Deep Subseafloor Biogeochemical Processes and Microbial Populations Potentially Associated with the 2011 Tohoku-oki Earthquake at the Japan Trench Accretionary Wedge (IODP Expedition 343)

Abstract

Post-mega-earthquake geochemical and microbiological properties in subseafloor sediments of the Japan Trench accretionary wedge were investigated using core samples from Hole C0019E, which was drilled down to 851‍ ‍m below seafloor (mbsf) at a water depth of 6,890 m. Methane was abundant throughout accretionary prism sediments; however, its concentration decreased close to the plate boundary decollement. Methane isotope systematics indicated a biogenic origin. The content of mole­cular hydrogen (H₂) was low throughout core samples, but markedly increased at specific depths that were close to potential faults predicted by logging-while-drilling ana­lyses. Based on isotopic systematics, H₂ appeared to have been abundantly produced via a low-temperature interaction between pore water and the fresh surface of crushed rock induced by earthquakes. Subseafloor microbial cell density remained constant at approximately 10⁵‍ ‍cells‍ ‍mL^–1. Amplicon sequences revealed that predominant members at the phylum level were common throughout the units tested, which also included members frequently found in anoxic subseafloor sediments. Metabolic potential assays using radioactive isotopes as tracers revealed homoacetogenic activity in H₂-enriched core samples collected near the fault. Furthermore, homoacetogenic bacteria, including Acetobacterium carbinolicum, were isolated from similar samples. Therefore, post-earthquake subseafloor microbial communities in the Japan Trench accretionary prism appear to be episodically dominated by homoacetogenic populations and potentially function due to the earthquake-induced low-temperature generation of H₂. These post-earthquake microbial communities may eventually return to the steady-state communities dominated by oligotrophic heterotrophs and hydrogenotrophic and methylotrophic methanogens that are dependent on refractory organic matter in the sediment.

The subseafloor biosphere is characterized by a highly abundant biomass and biodiversity (Morono et al., 2011; Kallmeyer et al., 2012; Hoshino et al., 2020). Scientific ocean drillings have provided insights into the phylogenetic diversity and function of a number of subseafloor environments, such as ultradeep (2.5‍ ‍km below the seafloor) sediments of the Japan Trench forearc basin (Inagaki et al., 2015), oligotrophic pelagic sediments of the South Pacific Gyre (D’Hondt et al., 2015), high-temperature sediments of the Nankai Trough (Heuer et al., 2020; Beulig et al., 2022), deeply-buried ridge flank basalt (Cowen et al., 2003; Lever et al., 2013), highly alkaline serpentinite bodies (Mottl et al., 2003; Kawagucci et al., 2018; Früh-Green et al., 2018), and deep-sea hydrothermal systems and volcanos (Yanagawa et al., 2016; de Ronde et al., 2019; Teske, 2020). These findings revealed that the development of a subseafloor microbial community is closely associated with the energy and nutrient states for growth and survival under the present physical and chemical conditions of these habitats and is affected by genetic and phenotypic adaptation histories to long-term geological and geochemical processes. To clarify a greater spectrum of the global subseafloor microbial ecosystem, further investigations are still required into previously unexplored subseafloor environments, which have different geological and geochemical processes and histories from those examined to date.

The Japan Trench Fast Drilling Project (JFAST) performed Integrated Ocean Drilling Program (IODP) Expedition 343 one year after the 2011 Tohoku-oki earthquake to investigate the subseafloor processes that occurred at the destructive earthquake and during its aftershocks (Chester et al., 2013a, 2013b; Fulton et al., 2013; Ujiie et al., 2013; Rabinowitz et al., 2015; Kodaira et al., 2021). The project drilled the hadal seafloor at a water depth of 6,890 m, reached 851‍ ‍m below seafloor (mbsf), and recovered a core from Hole C0019E (Core C0019E) covering the overlying Japan Trench accretionary prism and plate boundary decollement (Chester et al., 2013a). Drilling and post-drilling operations successfully detected the locations and structures of potential faults associated with a coseismic slip at the hole and a thermal anomaly due to frictional heating (Fulton et al., 2013). Furthermore, the pore-water chemistry of core samples indicated earthquake-induced environmental changes and subsequent microbial responses associated with the mega-earthquake and historical earthquake swarms. Onboard measurements of pore-water mole­cular hydrogen (H₂) revealed sharp peaks in a vertical profile (Chester et al., 2013a). Coseismic H₂ generation has been proposed by multiple studies that performed fault zone soil observations (Kita et al., 1980; Wakita et al., 1980), subland fault drilling (Wiersberg and Erzinger, 2008), rock crushing experiments (Kita et al., 1982), and fault sliding experiments (Hirose et al., 2011). The coseismic upwelling of deep subseafloor fluids (e.g. Kawagucci et al., 2012; Sano et al., 2014) is another example of earthquake-induced environmental changes with fluid migration. One of the scientific questions for the JFAST expedition was whether an abundant supply of coseismic H₂ at the mega-earthquake and even historical earthquake swarms affected the subseafloor microbial community compositions and functions inhabiting the accretionary prism and plate boundary decollement, particularly hydrogenotrophic chemolithoautotrophic populations, such as homoacetogens and methanogens (e.g. Wu and Lai, 2011).

The present study was primarily conducted to answer this question. We herein describe pore-water and sediment chemistries as well as the composition and functions of the microbial community in Core C0019E. In addition to onboard chemical ana­lyses performed according to the standard IODP protocol and H₂ measurements (Chester et al., 2013a), detailed onshore investigations, including stable isotope ana­lyses and culture-dependent and -independent microbiological characterizations, were conducted. To address the possible origins of the anomalously abundant H₂ enrichment at specific depths in the accretionary prism, we also performed an experimental simulation for isotope fractionation associated with H₂ generation through a low-temperature fluid-rock interaction. These comprehensive ana­lyses imply episodic and steady-state subseafloor geochemical and microbiological processes in the seismically active accretionary wedge of the Japan Trench.

Materials and Methods

Sampling

IODP Expedition 343 of D/V Chikyu was conducted between April and May 2012 (Fig. S1). Sampling procedures and preliminary results obtained during the expedition, including geological characteristics, physical properties, and chemical ana­lyses of the core, have already been reported (Chester et al., 2013a). A summary of this study is described below.

Twenty-one core sections (C0019E-1R to -21R) were recovered during the expedition. The core was classified into seven lithological units based on color, composition, grain size, and minor lithology. Unit 1 was shallower (176.5–185.2 mbsf; 1R) than the other units (648.0–836.8 mbsf; 2R-21R). Unit 3 (688.5–820.1 mbsf; 4R-16R), which was the longest, had a more terrigenous nature than Units 1 and 2 and contained abundant amounts of pyrite-like grains. Unit 4 (821.5–822.5 mbsf; 17R) consisted of sheared clay, recognized as decollement. Unit 5 (824.0–832.9 mbsf; 18R-20R) represented the underthrust incoming plate, while Units 6 and 7 (832.9–836.8 mbsf; 20R-21R) represented incoming plate sediments. In Unit 3, logging-while-drilling ana­lyses identified a series of low resistivities between 688–701 and 720 mbsf, suggesting the presence of faults. Structural geology also identified numerous fractures, including some potential faults, and the most probable major fault zones were identified at 720 and 820 mbsf (Chester et al., 2013a).

Sampling and analytical procedures on the major geochemical parameters of cations (K⁺, Ca²⁺, Mg²⁺, Mn²⁺, and NH₄⁺) and anions (Cl^–, SO₄^2–) in pore water and the composition of gas (CH₄ and H₂) as well as the solid-phase properties of Total Organic Carbon (TOC), Total Nitrogen (TN), and Total Sulfur (TS) were described in the expedition proceedings (Chester et al., 2013a). The present study used the values reported.

Pore-water was extracted from whole-round cores through a 0.2-μm disposable polytetrafluoroethylene filter by the squeezer assembly. Since most of the recovered core samples were fragmented during the coring operation and exposed to seawater and/or seawater-based drilling fluid during recovery, the pore-water sampling process, such as peeling of the outer layer, may have been insufficient to completely remove seawater contamination from the pristine inner parts of core samples. Sediment squeeze cakes were sealed in plastic bags for shore-based ana­lyses. Regarding gas ana­lyses, approximately 1‍ ‍mL of sediment or deposit was collected with a cut-off plastic syringe or corkscrew and extruded into a 20-mL glass vial containing 3‍ ‍mL of Milli-Q water and a small amount of HgCl₂ to prevent microbial activity. Onboard H₂ and hydrocarbon ana­lyses were conducted after heating the vials at 70°C for 30‍ ‍min followed by measurements with gas chromatographs (GL Science GC4000 and Agilent 6890N) equipped with a helium ionization detector (HID) and flame ionization detector (FID), respectively. Vials were then stored in a freezer until onshore isotope ana­lyses. The number of (sub)samples for pore-water (n=12) was lower than that for gas species (n=52) due to insufficient core recovery.

Subsamples for microbiological ana­lyses were taken from the inner part of a whole round core with a sterilized spatula. Microscopic observations of cell density, a cultivation test, and radioactive isotope tracer incubation experiments were conducted on 12 core samples (1, 4–8, 12–15, and 19–20R) and 16S rRNA gene amplicon sequencing on 9 (3, 6, 9–10, and 12–16). In microscopic observations, approximately 1‍ ‍g of fragments from each sample was placed into a plastic tube and fixed with 3‍ ‍mL of filtered phosphate-buffered saline (PBS; pH 7.2; filtered through a pore size of 0.22‍ ‍μm) containing 4% (w/v) paraformaldehyde at 4°C for 3 h. Fixed samples were stored at –80°C. To extract DNA, ~30‍ ‍g of the fragments from each core was placed into a plastic tube and stored at –80°C until onshore processing. In cultivation ana­lyses, approximately 75‍ ‍mL of core fragments was placed into a 100-mL autoclaved glass bottle and sealed with butyl rubber stoppers in an anaerobic glove chamber. To maintain samples under strict anaerobic conditions, the headspace of the bottles was pressurized with 200 kPa-N₂, 0.5‍ ‍mL of 5% (w/v) neutralized Na₂S solution was added, and bottles were stored at 4°C until onshore laboratory treatments.

Geochemical and microbiological analyses

Geochemical analyses

The carbon and hydrogen isotopic compositions of CH₄, H₂, and CO₂ in the headspace of vial subsamples for gas measurements were assessed using isotope ratio mass spectrometers (MAT253 and DELTA^XPADVANTAGE; ThermoFisher) at JAMSTEC as previously described (Kawagucci et al., 2010; Okumura et al., 2016). δD_H2 was quantified when subsamples contained a sufficient amount of H₂ (n=7 out of 52). The concentrations of gas species presented are the values obtained from onboard ana­lyses, whereas the mixing ratios of H₂ in the headspace gas measured in MAT253 ana­lyses were used to evaluate the magnitude of air-derived H₂ contamination by a Keeling plot (Keeling, 1958). The hydrogen and oxygen isotope ratios of interstitial H₂O were measured by cavity-ring-down spectroscopy (L2130-i; Picarro) at Northern Arizona University. Isotope ratios are presented by conventional δ‍ ‍notation in the permil scale. Analytical uncertainties (1-sigma) were within 0.3‰ (δ¹³C_CH4), 5‰ (δD_CH4), 10‰ (δD_H2), 0.5‰ (δ¹³C_CO2), 0.3‰ (δD_H2O), and 0.2‰ (δ¹⁸O_H2O). Isotope fractionation between two molecules, A and B, is expressed by α as follows:

$${{\mathrm{\alpha }}^{\mathrm{X}}}_{\mathrm{A}-\mathrm{B}}=\left(1000+{{\mathrm{\delta }}^{\mathrm{X}}}_{\mathrm{A}}\right)/\left(1000+{{\mathrm{\delta }}^{\mathrm{X}}}_{\mathrm{B}}\right)\text{,}$$

where X represents ¹³C or D.

The sulfur isotopic compositions of sulfate and sulfide in representative samples were assessed by a combination of sulfur conversion to SF₆ and a ThermoFisher MAT253 mass spectrometer with a dual inlet system at the Tokyo Institute of Technology. Sulfate dissolved in pore-water was initially precipitated as BaSO₄ by the addition of a 10% BaCl₂ solution followed by rinsing with acetone to remove native sulfur. Dried BaSO₄ was converted to Ag₂S using the Kiba reduction method (Sasaki et al., 1979). Sulfide-bearing squeezed cakes of sediments were dried and powdered using an agate mill. Powdered samples (2.0–5.5 g) were ultrasonically washed and soaked in 10% NaCl solution for 24 h. Samples were then rinsed with distilled water and centrifuged to remove soluble sulfate. The residue was washed and soaked with acetone for 24‍ ‍h to dissolve native sulfur and then rinsed with distilled water and centrifuged. The residue was dried at room temperature for >24‍ ‍h and sulfide was extracted using a modified method from Hsieh and Shieh (1997). The dried residue and alkaline Zn trap were placed in a 500-mL glass bottle. The bottle was purged with N₂ and the sample was reacted with 20‍ ‍mL of 5 M HCl followed by 20‍ ‍mL of chromium (II) solution at room temperature for >48‍ ‍h (Ueno et al., 2008). Chromium (II)-reducible sulfur (CRS) was reduced to H₂S and precipitated as ZnS in a 20-mL alkaline Zn trap. In this extraction procedure, acid volatile sulfur was also reduced to H₂S and precipitated as ZnS; however, our preliminary ana­lyses showed that negligible acid volatile sulfur was present in the samples (data not shown). ZnS was converted to Ag₂S by a reaction with 0.1 M AgNO₃ solution, cleaned via repeated centrifugation with distilled water, and dried at 70°C for >12 h. Ag₂S was reacted with excess F₂ at 300°C in a nickel reaction tube overnight to produce SF₆, which was purified using a cryogenic technique and gas chromatography. The multiple sulfur isotopic composition is presented using delta notation:

$${\mathrm{\delta }}^{\mathrm{x}}\mathrm{S}=\left\{{\left({}^{\mathrm{x}}\mathrm{S}/{}^{32}\mathrm{S}\right)}_{\mathrm{sample}}\right\}/\left\{{\left({}^{\mathrm{x}}\mathrm{S}/{}^{32}\mathrm{S}\right)}_{\mathrm{reference}}-1\right\}\times 1000(\mathrm{x}=33\mathrm{ }\mathrm{o}\mathrm{r}\mathrm{ }34)\text{,}$$

where (^xS/³²S)_sample and (^xS/³²S)_reference are the isotope ratios of the sample and reference material, respectively.

$$\mathrm{\Delta }{}^{33}\mathrm{S}={\mathrm{\delta }}^{33}\mathrm{S}-\left({\left(1+{\mathrm{\delta }}^{34}\mathrm{S}/1000\right)}^{0.515}-1\right)\times 1000\text{.}$$

The analytical reproducibilities of the δ³⁴S and Δ³³S values, as assessed by replicate ana­lyses of the international reference material IAEA-S1, were ±0.3‰ and ±0.01‰(1σ), respectively. The sulfur isotope effect during sulfate reduction is described as follows:

$${}^{34}\mathrm{\epsilon }=1000\bullet \left[1-{\left({}^{34}\mathrm{S}/{}^{32}\mathrm{S}\right)}_{\mathrm{sulfide}}/{\left({}^{34}\mathrm{S}/{}^{32}\mathrm{S}\right)}_{\mathrm{sulfate}}\right]$$

$${}^{33}\mathrm{\lambda }=\ln \left[{\left({}^{33}\mathrm{S}/{}^{32}\mathrm{S}\right)}_{\mathrm{sulfide}}/{\left({}^{33}\mathrm{S}/{}^{32}\mathrm{S}\right)}_{\mathrm{sulfate}}\right]/\ln \left[{\left({}^{34}\mathrm{S}/{}^{32}\mathrm{S}\right)}_{\mathrm{sulfide}}/{\left({}^{34}\mathrm{S}/{}^{32}\mathrm{S}\right)}_{\mathrm{sulfate}}\right]\text{,}$$

where (^xS/³²S)_sulfide and (^xS/³²S)_sulfate (x=33, 34) are the sulfur isotope ratios of sulfide and sulfate, respectively.

In textural observations of sulfide minerals near the fault zones in Unit 3, subsamples were collected from 7R02 and 13R02, which had localized high intensity parts that were observed by onboard X-ray and CT scans (Chester et al., 2013a). Collected samples were embedded in LR White resin after dehydration with an ethanol series; 70, 80, 90, and 99%. Embedded samples were polished, coated with carbon using a vacuum evaporation system (JEE-420; JOEL), and observed by a field emission Scanning Electron Microscope (SEM) (JSM-6500F; JEOL) at 15kV in Kochi University, with Energy Dispersive X-ray spectrometry (EDS) for the elemental distribution mapping.

Microbiological ana­lyses

To investigate cell density in core samples, fixed samples were washed with 1× PBS, resuspended in 1× PBS-ethanol (1:1) solution, and stored at –20°C. Just before the experiment, each sample (100‍ ‍μL) was diluted with 1× PBS (900‍ ‍μL), and the diluted sample was sonicated for 20‍ ‍s using a UH-50 ultrasonic homogenizer (SMT). Approximately 5‍ ‍mL of filtered (0.22‍ ‍μm) 1× PBS was placed into the filter tower prior to the addition of the diluted sample to ensure an even distribution of cells on the filter. The membrane was stained with SYGR green I (1/40‍ ‍[v/v] SYBR green I in TE) at room temperature for 10‍ ‍min in the dark. The membrane was rinsed with 5‍ ‍mL of TE buffer to remove excess dye. Filters were examined under epifluorescence using the phase-contrast microscope BX51 (Olympus). The average total cell count was obtained from more than 100 microscopic fields.

DNA extraction from subsamples was performed using the PowerMAX Soil DNA isolation kit (MoBio Laboratories). Samples were incubated at 65°C for 5‍ ‍min before mechanical shaking for 10‍ ‍min with a ShakeMaster (BioMedical Science), whereas subsequent steps were performed according to the manufacturer’s protocol. Extracted DNA was stored at –80°C. 16S rRNA gene fragments were amplified using the universal primer set Uni530F-U907R (Nunoura et al., 2012), and amplified 16S rRNA gene fragments were purified by agarose gel electrophoresis and the Min Elute PCR purification kit (GIAGEN). Sequence libraries were constructed using the Ion Xpress^TM Plus Fragment Library Kit (Life Technologies), and amplicon sequencing was performed on an Ion Torrent PGM sequencer (Life Technologies) equipped with an Ion 314 chip using 400-base read length chemistry by the Ion PGM™ Template OT2 400 Kit and Ion PGM^TM Sequencing 400 Kit (Life Technologies). Single-end reads from amplicon libraries were trimmed and filtered using PRINSEQ v0.20.4 with the following parameters: -trim_qual_left 20, -trim_qual_right 20, and -min_len 100. (Schmieder and Edwards, 2011). PCR primers were removed from the processed sequence using Cutadapt v1.10 (Martin, 2011). Cleaned sequences were analyzed using the QIIME2 v2019.4.0 pipeline (Bolyen et al., 2019). Between 77,105 and 146,575 high-quality SSU rRNA gene sequences were obtained, and 77,000 randomly subsampled sequences were used in further ana­lyses. Operational taxonomic units (OTUs) were constructed using a 97% similarity threshold. The taxonomic assignment of OTUs was conducted using the QIIME2 plugin feature-classifier classify-sklearn against the SILVA 138 database (Quast et al., 2013). OTUs presumed to be laboratory contaminants and members of human biomes were excluded from further ana­lyses as previously reported (Nunoura et al., 2016), which included members affiliated with the order Lactobacillales, the families Corynebacteriaceae, Micrococcaceae, Propionibacteriace, Enterobacteriaceae, Pseudomonadaceae, and Xanthomonadaceae, and the genera Leucobacter, Streptomyces, Staphylococcus, Brevundimonas, Bradyrhizobium, Methylobacterium, Pseudochrobactrum, Paracoccus, Sphingomonas, Aquabacterium, Burkholderia, Cupriavidus, Delftia, Diaphorobacter, Paucibacter, Providencia, Acinetobacter, and Pseudomonas. Several degenerate oligonucleotide primer sets, ME1/ME2 (Hales et al., 1996), mcrIRD, ANME-1 mcrI (Lever and Teske, 2015), and ME3MF/ME2r’ (Hales et al., 1996; Nunoura et al., 2008) were used to amplify the genes encoding methyl coenzyme M reductase (mcrA), thereby expanding phylogenetic coverage with the amplification conditions described in each study. To elucidate the taxonomic characteristics of the isolated pure cultures, DNA extraction, PCR amplification, and sequencing procedures were performed as previously described (Sakai et al., 2019). Amplicon sequence data were deposited into the DNA Data Bank of Japan nucleotide sequence database (DDBJ) Sequence Read Archive under DRA013695. All data were registered under BioProject ID PRJDB13255.

The population density of cultivated microorganisms represented by various physiological and metabolic characteristics was estimated by a series of quantitative cultivation tests (serial dilution cultivation method). To survey members capable of homoacetogenesis and methylotrophic/hydrogenotrophic methanogenesis, core fragments were cultivated at 20°C in MMJ medium (Takai et al., 2002) containing CH₃NH₂ (5‍ ‍mM) and acetate (5‍ ‍mM) under an atmosphere of H₂/CO₂ (80:20‍ ‍[v/v]). Approximately 0.5‍ ‍g of each core sample was resuspended in 5‍ ‍mL of sterile MJ synthetic seawater (Takai et al., 1999) containing 0.05% of Na₂S, and this was used as an inoculum. Five hundred microliters of slurry was inoculated into 15-mL tubes containing 5‍ ‍mL of MMJ medium, and 500‍ ‍μL of the medium was then serially transferred to another tube for the next dilution step. The positive dilution tube was subjected to dilution-to-extinction to obtain a pure culture. GenBank/EMBL/DDBJ accession numbers for the 16S rRNA gene sequences of Acetobacterium sp. strain JF_A and Methanolobus sp. strain JF_M are LC721303 and LC721304, respectively.

The potential microbial metabolic activities of hydrogenotrophic/methylotrophic/aceticlastic methanogenesis and homoacetogenesis were examined by the cultivation of core fragments with each of the radioactive isotope tracers (Tasumi et al., 2015; Kawagucci et al., 2018). Approximately 1‍ ‍g of core fragments was placed into a 30-mL sterilized glass vial followed by the addition of ~10‍ ‍mL artificial seawater medium (0.5 M NaCl, 10‍ ‍mM NaHCO₃, 0.001% resazurin, and 0.05% Na₂S) for slurry. The headspace of sealed slurry was purged by N₂ (or CH₄ for the methanotrophy test), and hydrogen gas was added only for the samples with ¹⁴C-bicarbonate, resulting in a final concentration of 1%. The ¹⁴C-labeled reagents of the possible metabolic substrate (NaH¹⁴CO₃, ¹⁴CH₄, ¹⁴CH₃¹⁴COOH, and ¹⁴CH₃NH₂) were individually added to each vial in order to set the final radioactivity level to 0.5 MBq per bottle. After the incubation of slurry at 20°C for one month, the ¹⁴C contents of CH₄, CO₂, and acetate in the vials were assessed by gas chromatography (Shimadzu GC-2014; Shimadzu) equipped with the radioisotope detector Raga Star (Raytest) and liquid chromatography (Agilent 1260 Infinity LC system; Agilent) coupled to an online radio flow detector (Ramona Star; Raytest).

Experiments and calculations

Low-temperature experiments

To simulate the isotope effect associated with H₂ generation through a water-rock reaction at a low temperature, we conducted a reaction between granular zerovalent iron and water, generally known as anaerobic corrosion, hereafter called the Fe-H₂O reaction. Although H₂ generation in nature involves Fe(II) in minerals, not Fe(0), the amount of H₂ experimentally generated through the Fe(II)-H₂O reaction at a low temperature was shown to be limited (Mayhew et al., 2013), which practically prohibited assessments of the isotope fractionation factor. Therefore, the Fe-H₂O reaction was adopted in the present study. Briefly, approximately 1‍ ‍g of autoclaved reagent-grade granular iron was placed into a 30-mL glass vial containing 5‍ ‍mL of pure water, δD_H2O of which was manipulated by the addition of D₂O. After purging the vial with pure He, the vial was sealed and stored at 4, 25, or 55°C. Generated H₂ reached more than 100 ppmv of the headspace of the vial within two weeks, and its δD_H2 value was measured. In this experiment, possible contamination by air, which contains 0.5 ppmv H₂ with a δD_H2 value of +150‰ (Gerst and Quay, 2000), was negligible against experimentally generated H₂.

Thermodynamic calculation

The dependence of the activity of dissolved species on the Gibbs free energy of possible metabolic reactions was estimated based on conventional methods using the B-dot activity model (McCollom and Shock, 1997; Amend et al., 2011; Shibuya et al., 2016). Three chemoautotrophic metabolic reactions involving H₂ and CH₄ were considered as follows:

4H₂+CO₂=CH₄+2H₂O

4H₂+2CO₂=CH₃COOH+2H₂O

CH₄+CO₂=CH₃COOH

The values of Gibbs free energy were computed according to the equation:

$${∆G}_r={∆G}_r^{°}+{RT\ln Q}_r\text{,}$$

where ${∆G}_r^{°}$ represents the standard Gibbs energy of the reaction, ΔG_r free energy, T the temperature in Kelvin, R the universal gas constant, and Q_r the activity product of the species involved in the reaction. The values of ${∆G}_r^{°}$ for the redox reactions were calculated with SUPCRT92 (Johnson et al., 1992). All calculations were conducted assuming the conditions of 25°C and 500 bar. Geochemical parameters given as constant values through the calculations are pH (7.9), ΣCO₂ (50‍ ‍mM), and [CH₄] (5‍ ‍mM) according to the observed values at Unit 3. The ΣCH₃COOH value of 1‍ ‍mM is given for homoacetogenesis to simulate the metabolic potential under acetate-rich conditions. ΣCH₃COOH values of 1 and 0.001‍ ‍mM are given for acetoclastic methanogenesis to simulate the transition of metabolic potentials along with acetate consumption.

Results

Pore-water chemistry

Geochemical data are presented in Supplementary Table S1. Fig. 1 shows the vertical profiles of the representative major ion Cl^–, gas concentrations (CH₄ and H₂), and carbon isotope ratios of CH₄ and CO₂ of C0019E. The average Cl concentration in samples (559±7.5‍ ‍mM) was similar to that in seawater (560‍ ‍mM), which demonstrated that pure-water contamination was negligible during the sampling and pore-water extraction processes. Sulfate concentrations measured from Unit 1–3 and Unit 5 samples ranged between 3–11 and 16–22‍ ‍mM, respectively. The significant amount of sulfate detected in deep subseafloor sediments appeared to be irregular because microbial sulfate reduction coupled with organic matter decomposition completely consumes sulfate in sedimentary environments through geological time. The significant amount of sulfate at these depths may have been due to contamination by seawater sulfate during sample recovery and processing (see Materials and Methods).

Fig. 1.

Subseafloor vertical profiles of representative parameters. The vertical profiles of (a, b, and c) the concentrations of Cl^–, H₂, and CH₄ and (d and e) the carbon isotope ratios of CH₄ and CO₂ are shown. The corresponding depths of core sections are shown by black horizontal bars with section names on the right side. Background color shades represent lithological units, while U4+ includes Units 4–7.

Sulfate isotope ratios were used as tracers of the origin and behavior of sulfur-bearing compounds. Only the two core samples (19R-1 and 20R-1) collected in Unit 5 below decollement were the least fragmented and provided the least contaminated pore-water samples (Chester et al., 2013a). The δ³⁴S and Δ³³S values of sulfate in the 1R-1 (Unit 1), 8R-2 (Unit 3), and 20R-1 (Unit 5) samples ranged between +19.7‰ and +23.9‰ and between +0.053‰ and +0.061‰, respectively (Supplementary File S1). These ranges were similar to the δ³⁴S and Δ³³S values of seawater sulfate (ca. +21.3‰ and +0.050‰, respectively) (Ono et al., 2012; Tostevin et al., 2014). Therefore, pore-water sulfate appeared to be derived from seawater and sulfur isotope ratios were not modified from the original values. Since a large amount of sulfate was previously reported at the depth of the oceanic basement below decollement (Gieskes et al., 1990, 1993; Hensen and Wallmann, 2005), pore-water sulfate in 20R-1 may have been derived from the subseafloor environment.

Fig. 2 shows the chemical composition of pore-water samples. Fluid compositions were generally similar within Unit 1–3 samples, but differed between Unit 1–3 and Unit 5 samples. By assuming that all the sulfate detected in Unit 3 sample ana­lyses was derived from seawater contamination during sample treatments, the genuine composition of the other chemical components of Unit 3 pore-water samples may be estimated by the extrapolation of sulfate concentrations to zero. The estimated indigenous compositions of the selected ions of Unit 3 were similar to the seawater composition (Fig. 2). The same assumption of seawater contamination during sample recovery and treatment as that described above was applied to Unit 5 data; however, the estimated indigenous composition was not similar to the seawater composition. For example, the estimated indigenous concentrations of Ca and K (Fig. 2) were very high and low, respectively. Therefore, this assumption does not appear to be applicable to Unit 5 and the pore-water chemical composition elucidated for Unit 5 was not markedly affected by seawater contamination.

Fig. 2.

Pore-water chemistry. The ion concentrations of Cl, K, Ca, Mg, NH₄, and Mn (a, b, c, d, e, and f) are plotted as a function of sulfate ion concentrations. Symbol colors represent the lithological units of samples. Broken lines connecting the seawater composition with the highest and lowest concentrations in Unit 3 samples are shown to indicate estimated endmember concentrations by the extrapolation of sulfate to zero, drawn by vertical grey bars.

The pore-water of Unit 5 showed a quantitative change between Mg depletion and Ca enrichment as well as the slight loss of K and gain of Mn. The characteristics of the pore-water chemistry of Unit 5, higher Ca and lower Mg than the ambient seawater composition with a sufficient amount of sulfate, were consistent with those of the oceanic basement below sediment reported by other drilling projects (e.g. Elderfield et al., 1999; Cowen et al., 2003). These characteristics of oceanic basement water have been attributed to a low-temperature fluid-basalt interaction (e.g. Mottl and Holland, 1978; Mottl and Wheat, 1994). In a similar manner, the compositional and isotopic characteristics of sulfate in Unit 5 pore-water were affected by the precipitation of sulfate minerals, such as anhydrite, as well as diffusive sulfate loss to microbial sulfate reduction in the overlying prism sediments. Ammonium, the most probable product from the decomposition of sedimentary organic matter, was enriched in Unit 1–3 samples (>1‍ ‍mM), but not in Unit 5 samples.

The measured δD_H2O and δ¹⁸O_H2O values of pore-water were within a narrow range of between –0.02‰ and +3.82‰ and between –0.16‰ and +0.37‰, respectively (Table S1). Even if seawater (δD_H2O and δ¹⁸O_H2O of +0‰) significantly contaminated indigenous pore-water during sample recovery and processing (e.g., up to 40% in volume based on sulfate concentrations. See Chester et al., 2013a), variations in the δD_H2O value of indigenous pore-water were estimated to range between –0.1‰ and +6.4‰. Since this variation was markedly smaller than those in δD_H2 and δD_CH4 values described later, we assumed that the δD_H2O value of indigenous pore-water was +0‰ for later estimations of the isotopic fractionation effects of αD_CH4-H2O and αD_H2-H2O (see below).

Carbon and hydrogen isotope systematics of CH₄

CH₄ concentrations varied between 1–20‍ ‍mM in Units 1–3, but were lower than 5‍ ‍mM from decollement to the depths of Units 4–6 (Fig. 1). Variations in CH₄ concentrations were previously reported to be significantly affected by degassing during core recovery and processing (Chester et al., 2013a); therefore, the measured concentrations may be underestimated from in situ concentrations. δ¹³C_CH4 values were constant at approximately –64‰ in the upper half of Unit 3 and gradually decreased to –84‰ with increases in depth from the bottom half of Unit 3. δD_CH4 values were almost constant at approximately –200‰ through Units 1–3 and increased along with a decrease in the δ¹³C_CH4 value and increase in depth (Fig. 3a). δ¹³C_CO2 values ranged between –‍30‰ and +0‰. Vertical changes in δ¹³C_CO2 were similar to those in δ¹³C_CH4. The vertical profiles of δ¹³C_CH4 and δ¹³C_CO2 values were continuous and smoother than that of CH₄ concentrations (Fig. 1), suggesting the negligible effects of degassing during core recovery and processing on the isotopic composition.

Fig. 3.

Methane isotope systematics. Each panel shows isotope composition plots between (a) δDCH₄ and δ¹³CCH₄, (b) δ¹³CCH₄ and δ¹³CCO₂, and (c) αDCH₄-H₂O and α¹³CCH₄-CO₂. Symbol colors and shapes represent the lithological units of samples. The diagonal dotted lines in panel (b) represent α¹³CCH₄-CO₂ values. The temperatures at which the isotope equilibrium between CH₄ and H₂O exhibited the corresponding αDCH₄-H₂O values are shown on the right side of the panel (c) (Turner et al., 2021). The diagonal arrow in panel (c) represents the isotope effect associated with methanotrophy (αD/α¹³C=10).

Regarding the origin and behavior of CH₄, isotope systematics between CH₄ and the relevant C- and H-bearing molecules are shown in Fig. 3. The isotopic variation in CH₄ appeared to be associated with variations in lithological units (Fig. 3a). As observed in vertical profiles, δ¹³C fractionation between CH₄ and CO₂ was constant at approximately 60‰ through Units 2–5 (Fig. 3b). It corresponded to an α¹³C_CH4-CO2 value of approximately 0.94. The α¹³C_CH4-CO2 value was the smallest at Unit 1 and the largest at Unit 6. α¹³C_CH4-CO2 and αD_CH4-H2O values increased as the number of lithological units became higher (Fig. 3c).

Stable isotope ratios of H₂: observation and experiment

H₂ concentrations in most of the gas samples analyzed (n=44 out of 52) were lower than 3.0‍ ‍μM (Fig. 1). In contrast, three serial samples at depths of 690–700 mbsf (4R-5R) showed H₂ concentrations higher than 8‍ ‍μM, forming a sharp peak up to 209‍ ‍μM in the vertical profile (Fig. 1). Another H₂ peak up to 26‍ ‍μM was detected at depths of 817–820 mbsf (15R-16R). Although extrinsic H₂ generation during drilling and coring operations and sample processing is not completely excluded as the origin of the abundant H₂ concentrations, all gas samples were processed by the same procedure and anomalies occurred among similar lithological samples (Unit 3). The amplitude of the increase in H₂ from the background H₂ level was significantly larger than that of variations in CH₄ concentrations at the same depths because the concentrations of both of these gases are affected in a similar manner by the degassing process during core and sample recovery. Therefore, these H₂ peaks may reflect indigenous H₂ enrichment at specific depths.

To discuss the possible source of the indigenous H₂ enrichment observed, δD_H2 values were quantified. The δD_H2 values observed ranged between –850‰ and –350‰. The Keeling plot ana­lysis (Keeling, 1958), a plot between δD_H2 values and the reciprocal of the H₂ mixing ratio of the headspace gas, revealed a bimodal mixing trend between a D-rich low-H₂ abundance endmember and a D-depleted H₂-rich endmember with δD_H2 of –850‰, corresponding to αD_H2-H2O of 0.15 (Fig. 4a). Since we recognized unavoidable contamination by H₂ from air (0.5 ppmv with a δD_H2 value of +150‰ [Gerst and Quay, 2000]) during head space gas sampling in the vial, the presence of a D-rich H₂ endmember was reasonable. The single bimodal mixing trend and the estimated δD_H2 value of approximately –850‰ for the H₂-rich endmember through the samples suggested the following interpretations for the subseafloor H₂ source: (1) on-site H₂ generation occurring at multiple depths of the prism sediments, each of which showed an identical isotope effect, and/or (2) off-site H₂ generation followed by the distribution of H₂ to specific locations.

Fig. 4.

Molecular hydrogen isotope characteristics. Panel (a) shows a “Keeling plot” for the H₂ observed. The symbols in panel (a) are the same as those in Fig. 3. The diagonal line represents the least square linear fit for the observation and atmospheric H₂. Vertical grey bars represent isotope ratios reported for H₂ generated by a fault sliding experiment (Hirose et al., 2011) and observed in high-temperature hydrothermal fluid (HTHF) (Proskurowski et al., 2006; Kawagucci et al., 2010). Panel (b) shows the results of the Fe-H₂O experiment in the present study. Each symbol represents a batch of experiments. Symbol colors represent reaction temperatures. Diagonal dotted lines represent the αD_H2-H2O values selected for comparison. Panel (c) is a compilation plot between the values resulting from the Fe-H₂O experiment (the dataset is the same as that in panel (b)), αD_H2-H2O values and temperature for the estimated endmember value of the C0019E core observation, and the theoretical value at the isotope equilibrium between H₂ and H₂O (Horibe and Craig, 1995). Each symbol represents a batch of experiments. Symbol colors represent δD_H2O values at the experiment (3k means ca. +3000‰).

To address possible H₂ generation and distribution processes in subseafloor environments, laboratory experiments were conducted to evaluate isotopic effects associated with H₂ generation through a low-temperature Fe-H₂O reaction. Experimentally generated H₂ showed various δD_H2 values, while clear correlations were observed between δD_H2 and δD_H2O values at each reaction temperature (Fig. 4b). The αD_H2-H2O values of each batch experiment exhibited temperature-dependent correlations of 0.13, 0.19, and 0.23 at 4, 25, and 55°C, respectively (Fig. 4b and c). αD_H2-H2O values from the Fe-H₂O experiment were lower than known αD_H2-H2O values at the H₂-H₂O isotope equilibrium at the corresponding temperatures, such as 0.23, 0.26, and 0.31 at 4, 25, and 55°C, respectively (Horibe and Craig, 1995). The αD_H2-H2O values of Fe-H₂O experiments at 4–25°C were similar to that of the estimated D-depleted H₂ endmember in the C0019E observation (Fig. 4a and c).

Solid-phase biogeochemistry

Fig. 5 shows TOC, TN, and TS contents in the solid phase and the sulfur isotope composition of CRS. Core samples from units other than Unit 3 had a TOC content as low as 0.1% with TOC/TN ratios lower than 6. In contrast, samples from Unit 3 had a uniform TOC content of ~0.6% and a TOC/TN ratio of ~7 (Fig. 5a and Table S3). These characteristics are consistent with the terrigenous origin of Unit 3 introduced by the lithological description (Chester et al., 2013a).

Fig. 5.

Solid-phase chemistry with sulfur isotope ratios. Panels (a) and (b) shows plots of the relative nitrogen content (TOC/TN) and relative sulfur content (TS/TOC) with the total organic carbon content (TOC). Panel (c) shows a plot between TS/TOC and the sulfur isotope ratios of CRS in squeezed cakes. Panel (d) shows three isotope plots for CRS as well as dissolved sulfate in pore-water (blue diamonds). A cross symbol represents the isotope composition of juvenile sulfur in the mantle. A dotted curve represents a theoretical mixing trend between juvenile sulfur and the sample with the highest Δ³³S value. The arrow represents an isotope fractionation pattern associated with microbial sulfate reduction with 34ε of ~60 and 33λ of ~0.51 (see 3.4). Three samples with an anomalously high sulfur content are circled in panels (b-d). The filled circle in each panel represents a sample (15R-1) and its sulfur source is enigmatic.

To trace the biogeochemical cycle of sulfur compounds in the deep subseafloor, the compositional, microscopic, and isotopic characteristics of solid-phase sulfur were examined. Relative sulfur contents (TS/TOC) were generally constant at approximately 0.5 through the depths examined regardless of TOC, and were >5-fold higher in samples collected at depths of ~705‍ ‍m (6R-2) and ~800–810‍ ‍m (13R-1 and 14R-1) (Fig. 5b). Locally high TS/TOC ratios were attributed to the occasional sampling of heterogeneously distributed sulfur-bearing aggregates, as suggested by the frequent occurrence of visible (>100‍ ‍μm) ellipsoidal to elongate aggregates in the dark mudstone of Unit 3 identified through onboard core imaging (Chester et al., 2013a). Onshore SEM-EDS observations of thin sections from samples collected at depths of 715‍ ‍m (7R-2) and 802‍ ‍m (13R-2) revealed the distribution of iron-sulfide aggregates with framboidal or acicular structures (Fig. 6). Based on appearances and Fe/S stoichiometry, these aggregates were mainly pyrrhotite with minor pyrite. Similar sulfide aggregates were identified in the same core sections by another microscopic study (Yang et al., 2018). Framboidal aggregates were scattered in the specimens, while acicular aggregates were only concentrated in discolored parts near the crack of core samples.

Fig. 6.

Two types of sulfide aggregates observed in Unit 3. (a) Framboidal aggregates and (b) an acicular aggregate with an EDS point ana­lysis and mapping. White/black images are back-scattered SEM images, and colored images are mapping results for each element. Each upper right spectrum is a point elemental ana­lysis at the point showing the black cross in each back-scattered image.

The δ³⁴S values of CRS in the analyzed interval were markedly lower (–50 to –10‰) than those of pore-water sulfate (ca. +22‰) (Fig. 5c). Moreover, some Δ³³S values of CRS (up to +0.16‰) were markedly higher than those of sulfate (Fig. 5d). It is important to note that the δ³⁴S and Δ³³S values of CRS in high TS/TOC samples were close to +0‰, corresponding to the values of juvenile sulfur from a mantle reservoir (e.g., Ono et al., 2012). Only one sample from 816.8‍ ‍m (15R-1) showed anomalous isotopic ratios with both positive δ³⁴S and Δ³³S values.

The apparent largest isotope fractionation between seawater sulfate and CRS, ³⁴ε of 67.8‰ and ³³λ of 0.514, was detected in Unit 1 (1R-1). This ³⁴ε value was larger than the maximum ³⁴ε value observed via microbial sulfate reduction in laboratory cultivations (Sim et al., 2011) and was within the ³⁴ε range in modern marine sediments (c.f., Wortmann et al., 2001; Canfield et al., 2010; Drake et al., 2015). CRS with the lowest δ³⁴S value in Unit 3 and seawater sulfate had ³⁴ε and ³³λ values of 55.3‰ and 0.514, respectively.

Microbial cellular biomass and taxonomic composition

Prior to descriptions of microbial cell density and 16S rRNA gene amplicon sequences, it is important to note that data were obtained from core samples with a low biomass and potent external seawater contamination, as discussed above for pore-water sulfate. Therefore, results may have been affected by contamination and the biases of external seawater microbial populations as well as analytical methods during core recovery, processing, and laboratory experiments. Possible contamination by drilling mud fluids in microbiological core subsamples was measured onboard (Chester et al., 2013a), and most of the core samples (except for 6R-2) used for microbial cell density and 16S rRNA gene amplicon sequencing ana­lyses were not markedly affected by contamination by drilling mud fluid (Chester et al., 2013a).

Microbial cell densities in core samples ranged between 4.4×10⁴ and 6.3×10⁵‍ ‍cells‍ ‍mL^–1 of sediment (Table S2). According to the global biomass distribution in subseafloor sedimentary environments, a similar cell density was detected in sediments collected at similar depths below the subseafloor (Magnabosco et al., 2018). The taxonomic compositions of core samples are shown in Fig. 7 and Fig. S2. The abundance of each phylogenetic group differed between lithological units; however, predominant members at the phylum level were common throughout the units; within Bacteroidota (11–66%), Pseudomonadota formerly named Proteobacteria (14–39%), Bacillota formerly named Firmicutes (4.2–32.3%), and Cyanobacteria (4–20%). In addition, Flavobacteriaceae members dominated within Bacteroidota, (80.7–96.9%) and accounted for approximately 60–70%, except 6R at 96%, in which the marine clusters NS2b, NS4, and NS5 and the genera Tenacibaculum, Formosa, Cloacibacterium, and Aurantivirga were core members. Gammaproteobacteria (5–18%) and Alphaproteobacteria (5–20%) members were the predominant components within Pseudomonadota, Oleispira, and Pseudoalteromonas were the predominant members of Gammaproteobacteria, and SAR11 and Rhodobacteraceae, including the genera Amylibacter, Ascidiaceihabitans, and Cognatiyoonia, were the dominant population for Alphaproteobacteria. On the other hand, Chloroflexota (0.2–6.4%), Atribacterota (0.4–18.5%), Planctomycetota (0.3–7.8%), and “Aerophobetes” (0.01–8.7%), which were frequently found in anoxic subseafloor sediments (Inagaki et al., 2006; Nunoura et al., 2016; Hoshino et al., 2020), were also detected as less predominant members than the above components, even though these members were very abundant populations in the core samples of 10R and 14R. Bacteroidota, Pseudomonadota, and Cyanobacteria are key members of planktonic microbial communities in global seawater environments (Hoshino et al., 2020), similar to the aforementioned pore-water sulfate contamination by seawater sulfate; therefore, the OTUs within these bacteria may be attributed to the contamination of seawater samples during sample recovery and processing (see Materials and Methods). Bacteroidota and Pseudomonadota were previously reported as major taxonomic members within drilling fluids (Yanagawa et al., 2013; Martino et al., 2019), even though the onboard contamination ana­lysis indicated that contamination by drilling mud fluid was negligible. Members of Pseudomonadota and Flavobacteriaceae were also detected in the shallow sediments of the Japan Trench (Hiraoka et al., 2020), Bacillota, Bacteroidota, and Pseudomonadota were predominantly found in the ultradeep sediments of the Japan Trench forearc basin, and members of Chloroflexota and Atribacterota were identified in the same environments (Inagaki et al., 2015); therefore, 16S rRNA gene compositions in prism sediments may be affected by the biodiversity of indigenous microbial communities.

Fig. 7.

Summary of microbiological ana­lyses.

Taxonomic composition of 16S rRNA gene amplicon sequences in sediments at site C0019E. Phylum-level taxonomic composition based on SSU rRNA gene amplicon sequencing using universal primers.

The physiological properties of microorganisms corresponding to most of the OTUs constructed remain unknown, and the in situ functions of these potentially indigenous microbial populations remain unclear. However, if most of the OTUs within Bacteroidota and Pseudomonadota came from core samples, the predominant OTUs were most likely derived from the microbial taxa represented by a limited number of cultivated heterotrophic members, and some of the members of Atribacterota and Chlorofrexiota (e.g. Anaerolineae, Dehalococcoidia) have recently been identified as anaerobic heterotrophs (Kaster et al., 2014; Katayama et al., 2020; Yang et al., 2020). Most of the microbial populations in Japan Trench prism sediments represented by the 16S rRNA gene amplicon sequencing ana­lysis appeared to be associated with the decomposition of refractory organic matter. Anaerobic methane oxidation coupled with sulfate reduction were experimentally inferred from carbon and sulfur isotope ana­lyses, and a few populations of sulfate reducers within the families Desulfobacteraceae and Desulfobulbaceae, both of which are associated with methanotrophic archaea (Dekas et al., 2016), were detected in 13R and 16R samples. The presence of methanogens was not indicated from amplicon sequences and additionally performed PCR amplifications using specific primers of the genes for mcrA.

Metabolic potentials and cultivations

The metabolic potentials of chemolithoautotrophic populations were examined using ¹⁴C-labeled substrates and detectable activities were found in several core samples (Table S2). In incubations with ¹⁴C-labeled bicarbonate, active homoacetogenesis (¹⁴C-labeled CO₂ to CH₃COOH) was observed in 5R, 8R, and 14R samples, while hydrogenotrophic methanogenesis (¹⁴C-labeled CO₂ to CH₄) was not present in any samples. Methanogenesis from ¹⁴C-methylamine was detected in the 13R sample and the anaerobic oxidation of ¹⁴C-labeled CH₄ to CO₂ in 1R, 6R, and 13R samples.

Based on the results obtained from the metabolic activity assessment of the microbial community, cultivations for homoacetogens as well as hydrogenotrophic and methylotrophic methanogens were conducted. Successful cultivation was achieved in the homoacetogenic medium with the 1R, 5R, 6R, 7R, 8R, and 13R samples, and the dilution cultivation technique estimated the cultivable population densities of homoacetogens in indigenous core samples as between 9.2×10² and 9.6×10⁴‍ ‍cells‍ ‍mL^–1 of sediment (Table S2). In addition, a positive cultivation for the methylotrophic methanogen population was obtained from the 13R sample and the estimated cultivable population density was 9.6×10⁴‍ ‍cells‍ ‍mL^–1 of sediment. Enriched cultures in the most diluted media were subject to dilution-to-extinction techniques and the 16S rRNA gene sequences of isolates were elucidated. The 16S rRNA gene sequences of isolates from the homoacetogenic medium showed >99.9% similarity with each other, and were closely related to Acetobacterium carbinolicum (99%). The methylotrophic methanogen isolate was closely related to Methanolobus taylorii (99%). Although A. carbinolicum was isolated from subseafloor sediments of the Nankai Trough (at a water depth of 4,791‍ ‍m and at 4.15 mbsf) (Toffin et al., 2004) and shallow marine sediments (Purdy et al., 2003), there has been no cultivation example of related strains from hadal subseafloor sediments. Although the 16S rRNA amplicon sequence ana­lysis did not detect the signatures of the genera Acetobacterium and Methanolobus through core samples, they may represent indigenous populations not detected by the amplicon sequence ana­lysis. In cultivation tests, no hydrogenotrophic methanogens were obtained from any core samples. The cultivation of anoxic methanotrophs with sulfate reduction was not performed.

Thermodynamic calculation

The energetic potentials of hydrogenotrophic methanogenesis, homoacetogenesis, and aceticlastic methanogenesis were theoretically evaluated as ΔG_r changes in response to changes in activities, a[H₂] and a[Ac] values (Fig. 8). Homoacetogenesis and hydrogenotrophic methanogenesis were both exergonic at a[H₂] as low as the detection limit of the H₂ concentration used in the present study (~10^–6 M), but were not exergonic under a possible H₂ background level of ~10^–8 M, quantified at other deep subseafloor areas with a more sensitive method (Lin et al., 2012; Heuer et al., 2020). When the concentration of acetate was constant as low as 1‍ ‍μM, corresponding to a[Ac] of 2.8×10^–10, the ΔG_r pattern with changes in a[H₂] was similar between homoacetogenesis and hydrogenotrophic methanogenesis, while acetoclastic methanogenesis did not progress at any a[H₂]. When the concentration of acetate was constant as high as 1‍ ‍mM, corresponding to a[Ac] of 2.8×10^–7, the minimum threshold of a[H₂] for homoacetogenesis (~10^–7) was an order of magnitude higher than that of hydrogenotrophic methanogenesis (~10^–8), while acetoclastic methanogenesis progressed at any a[H₂]. The minimum threshold of a[Ac] for acetoclastic methanogenesis was ~10⁹, corresponding to an acetate concentration of ~10‍ ‍μM, as high as the background level of the acetate concentration in high-TOC sedimentary environments (Heuer et al., 2009; Ijiri et al., 2012).

Fig. 8.

Gibbs free energy changes along with substrate activities for representative metabolism. Changes in Gibbs free energy changes with (a) a[H₂] and (b) a[Acetate] for hydrogenotrophic methanogenesis, homoacetogenesis, and acetoclastic methanogenesis under the given conditions (pH=7.9, ΣCO₂=50‍ ‍mM, and CH₄=5‍ ‍mM) are shown. Horizontal dashed-dotted and dotted lines with dark and light green colors represent acetoclastic methanogenesis at ΣCH₃COOH of 1 and 0.001‍ ‍mM, respectively.

Discussion

CH₄ generation and maturation related to subseafloor microbial communities

CH₄ was abundant in core samples of all accretionary prism sediments and appeared to be derived from microbial methanogenesis in situ and/or in proximal environments because of the in situ low-temperature condition (lower than 30°C) and the relatively low geothermal gradient of the Japan Trench accretionary wedge (Fulton et al., 2013). The contribution of thermogenic hydrocarbons was previously suggested to be negligible due to the low temperature, which was reflected by the high methane/ethane ratios (~10³) observed (Chester et al., 2013a).

CH₄ generation and maturation related to subseafloor microbial communities in the prism sediment are hereafter discussed based on isotope systematics between CH₄ and the relevant C- and H-bearing molecules. Low δ¹³C_CH4 values (–‍84‰–64‰) and α¹³C_CH4-CO2 values (0.92–0.94) (Fig. 3b) may both be attained through large kinetic isotope effects via hydrogenotrophic methanogenesis under the H₂-limited condition (α¹³C_CH4-CO2 of 0.91–0.94) (Valentine et al., 2004; Penning et al., 2005; Okumura et al., 2016). Although the microbial population and function of hydrogenotrophic methanogens were not justified by cultivation-dependent or -independent ana­lyses in the present study, hydrogenotrophic methanogenesis associated with the fermentative decomposition of sedimentary organic matter may be operative in the accretionary wedge, which has frequently been suggested in similar subseafloor sedimentary environments (e.g. Inagaki et al., 2015; Ijiri et al., 2018; Beulig et al., 2022).

Besides hydrogenotrophic methanogenesis, methylotrophic methanogenesis using methyl group carbons may also exert a large kinetic isotope effect (α¹³C_{CH4-substrate} of 0.93–0.95) (Krzycki et al., 1987; Summons et al., 1998). The activity of methylotrophic methanogenesis was detected in the core sample of 13R by both radioactive isotope tracer and cultivation tests (Fig. 7). Due to the terrigenous origin, the sediments of Unit 3 may host the typical metabolic products of coastal biota, such as trimethylamine N-oxide (TMAO) (Yancey et al., 1982) and dimethylsulfoniopropionate (DMSP) (Kiene et al., 2000), which may be substrates for methylotrophic methanogenesis. Therefore, methylotrophic methanogenesis may contribute to the abundance of biogenic CH₄ in prism sediments as well as hydrogenotrophic methanogenesis.

The present study showed that the homoacetogenic population and its functions dominated the methanogenic population. The microbial processes initiated from homoacetogenesis may be potential sources of biogenic CH₄. One example is aceticlastic methanogenesis directly using the end product (acetate) of homoacetogenesis, while another is hydrogenotrophic methanogenesis syntrophically coupled to acetate oxidation (to H₂ and CO₂) (Hattori, 2008), which may be catalyzed by different acetate-oxidizing populations from homoacetogens and even by the reverse acetogenesis reaction of homoacetogens. The apparent kinetic isotope effects of these methanogenic processes via homoacetogenesis are caused by the sequence of homoacetogenesis (α¹³C_Ac-CO2=0.94) (Gelwicks et al., 1989) and aceticlastic methanogenesis (α¹³C_CH4-Ac=0.98–0.99) (Valentine et al., 2004; Vinson et al., 2017), which may have provided the low δ¹³C_CH4 and α¹³C_CH4-CO2 values observed.

While stable carbon isotopic signatures (e.g., α¹³C_CH4-CO2 values) provide an insight into the source of extant CH₄, δD_CH4 and αD_CH4-H2O values are associated with not only the source, but also the history of maturation. The constant αD_CH4-H2O values observed in the present study (~0.81) (Fig. 3c) suggest that extant CH₄ in the prism sediments was not produced in recent years, it had accumulated through geological time. Biogenic CH₄ has been shown to exhibit αD_CH4-H2O values of ~0.7 at its generation (e.g. Sugimoto and Wada, 1995). On the other hand, a compilation of αD_CH4-H2O values from various experiments and observations (Okumura et al., 2016) revealed that αD_CH4-H2O values higher than 0.8 were the typical signatures of aged CH₄ in geological samples regardless of their origin, as indicated by α¹³C_CH4-CO2 values. The αD_CH4-H2O values of geological CH₄ are consistent with the values of isotope equilibrium between CH₄ and H₂O at in situ temperatures (Gropp et al., 2021; Turner et al., 2021) (Fig. 3c). This was also the case for Units 1–3, for which αD_CH4-H2O thermometer values (0–50°C) (Fig. 3c) were consistent with the in situ temperature (<30°C) (Fulton et al., 2013).

The slightly high δD_CH4 values (>–180‰) found in the bottom of Unit 3 as well as Units 4–5 (Fig. 3a) appeared to be affected by microbial methane consumption, which enriches D and ¹³C with a δD/δ¹³C ratio of ~10 in remnant CH₄ (Feisthauer et al., 2011; Kawagucci et al., 2021) (Fig. 3c). The metabolic function of anaerobic methane oxidation coupled with sulfate reduction was verified at several depths of the prism sediments by the radioactive isotope tracer assay (Table S2) despite this function not being detected from the boundary layer between Units 3 and 4. In addition, the vertical gradient of methane concentrations through the sediments between the bottom of Unit 3 and Units 4–5 (Fig. 1 and 2) indicated the occurrence and function of microbial methanotrophy.

H₂ generation and its distribution related to earthquakes and faults

In many oceanic and terrestrial subsurface sediments, H₂ concentrations are generally lower than 0.1‍ ‍μM (Lin et al., 2012). In the present study, H₂ concentrations in most parts of C0019E core sediments were below the limit of quantification (<3‍ ‍μM) (Fig. 1). Unless local H₂ inputs are present, a low H₂ level is maintained most likely by subseafloor microbial functions, particularly by syntrophic microbial communities between hydrogenogenic fermenters and hydrogenotrophic chemolithotrophs. Therefore, kinetic properties for their utilization of H₂ contribute to a steady-state H₂ concentration (e.g. Lovley and Goodwin, 1988).

The sharp peaks observed in H₂ concentrations at 690–700 and 817–820 mbsf in the present study (Fig. 1) demonstrated potential local H₂ inputs. At approximately 700 mbsf, significant low gamma ray and resistivity signals were also identified by logging-while-drilling measurements (Chester et al., 2013a). These findings suggest relationships between fault formation and local H₂ inputs, leading to possible interpretations such as (1) on-site H₂ generation around the faults and/or (2) remote H₂ generation(s) and transportation via fault formation and fluid migration. Several processes have been reported for the abundant generation of H₂ associated with fault activity: the water-rock redox reaction (Mayhew et al., 2013), water radiolysis (Lin et al., 2005), and mechanochemical water reduction associated with fault activity (Kita et al., 1982; Hirose et al., 2011) as well as the thermal decomposition of sedimentary organic matter (e.g. Kawagucci and Seewald, 2019).

Based on the field observations and possible H₂ generation mechanisms described above, we hereafter discuss the source and distribution of H₂ anomalies at specific depths using the hydrogen isotope ratio as a tracer. The highly D-depleted signature of H₂ found at the peak of 700 mbsf (δD_H2=–850‰ and αD_H2-H2O=0.15) (Fig. 4a) strongly suggested that H₂ was generated by a low-temperature process because of greater isotope effects (smaller αD_H2-H2O) at lower temperatures. Based on the simulated experiments in the present study (Fig. 4b), a strong kinetic isotope effect (αD_H2-H2O value=0.15) was attainable by the generation of H₂ via the low-temperature Fe-H₂O reaction. Although zero valent iron used in the experiment was absent in the sedimentary environment, the strong kinetic isotope effect on H₂ generation associated with the Fe-H₂O reaction is of importance for the isotope effect associated with H₂ generation through other low-temperature water-mineral interactions possibly occurring in the subseafloor, such as the Fe(II)-H₂O interaction (Mayhew et al., 2013).

In contrast, other processes for H₂ expected around the fault zone cannot cause the D depletion in H₂ observed. A simulated experiment of high-velocity fault sliding, resulting in the partial melting of rocks at the sliding surface, generated H₂ with the D-rich signature of αD_H2-H2O of 0.8 (Hirose et al., 2011) (Fig. 4a). High-temperature hydrothermal fluids in typical deep-sea hydrothermal vents contain H₂, the isotope signatures of which are identical to the isotope equilibrium between H₂ and H₂O at in situ temperatures, namely, αD_H2-H2O of 0.6–0.7 at 300–400°C (e.g. Proskurowski et al., 2006; Kawagucci et al., 2010) (Fig. 4a). A previous thermal decomposition experiment on sedimentary organic matter also generated H₂ isotopically equilibrated with H₂O at given in situ high temperatures (Kawagucci and Seewald, 2019). Furthermore, microbial H₂ metabolism, including its production and consumption, has been shown to promote the isotope equilibrium between environmental H₂ and H₂O (Walter et al., 2012; Kawagucci et al., 2014), while the H₂-H₂O isotope equilibrium results in a αD_H2-H2O value of not lower than 0.2 even at 0°C (Fig. 4c). These findings suggest that D-enriched H₂ produced by high-temperature processes cannot be transformed to highly D-depleted H₂ in the present study even through the later low-temperature isotopic equilibrium was mediated by the microbial metabolism of H₂. This is one reason why we excluded the interpretation that remote H₂ generation is associated with fault sliding and a subsequent hydrothermal reaction followed by H₂ transportation via fault formation and fluid migration. Nevertheless, remote hot H₂ generation followed by the subsequent stimulation of hydrogenotrophic activity that had previously occurred and had since ceased may be associated with possible geochemical and microbiological processes, such as methanogenesis and homoacetogenesis, other than the specific peaks of D-depleted H₂.

H₂ generation through a low-temperature water-rock interaction may progress in natural fault zones. The significant generation of H₂ was confirmed in simulated experiments between spinel-bearing rocks and water at 55°C within 10 days (Mayhew et al., 2013). H₂ generation was previously attributed to electron transfer from Fe(II) to H₂O adsorbed on fresh spinel surfaces (Mayhew et al., 2013). Since the rock surface in the subseafloor sedimentary lithosphere has been virtually passivated by changes through geological time, the Fe(II)-H₂O interaction was generally not expected (Fig. 9A). On the other hand, even after these changes, sedimentary rock fragments still possessed an intact mineral composition in the inner part surrounded by the changed skin. Earthquake rupture exposes intact minerals to the rock surface and this is followed by subsequent interactions between the new surface and pore-water (Fig. 9B). Therefore, we speculate that relatively recent earthquakes (aftershocks after the 2011 Tohoku-oki earthquake) may have induced fault rupture at a depth of 700 mbsf at the drilling site and induced on-site H₂ generation around the faults through the low-temperature water-rock redox interaction.

Fig. 9.

Illustration for the hypothesis. A hypothesis for microbial ecosystem dynamics at C0019E is shown. Panel A: The steady-state mode. [a] A syntrophic relationship between H₂-generating fermentation and hydrogenotrophic methanogenesis as well as [b] methylotrophic methanogenesis continue to the accumulation of CH₄. Panel B: Earthquake event. [c] On-site H₂ generation associated with a low-temperature fluid-rock interaction and [d] H₂S and H₂ accompanied by coseismic hydrothermal fluid upwelling. Panel C: [e] homoacetogenesis and [f] accumulated acetate-based acetoclastic methanogenesis and reverse acetogenesis followed by hydrogenotrophic methanogenesis.

Generation and distribution of solid-phase sulfur related to coseismic hydrothermal fluid flows

Processes associated with fault activity, suggested by carbon and hydrogen biogeochemistry as well as microbial functions, were examined in more detail based on solid-phase sulfur observations. The very low δ³⁴S values with positive Δ³³S values for CRS observed in Units 1 and 3 strongly suggest the involvement of microbial sulfate reduction in CRS production (Fig. 5d). Microbial sulfate reduction commonly occurs in relatively shallow subseafloor sediments during the early diagenesis of sedimentary organic matter and/or the sink of upwelling (seeping) CH₄ from deep subseafloor reservoirs. CRS in Units 1 and 3 had been generated under sulfate-available conditions at shallow depths in the past and was then further buried with time.

Sulfur isotope systematics and TS/TOC ratios (Fig. 5d) as well as the sporadic distribution of aggregates (Fig. 6) may be attributed to the combination of juvenile sulfur transported by coseismic hydrothermal fluids (Yang et al., 2018) and biogenic sulfide. Deep earthquake-induced processes near seismic centers, such as the frictional heating of basalt and peridotite, leads to a high-temperature fluid-rock interaction that forms coseismic hydrothermal fluids containing H₂S with the juvenile sulfur isotope signature. Coseismic hydrothermal fluids then migrate through permeable fractures around the fault zones, followed by cooling, resulting in the local precipitation of sulfide minerals. Yang et al. (2018) indicated that framboidal pyrite/pyrrhotite aggregates formed when coseismic hydrothermal fluids encountered previously formed framboidal pyrite, while acicular pyrite/pyrrhotite aggregates formed when coseismic hydrothermal fluid leached the pre-existing pyrite and cooled within the fractures (Fig. 6). Although the effluent of coseismic hydrothermal fluids with deep-sourced CH₄ and ³He-rich helium was captured around the epicenter 36 days after the 2011 Tohoku-oki Earthquake (Kawagucci et al., 2012; Sano et al., 2014), it remains unclear whether the sulfides observed in the present study originated from historical or recent earthquakes.

The presence of sulfides derived from coseismic hydrothermal fluid is consistent with the absence of hydrothermal fluid-derived H₂ observed in the core samples as described above. Although seafloor high-temperature hydrothermal fluids contain both H₂ and H₂S (German and Seyfried, 2014; Nakamura and Takai, 2014), hydrothermal fluid-derived H₂ is rapidly consumed by microbial metabolism without the precipitation of minerals. Therefore, any trace (even the D-enriched isotopic signature of high-temperature H₂ generation) of hydrothermal fluid-derived H₂ is expected to disappear, in contrast to H₂S and sulfide minerals.

Subseafloor microbial populations and functions

Subseafloor microbial communities in Japan Trench accretionary prism sediments without H₂ enrichment may be sustained by sedimentary organic matter, which fuels the syntrophic populations of hydrogenogenic fermenters and hydrogenotrophic microbes, including methanogens (e.g. Lovley and Goodwin, 1988; Imachi et al., 2011) (Fig. 9A). This sedimentary organic matter-based community, which may consist of Bacteroidota, Pseudomonadota, Chloroflexota, Atribacterota, and Planctomycetota, has also been identified in other deep subseafloor biospheres at the forearc basin of the Japan Trench and the accretionary prism of the Nankai Trough (Inagaki et al., 2015; Nunoura et al., 2016; Ijiri et al., 2018; Beulig et al., 2022). In addition, the population and metabolic functions of the methylotrophic methanogens detected in both the tracer activity ana­lysis and cultivation tests may be representative of the steady-state ecosystem in prism sediments (Fig. 9A). These steady-state communities and functions may result in the accumulation of CH₄ at very slow rates through geological time since the oxidants capable of exothermically metabolizing CH₄, such as sulfate, are not sufficient due to the distance from the seafloor and decollement.

The predominant potential of homoacetogenesis, not hydrogenotrophic methanogenesis, in the accretionary prism sediments under H₂-rich conditions (Fig. 9C) was demonstrated by both radioactive isotope tracer ana­lyses and cultivation tests, while both shared the same substrates, H₂ and CO₂. Since the thermodynamic evaluation revealed that homoacetogenesis and hydrogenotrophic methanogenesis are energetically available under the given conditions of the laboratory experiments as well as the in situ H₂-rich sediments (Fig. 8), the out-competition of homoacetogenesis may be attributed to differences in the kinetic nature between these hydrogeneotrophic chemolithtrophic populations under H₂-rich conditions. The successful cultivation of homoacetogens was achieved by the medium with the 200-kPa headspace gas of 80% H₂ (an order of 10⁵ Pa-H₂ or 10³ μM H₂). Significant acetogenic activities based on the radioactive isotope tracer assay were also obtained from media with the 100-kPa headspace gas of 1% H₂ (an order of 10³ Pa-H₂ or 10¹ μM-H₂). To date, the kinetic out-competition of homoacetogenesis under these H₂-enriched conditions at low temperatures has been verified in laboratory cultivation experiments using several homoacetogen and methanogen isolates (Kotsyurbenko et al., 2001).

At the time of the drilling expedition 1 year after the 2011 Tohoku-oki earthquake, we found H₂ enrichment (up to an order of 10⁴ Pa-H₂) and the significant potential of homoacetogenesis-predominant subseafloor microbial communities and functions (Fig. 9A). It remains unclear whether the tempo and mode in microbial communities and functions rapidly change in short periods or gradually in a geological time scale. However, from a local scale point of view, earthquake-induced H₂ enrichment may be rapidly consumed by the response and enhancement of home-acetogenic populations and functions. Even after H₂ consumption by homoacetogenic populations, the reduced concentrations of H₂ at the minimum threshold for acetogens may still be available for other hydrogenotrophic populations. The theoretical calculation of metabolic thermo­dynamics indicates the exothermicity of hydrogenotrophic methanogenesis at lower concentrations of H₂ than the thermo­dynamic minimum threshold for acetogens (Fig. 8). Labo­ratory cultivation experiments (Kotsyurbenko et al., 2001) reported that the minimum threshold for H₂ uptake was approximately 10¹ Pa-H₂ for acetogens and <10⁰ Pa-H₂ for‍ ‍methanogens. In addition, a number of experiments using natural sediments demonstrated that the minimum threshold for mesophilic hydrogenotrophy was 10^–1–10⁰ Pa-H₂ (e.g. Lovley and Goodwin, 1988). Although the hydrogenotrophic metabolic activity in accretionary prism sediments at this low H₂ level was not examined in the present study, multiple lines of evidence support the potential occurrence of hydrogenotrophy other than homoacetogenesis under the low H₂ condition in the far-post-earthquake/pre-earthquake event era as a part of the steady-state ‘normal’ ecosystem.

Conclusions and future perspectives

The geochemical and microbiological characteristics of the accretionary prism sediments of the Japan Trench 1 year after the 2011 Tohoku-oki earthquake were investigated. The 2011 Tohoku-oki earthquake was a type of millennial mega-earthquake occurring in similar segmentations of Japan Trench seismogenic zones (Ikehara et al., 2016; Kodaira et al., 2021). However, the geochemical and microbiological processes discussed in the present study were more frequent, not millennial, events because they progressed in a local scale and earthquake events follow the Gutenberg and Richter relationship (lower magnitude earthquakes occurring at a higher frequency). It is important to note that the sequence of seismicity, fluid-rock interactions, and microbial responses may occur anywhere in seismically-active subseafloors, not only in the Japan Trench. Nevertheless, since the primary chemical force is highly mobile and metabolizable H₂, post-earthquake biogeochemical processes and microbial populations may be episodic and rapidly return to the steady-state ‘normal’ ecosystem (Fig. 9A) and, thus, may be difficult to detect even by scientific drilling operations, except in expeditions just after a mega-earthquake.

When an earthquake occurs in deep seismogenic zones and fault formation and sliding are activated and widespread, coseismic hydrothermal fluids are generated by frictional heating and migrate to accretionary prism sediments along the faults, which brings deep sources of H₂ and H₂S followed by sulfide precipitation due to the conductive cooling of the fluid (Fig. 9B). At accretionary prism sediments, the fresh surface of crushed rock around the faults interacts with pore-water, which serves as on-site H₂ generation until the surface is completely changed. Earthquake-induced H₂ inputs episodically drive the development of homoacetogenesis-predominant subseafloor microbial communities (Fig. 9C). When in situ H₂ concentrations are below the minimum threshold for homoacetogenic metabolism, other hydrogenotrophic populations, such as methanogens, further consume H₂ until their minimum threshold. Acetate, which accumulates as a result of temporal homoacetogenesis at a high H₂ level, is also finally decomposed at a low H₂ level into CH₄. Thermodynamic calculations (Fig. 8) suggest the potential of acetoclastic methanogenesis and/or the acetate-dependent H₂-syntrophic metabolism of reverse homoacetogenesis coupled with hydrogenotrophic methanogenesis (Fig. 9C). When episodic acetate supplied by homoacetogenic metabolism is virtually consumed, subseafloor microbial communities return to the steady-state tempo and mode (Fig. 9A).

However, there is still a lack of direct evidence to justify the hypothesis proposed here. For example, we did not obtain any data on in situ acetate concentrations for a quantitative discussion about microbial acetate production and consumption. The scare and coarse intervals of core sampling for onshore microbiological studies do not allow us to clarify the possible spatial transition of the metagenomics-based community structure around the H₂ anomaly. If larger volumes of core samples are available, more specific cutting-edge ana­lyses, such as clumped isotopes of hydrocarbons (Eiler et al., 2014; Gilbert et al., 2019; Taguchi et al., 2020) as well as metatranscriptomic comparisons of microbial communities (Orsi et al., 2013; Zinke et al., 2017, 2019), may be performed to justify the hypothesis. Nevertheless, JFAST drilling and geochemical and microbiological ana­lyses of core samples provide insights for the first time into the hadal and seismically active deep subseafloor biosphere. The relationship between earthquakes and the deep subsurface ecosystem will be the focus of future scientific drilling projects.

Citation

Kawagucci, S., Sakai, S., Tasumi, E., Hirai, M., Takaki, Y., Nunoura, T., et al. (2023) Deep Subseafloor Biogeochemical Processes and Microbial Populations Potentially Associated with the 2011 Tohoku-oki Earthquake at the Japan Trench Accretionary Wedge (IODP Expedition 343). Microbes Environ 38: ME22108.

https://doi.org/10.1264/jsme2.ME22108

Acknowledgements

We would like to express our deepest appreciation to Dr. Uta Konno for his efforts with carbon and hydrogen isotope ana­lyses in JAMSTEC. Dr. Kenji Shimizu kindly contributed to solid-phase sulfide observations, sample preparation, and ana­lyses and Dr. Takayuki Ushikubo provided advice on observations. We are grateful to all the people realizing and supporting the JFAST expedition of DSV Chikyu. Parts of this study were supported by the Ministry of Education, Culture, Sports, Science, and Technology (MEXT) Grants-in-Aid (KAKEN-HI) nos. 17H01869 (to SK and SS), 20H02020 (to SK), 17H06455 (to TS), 23224013 (to NY), and 17H06105 (to NY).

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