

Figure 3

The porewater DIC concentration increases with the depth, and the δ¹³C-DIC values are negatively correlated with DIC concentration at all sites in Zone 1 (Figure 4A). The porewater DIC concentrations at sites G01, G16, G19 and G34 vary from 11.85 – 24.29 mM, 5.47 – 18.26 mM, 4.24 – 15.45 mM and 3.79 – 18.87 mM, and the δ¹³C-DIC values range from −29.57 to −15.82 ‰, −35.04 to −15.70 ‰, −39.60 to −12.20 ‰ and −27.39 to −11.27 ‰, respectively. The most negative δ¹³C-DIC occurs at the depth of maximum sulfate depletion. For site G58 in Zone 2 (Figure 4B), δ¹³C-DIC values decrease almost linearly from −8.48‰ at the top to a minimum of −15.52‰ at the bottom. For site G57 in Zone 2 (Figure 4B), which has a much deeper SMT, the change in δ¹³C-DIC values is much more gradual since the site of AOM contributions to DIC is located deeper.

Figure 4

3.2 The depth profiles of sediment TOC, TIC and δ¹³C-TOC

Figure 5 displays the depth profiles of sediment TOC, TIC and δ¹³C-TOC. At the sites of Zone 1, the TOC shows a decreasing trend as depth increases, varying from 2.02% to 0.28%, while the TIC contents are in the range of 2.03% – 4.21% and exhibit a small increase trend with depth. In contrast, the TOC and TIC at site G58 in Zone 2 show small fluctuations in the depth profile, with ranges of 0.59% – 1.32% for TOC and 2.15% – 2.69% for TIC, respectively. At all sites, there is no clear trend of the δ¹³C-TOC variation in the depth profile, and their values range from –24.14‰ to –20.59‰.

Figure 5

3.3 Morphology and mineralogy of the carbonate phase

Based on the vertical distribution of porewater geochemical profiles (Figures 2, 3), we selected the sediment samples from three different depths of Zone 1 (sites G16 and G34) and Zone 2 (site G58) for XRD analysis, and the results are shown in Figure 6 and Supplementary Table S1. Quantitative analyses of minerals were performed using the default structure repositories in the software Profex (version 5.0.2). Overall, the carbonates at sites G16, G34 and G58 are composed of calcite, aragonite, and dolomite. The percentage of each carbonate mineral within the bulk sediment and the relative abundance of each carbonate type in total carbonates are listed in Supplementary Table S2. By comparison, the ratio of calcite/∑carbonate is the highest in all selected samples (66.48% ~ 85.18% at site G16; 51.56% ~ 81.69% at site G34; 75.59% ~ 82.68% at site G58). Aragonite is almost undetectable in the shallow sediments (55 cmbsf depth) of sites G16 and G34 (aragonite/∑carbonate = 0), and not detected at all selected depths of site G58. The ratio of aragonite/∑carbonate decreases from 20.61% at 415 cmbsf depth to 6.94% at 475 cmbsf depth of site G16, and it increases dramatically from 4.50% at 415 cmbsf depth to 35.26% at 535 cmbsf depth of the site G34. However, the difference in the ratio of dolomite/∑carbonate is very small in all selected samples (12.91% ~ 14.82% at site G16; 13.18% ~ 18.64% at site G34; 17.32% ~ 24.41% at site G58).

Figure 6

In addition to XRD analysis, the SEM images and EDS signals can also be utilized to identify the various carbonate phases. In this study, most of the carbonate concretions were observed in the deep layer of sites G16 and G34 (Figure 7; Supplementary Table S3). The mineral morphology at the 415 cmbsf depth of site G16 is analogous to that at the 535 cmbsf depth of site G34. Both samples contain aragonite and calcite, where aragonite occurs in the habit of acicular CaCO₃ with roughly strip-shaped bodies ranging from 1 to 3 μm in diameter (Figures 7B, G), and calcite appears in a massive form (Figures 7C, H). At the other depths of the sites G16 and G34, no strip-shaped aragonite but only calcite was found (Figures 7A, D–F).

Figure 7

4 Modeling results of porewater geochemical profiles

The porewater geochemical profiles of sulfate, methane, calcium, magnesium, and DIC concentration were simulated by a reaction-transport model (RTM) under steady state (e.g., Berner, 1980; Chatterjee et al., 2011; Meister et al., 2019; Xu et al., 2023). The RTM was written using MATLAB R2023a and its code is provided in the Supplementary Material.

The instantaneous reaction rate equation (Equation 4) is as follows:

where C is the concentrations, t is time, ω is the sedimentation rate, φ is the porosity of the marine sediment, Ds is the effective diffusion constant and is corrected by advection and tortuosity (τ²) as τ² = 1-2·lnφ (Boudreau, 1997). The specific values of simulation parameters are shown in Table 1, and part of the relevant data in the table refers to previous studies in this area (Luff et al., 2000; Hu et al., 2018).

The time of simulation is related to the sedimentation rate and porosity. Here we used an exponential function to describe porosity:

where the value of porosity (φ₀) of sediment at the sediment-water interface (SWI) is set to 0.8, λ (-0.01) is the attenuation coefficient and x is the depth. The burial time can be calculated as (Meister et al., 2019):

Sources and sinks of methane, sulfate, and DIC are stoichiometrically coupled to rates of organic matter (OM) degradation via the simplified reactions for sulfate reduction (SR) and methanogenesis (ME).

For OM degradation, the reactive continuum model (RCM) was applied (Boudreau and Ruddick B., 1991):

where the parameter a describes the average lifetime of OM, v defines the shape of the distribution, t is burial time calculated by Equation 6 and G represents the content of OM.

Based on the OM degradation rate, the SR and ME rates can be written as follows:

where fs is an error residual function (Chuang et al., 2019; Xu et al., 2023), as:

where C_s is the sulfate concentration, C_S^* is the threshold sulfate concentration for methanogenesis (~ 1 mM), and b (~ 0.001) is a parameter controlling the steepness of f_S. When OM is degradated by OSR, fs = 1; and when OM is degradated by methanogenesis, fs = 0.

The ratio of DIC production from sulfate consumption via AOM is 1:1, and the rate of AOM can be described as:

where C_M is methane concentration, C_S is sulfate concentration and k_AOM is the first order rate constant for AOM.

The precipitation of carbonate is the sink of calcium and DIC (Equation 3), and its rate was calculated as Equation 12 (Luff et al., 2001):

where C_calcium and C_carbonate represent the contents of calcium ions and carbonate, K^*_SP is the stoichiometric solubility constant of solid carbonate, k_Ca is the first-order kinetic rate constant for carbonate precipitation. The precipitation rate of magnesium (Mg²⁺) in porewater is correlated with the precipitation rate of Ca²⁺ (Fantle and DePaolo, 2006; Fantle and DePaolo, 2007; Zhang and DePaolo, 2020), which can be written as following:

where C_Mg and C_Ca are the porewater concentrations of Mg²⁺ and Ca²⁺, and k_Mg is the first-order kinetic rate constant for magnesium precipitation.

The absolute concentrations of [¹³DIC] and [¹²DIC] were computed by separate reaction-transport (Equation 14) for each isotope (R = R₁₂ + R₁₃). Minor carbon isotope fractionation occurs during OSR, AOM and carbonate precipitation (Meister et al., 2013; Chuang et al., 2019; Meister et al., 2019). Hence, the sources and sinks are assumed proportional to the isotopic composition of the source pool, and the rate of CH₄ can be calculated as follows:

Methanogenesis can occur via two main pathways: the autotrophic pathway using CO₂ and H₂ and the acetoclastic pathway using acetate. However, independent of the pathway, the sources of methane and DIC are in isotopic proportion with the OM by the fractionation factor α and 2-α, respectively (Meister et al., 2019).

The concentration profiles of diverse species and R_AOM from Zone 1 and Zone 2 were simulated in this study. The modeling results are almost consistent with the measured concentrations of SO₄²⁻, Ca²⁺, and Mg²⁺ in Zone 1 and Zone 2 (Supplementary Table S4; Figure 8). By contrast, DIC concentration and δ¹³C-DIC have a poor match with the simulated curve (Figure 8) and the possible explanation is that DIC samples undergo degassing during collection and storage, which will cause CaCO₃ to precipitate from the solution and lead to changes in physiochemical conditions (Miller et al., 2017).

Figure 8

Considering the uncertainty in methane sampling, the simulation of methane profiles is not only based on methane concentration profiles but also the porewater profiles of SO₄^2- and δ¹³C-DIC. As shown in Figure 8, the methane concentration at sites G01, G16, G19, and G34 in Zone 1 increases significantly at the bottom of sediments, while there is no methane occurring within the depth profiles of the sites G57 and G58 in Zone 2. In addition, the R_AOM of Zone 2 is zero and the decreasing trend of Ca²⁺ and Mg²⁺ is not as obvious as that of Zone 1. Based on the calculation of the reaction-transport model, site G19 has the highest value of R_AOM with 71.55 mmol/m²·yr. Furthermore, the rates of authigenic carbonate precipitation are simulated as 7.54 mmol/m²·yr at site G01, 8.12 mmol/m²·yr at site G16, 7.20 mmol/m²·yr at site G19, and 10.11 mmol/m²·yr at site G34, respectively. In contrast, the rates of authigenic carbonate precipitation in Zone 2 are much lower than those in Zone 1. According to the OM degradation rate and calculation by the reaction-transport model (Equation 9), the methanogenesis rates at sites G01, G16, G19 and G34 are 1.70 mmol/m²·yr, 2.30 mmol/m²·yr, 1.40 mmol/m²·yr and 0.92 mmol/m²·yr, respectively, however, the rates at sites G57 and G58 are nearly 0 (Supplementary Table S5).

5 Discussion

5.1 The relationship of the fluxes of DIC and SO₄^2- with SMT depth

The SMT, an important biogeochemical zone characterized by the simultaneous consumption of sulfate and methane (AOM, Equation 1), is commonly observed in methane-rich marine sediments at relatively shallow depths below the seafloor (1 ~ 30 mbsf) (Malinverno and Pohlman, 2011). Within or near the SMT zone, the depletion of sulfate can lead to the dissolution of barite, thereby causing an increase in the concentration of porewater Ba²⁺ (Aloisi et al., 2004; Riedinger et al., 2006; Snyder et al., 2007). In this study, we also observed an increasing trend of Ba²⁺ concentration at or below the depth of SMT, but its concentration was relatively low, which may be mainly due to the lower content of barite minerals in the sediments (Supplementary Table S1). Here, we define the depth of SO₄²⁻ becomes near depleted (1 mM) as the SMT depth, which can be derived using the least square method by fitting the SO₄^2– depth profiles of all sampling sites (Borowski et al., 1999; Arndt et al., 2013; Chuang et al., 2016; Feng et al., 2018). When the sediment core sampling does not cross the SMT, the depth of the SMT can be achieved by extrapolating the SO₄²⁻ concentration gradient to a depth where porewater SO₄²⁻ concentrations are near depleted (Smith and Coffin, 2014).

It should be noted that a shallow SMT often occurs a few centimeters below the seafloor with significant advection of upward fluids, such as seeps and vents, because the advecting fluids typically have a higher concentration of CH₄ than surrounding sediment, and advection often involves multiphase fluid flow (free gas and liquid) that may be episodic (Luff and Wallmann, 2003; Miller et al., 2017). However, the depths of SMT at our sampling sites are greater than 4 m and the porewater SO₄^2- concentrations exhibit a linear decrease in depth profiles, thereby suggesting that the impact of advection of upward fluids is not significant in this study area. Assuming under steady-state conditions, the diffusion rates of diverse species such as SO₄^2–, Ca²⁺, Mg²⁺ and DIC, can be calculated from the linear fit to the species concentration (Supplementary Table S4) gradient according to Fick’s first law (Equation 17) (e.g., Berner, 1978; Coffin et al., 2015):

where J is the diffusive flux (mmol/m² yr); φ is the porosity; Ds is the diffusion coefficient (m²/s) which is related to temperature; c is the concentration of diverse species in pore water (mM), and x is the depth (mbsf). The bottom water temperatures at the study sites are shown in Table 1, and porosity is obtained by a functional relationship with depth (Equation 5). In this study, we set that the downward flux (SO₄^2–) is negative (into the sediments) corresponding to an upward flux (positive) for DIC out of the sediments. Due to the core sampling not crossing the SMT, it is difficult to estimate the flux of DIC upward from below the SMT. Nonetheless, our calculation results of the reaction-transport model show that the methanogenesis rate is much lower than the rates of AOM and OSR, indicating the flux of DIC from below SMT is much smaller than that from above SMT (Equation 9; Supplementary Table S5). Thus, the DIC_net flux is approximated as DIC diffusion upward through the sediments above the SMT (J_DIC) and the loss of DIC by authigenic carbonate precipitation (Table 2), where the loss flux of DIC can be estimated using the porewater concentration profiles of Ca²⁺ and Mg²⁺ (Equation 18) (Smith and Coffin, 2014):

The fluxes of SO₄^2– and DIC_net in Zone 2 are significantly lower than those in Zone 1, and the larger fluxes of SO₄^2– and DIC_net correspond to the shallower SMT depth (Table 2). We combined the data from this study with the published data (Figure 9), and found that the DIC_net and SO₄^2– fluxes of all comparing sites have an exponential relationship with SMT depths (y = 273.63x^−1.17 and y = −267.89x^−1.37, Figure 9), which is very consistent with the report by Miller et al. (2017). The explanation for this correlation potentially lies in the fact that the fluxes of DIC and SO₄^2– to the SMT depth are mainly controlled by the reactivity and content of sediment OM in depth profiles (Egger et al., 2018; Xu et al., 2023). The diffusion flux of SO₄^2– is approximately equal to that of CH₄ within the SMT (Equation 1), so the depth of the SMT can provide a relatively qualitative prediction of upward CH₄ flux in the sediments.

Figure 9

5.2 Sulfate consumption pathways (OSR vs. AOM)

The depth profiles of porewater δ¹³C-DIC exhibit the obvious characteristic pattern reflecting the occurrence of AOM. Especially, the minimum values of δ¹³C-DIC (−27.40 ‰ to −39.60 ‰, Figure 4) at the depth of sulfate depletion are significantly more negative than that of the δ¹³C of sedimentary OM found in this region (~ −21.0 ‰, von Rad et al., 1996). If only the degradation of organic carbon occurs within this depth range, then the δ¹³C-DIC value should lie between that of seawater (0 ‰ PDB) and that of organic carbon (~−21.0 ‰). Thus, the isotopic signature of DIC of this study indicates the additional source of light carbon in Zone 1 (Figure 4). The most likely source is AOM, which produces δ¹³C-DIC that is isotopically lighter than that predicted from the simple mixing of the carbon derived from the OM degradation and seawater. Besides this, the decomposition of OM through bacterial sulfate reduction (OSR) can generate DIC and NH₄⁺ and enter the surrounding porewater in the methane-bearing sediments (e.g., Luo et al., 2013; Zha et al., 2022), so NH₄⁺ is an effective indicator of the extent of OM degradation because NH₄⁺ is mainly produced via anaerobic bacterial decomposition of nitrogen-bearing organic compounds (Wehrmann et al., 2011; Mazumdar et al., 2014). It should be mentioned that the distribution of porewater NH₄⁺ is influenced by many other factors. The first is the dilution of freshwater released by methane-hydrate decomposition. However, no significant fluctuations in Cl⁻ concentration were observed in the depth profiles, indicating that porewater NH₄⁺ concentration is not affected by this effect. Secondly, methanogenesis in the sediments can also increase porewater NH₄⁺ concentration (Mazumdar et al., 2014), but our results from the species concentration profiles and reaction-transport modeling confirm that the effect of deep methanogenesis is very limited (Supplementary Table S5). In addition, the temperature can affect ion diffusion and microbial-mediated OM decomposition, leading to changes in porewater NH₄⁺ concentration in depth profiles (Sivan et al., 2007; Zhang et al., 2021b), but the temperature variation is very small due to a relatively short sediment core (< 8.5 m) in our sampling sites. Thus, the influence of temperature on NH₄⁺ concentration can be ignored.

As shown in Figures 10A and B, the cross-plots of SO₄²⁻ vs. NH₄⁺ display a significant negative relationship, indicating the OSR process is accompanied by OM decomposition and NH₄⁺ release in the study area. By comparison, the slope of SO₄^2-–NH4⁺ at site G01 (k = −0.033) is larger than those at all other sites (Figure 10), and the possible reason is due to the adsorption of NH₄⁺ by clay minerals in the sediment at this site (Lukawska-Matuszewska and Kielczewska, 2016). Avinash et al. (2016) showed that the Arabian Sea sediments contain abundant illite with a large capacity for potassium-ammonium cation exchange, which originates from the fluvial or aeolian transport (wind-blown dust from the adjacent deserts). Therefore, we speculate that site G01 is more likely to receive abundant clay minerals like illite to absorb more NH₄⁺ than other sites because it is located in a shallower water depth (844 m, Tabel 2). Although the OSR process occurs, our calculated flux of NH₄⁺ (3.35 ~ 17.78 mmol/m²·yr) is much smaller than that of SO₄^2- (77.82 ~ 101.38 mmol/m²·yr), indicating that OSR does not play a dominant role in porewater sulfate reduction.

Figure 10

Distinguishing the relative contributions of OSR and AOM to sulfate consumption is crucial to accessing methane flux into the SMT and the associated biogeochemical processes occurring in marine sediments (Malinverno and Pohlman, 2011; Luo et al., 2013). In the reaction of AOM, the molecule ratio of the sulfate consumption vs. DIC generation is 1:1 on the basis of the chemical reaction formula (Equation 1), while in the OSR process, the consumption of one molecule of sulfate produces two molecules of DIC (Equation 2). Therefore, the scatter plot of the flux of SO₄²⁻ vs. DIC_net can be used to estimate the relative contribution fractions of AOM and OSR to total porewater SO₄²⁻ consumption (Smith and Coffin, 2014).

As shown in Figure 11, dots of Zone 1 (G01, G16, G19 and G34) in the methane-hydrate area fall near the AOM line, indicating the AOM reaction is the major process for sulfate reduction (Akam et al., 2020; Zhang et al., 2021b). In comparison, dots of G57 and G58 of Zone 2 in the background area are closer to the OSR line, thus proving that OSR is the dominant process for sulfate consumption. This is further supported by simulation results from the reaction-transport model, as the rates of AOM in Zone 1 are higher than those in Zone 2 (Figure 8; Supplementary Table S5).

Figure 11

5.3 Sources of porewater DIC

Porewater DIC mainly originates from AOM, OSR and the seawater that trapped during sediment accumulation (e.g., Borowski et al., 2000; Chatterjee et al., 2011; Feng et al., 2018; Meister et al., 2019; Akam et al., 2020). In some cases, the incorporation of DIC from deep sediments may affect the δ¹³C-DIC within the SMT (e.g., Chatterjee et al., 2011; Hu et al., 2022; Zha et al., 2022), but based on the modeling results of methane turnovers, the calculating depth-integrated AOM rates at sites G01, G16, G19 and G34 are about 25 to 51 times of the in situ methanogenesis rates (Supplementary Table S6). Therefore, in situ methanogenesis constitutes a minor proportion of the total DIC, whose effect on deep DIC production can thus be ignored in this study.

Here, an isotope mass balance model (Equation 19) was used to distinguish the source of porewater DIC in the closed systems (Borowski et al., 2000; Chen et al., 2010; Zhang et al., 2021b):

where X is the proportion of DIC from different sources, and δ¹³C is the carbon isotope composition. The subscripts AOM, SW and OM refer to DIC derived from the sources listed above and the value of Xsw is obtained through dividing the DIC concentration of the overlying seawater by the DIC value within the SMT (Borowski et al., 2000). In this study, the average value of δ¹³C-TOC for each core sediment is taken as the δ¹³C_OM. The δ¹³C value of CH₄ (−68.5 ‰) is followed as reported by Römer et al. (2012). The δ¹³C_DIC-added is defined as the isotopic value of the DIC being added to the initial porewater since the burial, which can be obtained from the slope of the linear regression graph of DIC vs. δ¹³C-DIC*DIC (Ussler and Paull, 2008; Chen et al., 2010; Zhang et al., 2021b).

The δ¹³C_DIC-added at sites G57 (−16.7 ‰) and G58 (−17.2 ‰) are more positive than δ¹³C_OM (Figure 12B), indicating that the porewater DIC in Zone 2 is less affected by AOM. In Zone 1, however, the relative contributions of the AOM process to total porewater DIC are 60% at site G01, 52% at site G16, 68% at site G19, and 25% at site G34 (Table 3). By comparison, the δ¹³C-DIC_added value at site G34 (–31.6 ‰) is similar to that at site S3 in Makran from Zhang et al. (2021b) (–33.1 ‰) indicating that the two adjacent sites have the same DIC source (Figure 12A; Table 3). The difference in the percentage of porewater DIC sources at different sites is likely related to the flux of the upward diffusion of CH₄ in the methane-hydrate area (Chuang et al., 2013; Wei et al., 2019; Feng et al., 2020).

Figure 12

5.4 Removal of DIC by authigenic carbonate precipitation

The formation of authigenic carbonate is an important pathway to remove porewater DIC, particularly in continental shelf regions (e.g., Sun and Turchyn, 2014; Bradbury and Turchyn, 2019; Wang et al., 2019; Ruban et al., 2020; Loyd and Smirnoff, 2022). In this study, porewater Ca²⁺ and Mg²⁺ concentrations decrease with depth, suggesting the authigenic carbonate precipitation at sites in the study area (Ritger et al., 1987; Aloisi et al., 2002; Snyder et al., 2007; Nöthen and Kasten, 2011). Numerical simulation by the reaction-transport model shows that the rate of authigenic carbonate precipitation is 10 mmol/m²·yr at site G34, slightly higher than that at sites G01 (8 mmol/m²·yr), G16 (8 mmol/m²·yr), and G19 (7 mmol/m²·yr). The rates of authigenic carbonate precipitation at the sites in Zone 1 (average 8.25 mmol/m²·yr) are much higher than the global average (~5 mmol/m²·yr, Sun and Turchyn, 2014), emphasizing the significance of the methane hydrate area in the formation and burial of authigenic carbonate in global marine sediments. Compared with Zone 1, the rate of authigenic carbonates at sites G57 and G58 in the background area (Zone 2) is much lower (average 4 mmol/m²·yr).

The XRD analysis shows the mineral species of authigenic carbonate precipitation at sites G16, G34 and G58, which are mainly composed of calcite, aragonite, and dolomite (Figure 6; Supplementary Table S1). The molar fraction of MgCO₃ in calcite can be obtained by lattice parameters of calcite (Titschack et al., 2011; Dos Santos et al., 2017; Fahad and Saeed, 2018). Based on the XRD/Rietveld analysis, we found that the molar fraction of MgCO₃ in calcite is less than 10% at all sites in our study, indicating the absence of high-Mg calcite (Supplementary Table S6). In Zone 1, authigenic carbonate phases exhibit notable variations at different depths (Supplementary Table S1; Figure 6) of sites G16 and G34. At the surface sediments (55 cmbsf depth), calcite is the main component of carbonate phases, and aragonite is almost undetectable. However, the highest aragonite content occurs around the SMT with a depth of 415 cmbsf at site G16 (5.27%) and 535 cmbsf at site G34 (12.09%) (Supplementary Table S1), resulting in the highest ratio of aragonite/∑carbonate and the lowest ratio of calcite/∑carbonate appear at the SMT depth (Supplementary Table S2; Supplementary Figure S1). Many factors, such as sulfate inhibition and the presence of extracellular polymeric substances (EPS), can affect the formation of carbonate mineral phases (e.g., Formolo et al., 2004; Himmler et al., 2015; Li et al., 2018; Wang et al., 2019). Based on the results of porewater DIC and XRD analysis, however, we speculate that the solubility of carbonate minerals may play a significant role in the aragonite precipitation. According to carbonate solubility, calcite (K’= 7.98×10^–7 at 5°C in 34.5‰ seawater) precipitation takes priority over aragonite (K’= 11.4×10^–7 at 5°C in 34.5‰ seawater) (Berner, 1976). Therefore, at surface depth with relatively low DIC concentrations, calcite precipitates first. With the increase of DIC generated by the microbial-mediated AOM occurring around SMT depth (Chafetz et al., 1991; Peckmann et al., 2001; Liu et al., 2016), aragonite starts to precipitate, leading to the highest ratio of aragonite/∑carbonate around the SMT depth compared to other depths. Nevertheless, this ratio remains lower than that of calcite/∑carbonate at all depths. By contrast, aragonite does not appear while calcite is the dominant carbonate phase in Zone 2 (site G58) (Supplementary Tables S1, S2), which is similar to the distribution of carbonate phases in surface sediments of Zone 1, as it is less affected by AOM (Liu et al., 2016; Li et al., 2018; Wei et al., 2020). Although dolomite is one of the carbonate phases in the study area, its content is relatively low (Zone 1: 2.73% ~ 4.47% at site G16; 2.47% ~ 4.52% at site G34; Zone 2: 2.52% ~ 3.10% at site G58) (Supplementary Table S1). Currently, the formation mechanism of dolomite is very complex and remains elusive (Cai et al., 2021). Microbially-mediated dolomite precipitation could possibly occur in modern reductive saline and anoxic environments (Warren, 2000; Cai et al., 2021), such as lagoons and gas hydrate areas (e.g., Casado et al., 2014; Li et al., 2018). Therefore, we infer that it is possible in the past that the accumulation of gas hydrates resulted in saline, sulfate-free porewaters which could host the small amounts of dolomite that we observed in this study.

5.5 Release of DIC to the overlying seawater

Theoretically, DIC that is not precipitated into authigenic carbonates can enter the upper sediment and eventually the overlying seawater through advection or diffusion (Chatterjee et al., 2011; Rassmann et al., 2018; Akam et al., 2020; Zhao et al., 2020). Based on extensive seafloor observations and geochemical analysis, methane hydrate systems contribute considerable proportions of the local carbon budget to the overlying seawater by emitting a large amount of DIC (Aharon et al., 1992; Suess et al., 1999; Garcia-Tigreros and Kessler, 2018; Feng et al., 2020). DIC released from marine sediments has an important influence on the carbonate equilibrium system in seawater (Chen, 2002; Krumins et al., 2013; Akam et al., 2020; Feng et al., 2020). Previous studies have shown that global DIC flux released from sediments to overlying seawater averages 6.5 Tmol/yr (3.2 ~ 9.2 Tmol/yr), of which 98% is released from continental margin sediments with the depth of SMT less than 13 m, equivalent to approximately 20% of the global river DIC (~ 33 Tmol/yr) flowing into the ocean (Meybeck, 1993; Amiotte Suchet et al., 2003; Treude et al., 2005; Aufdenkampe et al., 2011; Egger et al., 2018; Akam et al., 2020).

In this study, we use the reaction-transport model to quantitatively assess the DIC fluxes released from the sediments into the overlying seawater of the Makran continental margin (Table 4). In this calculation, DIC_OSR, DIC_AOM and DIC_carbonate are simulated by the rate of OM decomposition, the rate of AOM within SMT, and the depth profile of Ca²⁺ concentration, respectively. As described previously, the situ methanogenesis via shallow OM decomposition is very low (Supplementary Table S5), so the contribution of DIC generated through this process is not considered. The flux of DIC_burial is constrained by the sedimentation rate and DIC concentration (Zha et al., 2022), thus it can be estimated by the reaction-transport model (Table 4). Furthermore, the DIC released from sediment to overlying seawater can be defined as Equation 20 (Wallmann et al., 2008; Dickens and Snyder, 2009; Chatterjee et al., 2011; Solomon et al., 2014; Akam et al., 2020):

Based on the Equation 20, we find that DIC_out differs greatly between the site G19 (90.4 mmol/m²·yr) and G34 (51.3 mmol/m²·yr), although both sites are located in the methane-hydrate area (Figure 1). In addition, we compared the DIC_out values in Zone 1 (68 ± 14 mmol/m²·yr) and Zone 2 (22.4 mmol/m²·yr), and this difference is likely caused by the different DIC cycling patterns between the two Zones (Table 4; Figure 13). As shown in Table 4 and Figure 13, the DIC fluxes produced by AOM in Zone 1 is 55 ± 11 mmol/m²·yr, which is significantly higher than that in Zone 2 (DIC_AOM = 0). Although the rate of authigenic carbonate formation is also relatively high in Zone 1, the process of carbonate precipitation is insufficient to prevent the release of large amounts of DIC from the sediment into the overlying seawater. By comparison, the fluxes of DIC released from sediments range from 51.3 to 90.4 mmol/m²·yr in Zone 1, which is higher than those of other methane-hydrate areas, such as the Gulf of Mexico (9 ~ 33 mmol/m²·yr, Smith and Coffin, 2014), Dongsha area of the southern South China Sea (13.1 ~ 26.1 mmol/m²·yr, Chen et al., 2017), Shenhu area of the southern South China Sea (10.1 ~ 31.7 mmol/m²·yr, Wu et al., 2013), and Beikang Basin (32.3 ~ 50.1 mmol/m^2·yr, Feng et al., 2018).

Figure 13

This upward DIC flux across the seafloor accounts for an average of 88% of the total DIC in Zone 1. Importantly, the process of AOM contributed a large proportion of DIC_out at sites G01, G16, G19, and G34 in Zone 1 (Table 4) due to the upward flux of CH₄. High rates of AOM can lead to an increase in DIC_out flux and total alkalinity (TA) in the overlying seawater (Akam et al., 2020; Zha et al., 2022), which is beneficial to reduce ocean acidification and increases the CO₂ absorption capacity of seawater (Chen, 2002; Krumins et al., 2013; Akam et al., 2020; Feng et al., 2020). In future, more related work on carbon cycling in methane-bearing sediments should be conducted to quantitatively assess the DIC and TA released from global marine sediments to the overlying seawater and their impact on carbonate equilibrium in marine systems.

6 Conclusion

Based on the active density of flares at the Makran continental margin, we divided the study area into Zone 1 (an active accretionary environment with high methane flux) and Zone 2 (a farther offshore background area). The differences in geological structures and sampling locations between the two zones have a potential impact on the porewater geochemistry and DIC flux. Our results show that Zone 1 has the larger CH₄ flux with the shallower SMT depth, and a noticeable increase in porewater Ba²⁺ concentration was observed within or near the SMT zone. The flux of SO₄²⁻ vs. DIC_net was used to determine the pathway of sulfate consumption at each site, revealing that almost all porewater SO₄²⁻ at sites G57 and G58 in Zone 2 is consumed by OSR, the AOM reaction, however, is the main process of sulfate consumption affected by methane diffusion at sites in Zone 1. The relative contributions of AOM to the porewater DIC calculated by the isotope mass balance model are 60% (G01) and 52% (G16), 68% (G19), and 25% (G34) in Zone 1, respectively. The TIC, XRD and SEM analyses show that the total carbonate content increases with depth, and the carbonate phases are composed of calcite, aragonite, and dolomite. By comparison, calcite is the main carbonate phase while the dolomite content is relatively low. Due to high concentration of DIC induced by AOM, aragonite appears at or below the depth of SMT. The reaction-transport model results and calculation reveal that the flux of DIC released from sediments is 51.3 to 90.4 mmol/m²·yr in Zone 1, which is significantly higher than that in Zone 2 (22.4 mmol/m²·yr). Overall, this study highlights that the methane-related carbon pools are highly heterogeneous in the methane-hydrate area, and the contribution of AOM is essential to the authigenic carbonate burial and the bottom seawater chemistry.

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