1 Introduction

Active hydrothermal systems in the Okinawa Trough (OT) are representative of one of the four TAIGAs, i.e., the ‘TAIGA of Methane’ (Urabe et al. Chap. 1). Numerous cruises for more than two decades have investigated the hydrothermal activities and fluid geochemistry of high-temperature hydrothermal fields within the OT, such as Minami-Ensei (Chiba et al. 1993; Kawagucci et al. 2013a), Iheya North (Nakagawa et al. 2005; Kawagucci et al. 2011, 2013b), Jade (Sakai et al. 1990a, 1990b; Ishibashi et al. 1995), Hakurei (Ishibashi et al. 2014), Hatoma (Nakano et al. 2001; Saegusa et al. 2006; Kawagucci et al. 2010b; Toki et al., in preparation), and Daiyon-Yonaguni (Yonaguni IV) (Kishida et al. 2004; Konno et al. 2006; Suzuki et al. 2008). The most notable characteristics of the fluid chemistry of the OT hydrothermal fields are abundant millimolar levels of CH4 (Fig. 30.1 and Table 30.1). This is why the OT hydrothermal systems are representative of the ‘TAIGA of Methane.’

Fig. 30.1
figure 1figure 1

Compilation of the physical properties and fluid chemistry of the Okinawa Trough and global hydrothermal fields. The vertical axis (methane concentration) is in logarithm scale. Horizontal axes are common for all panels and represent each hydrothermal field as follows: 1-Minami-Ensei, 2-Iheya North, 3-Hakurei, 4-Jade, 5-Hatoma, 6-Daiyon-Yonaguni, 7-vacant, 8-Endeavour, 9-Escanaba, 10-Guaymas, 11-Middle Valley, 12-vacant, 13-Suiyo, 14-North Fiji, 15-Roman Ruins, 16-Satanic Mills, 17-Vienna Woods, 18-Kilo Moana, 19-Tow Cam, 20-ABE, 21-Mariner, 22-Brother, 23-vacant, 24-Kairei, 25-Edmond, 26-vacanct, 27-Manez Gwen, 28-Lucky Strike, 29-Broken Spur, 30-TAG, 31-Snake Pit, 32-Rainbow, 33-Logatchev, 34-Ashadze, 35-vacant, 36-EPR21N, 37-EPR13N. Vertical dark and light gray bars separate different geological/tectonic settings. Horizontal gray bars in several panels represent seawater levels. References for this figure are presented in Sect. 30.2.2

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Table 30.1 Fluid chemistry of the Okinawa Trough hydrothermal fields
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The chemical characteristics of venting hydrothermal fluid are primarily constrained by high-temperature fluid-rock interactions at deep hydrothermal reaction zones (Seewald and Seyfried 1990; Seyfried et al. 2003; German and von Damm 2003). Phase separation and subsequent phase segregation occurring at deep high-temperature region also affect the chemistry of the venting fluid (Gamo 1995). Furthermore, certain microbial and chemical effects on hydrothermal chemistry have been identified in the branched low-temperature diffusing fluid in the discharge stage (Butterfield et al. 2004; von Damm and Lilley 2004; Toki et al. 2008). These key processes, however, have been hypothesized to occur in the high-temperature hydrothermal reaction zones and during the discharge stage, not in the fluid recharge stage, where cool seawater penetrates into the subsurface and its chemistry changes with increasing temperature via relevant (bio)geochemical processes.

Recently, I and my colleagues have claimed that chemical and biological processes occurring in the recharge stage significantly affect the chemical composition of the venting fluid in the case that the product is inert under the conditions of the high-temperature hydrothermal reaction zone, such as methane (Kawagucci et al. 2011, 2013a). From this point of view, we have considered Microbial Methanogenesis at Recharge stage (MMR) in hydrothermal circulation as a possible predominant source of CH4 in the high-temperature venting fluid, and refer to the concept as the ‘MMR model’ (Kawagucci et al. 2011). The MMR model is relevant particularly in a sediment-associated hydrothermal system, where the recharge stage of hydrothermal fluid circulation is expected to include an organic-rich sedimentary environment that promotes the widespread occurrence of functionally active microbial communities and impacts on fluid chemistry. Although the characteristic fluid chemistry in sediment-associated hydrothermal systems has occasionally been explained by equivocal ‘fluid-sediment interaction’, the MMR model argues that “fluid-sediment interaction” has been further classified into three types with respect to processes (microbial or chemical) and stages of subseafloor fluid circulation (recharge or discharge) (Kawagucci et al. 2013a).

In this chapter, I attempt to discover how and where the ‘TAIGA of Methane’ forms mainly based on fluid chemistry, following the recently proposed MMR model (Kawagucci et al. 2011, 2013a). For this purpose, general characteristics of the hydrothermal fluid chemistry of the OT fields are reviewed at first and then compared with those of other hydrothermal systems in Sect. 30.2. After assessing the utility of geochemical tracers to deduce the origin of hydrothermal fluid methane in Sect. 30.3, I discuss the geochemical origins of CH4 dissolved in hydrothermal fluid in each OT field and their relevance with the geological setting and geographical distribution of hydrothermal fluid circulation in Sect. 30.4.

2 Major Fluid Chemistry

2.1 General View of the OT Hydrothermal Fluids

The OT hydrothermal fluids show significant enrichment of not only CH4 but also of some other components (e.g., CO2, K, NH4, etc.) compared to hydrothermal fluids collected from sediment-starved Mid-Ocean Ridge (MOR) fields. Pioneering works (Sakai et al 1990a, 1990b; Gamo et al. 1991; You et al. 1994; Ishibashi et al. 1995) noted that the distinctive chemistry of the OT hydrothermal fluids is linked with the organic-rich continent-derived sediment filling the OT seafloor and the volatile-rich dacitic-rhyolitic (silicic) magma beneath the OT. Moreover, relatively shallow water depths of the OT hydrothermal systems (Fig. 30.1) serve as a physical factor that induces frequent boiling (subcritical phase separation) and subsequent phase segregation (Fig. 30.2), which results in chemical compositions of the venting hydrothermal fluids that are quite different from the deep source fluid (Gamo 1995). Phase separation forces volatile species into the resulting ‘vapor’ phase and leaves ion species in the ‘liquid’ phase. The source hydrothermal fluid composition prior to the phase separation can be estimated by a simple correction using the venting fluid chlorine concentration (i.e., the ion-element/Cl ratio) (Butterfield et al. 2003). Despite potential influences of multiple factors due to the geological and tectonic backgrounds of the OT, limited observations at each sediment-associated basalt-hosted MOR system and sediment-starved silicic rock-hosted arc-back-arc (ABA) system in the early 1990s inhibited the understanding of how the sediment filling the OT seafloor and volatile-rich silicic host rocks impact the OT hydrothermal fluid chemistry. Over the 35 years since the discovery of deep-sea hydrothermal activity in 1977, the current global dataset of hydrothermal fluid chemistry has enabled us to compare the fluid chemistry of the OT fields with that of hydrothermal systems of various geological and tectonic backgrounds. Field-to-field variations (inter-field variation) of fluid chemistry among the global hydrothermal fields help to clarify the characteristic fluid chemistries of the OT hydrothermal fluids and their relationships with geological and tectonic backgrounds.

Fig. 30.2
figure 2

Two-phase curve for seawater. Each symbol represents P-T conditions of venting fluids at each field from the Okinawa Trough (filled circles), sediment-associated systems (filled squares), sediment-starved ABA systems (open circles), and sediment-starved MOR systems (open squares). An open star (CP) represents a critical point of seawater. References for this figure are presented in Sect. 30.2.2

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This chapter discusses only the ‘endmember’ high-temperature fluid chemistry, estimated by a Mg-diagram and extrapolation of Mg concentrations to zero (von Damm et al. 1985), to avoid losing focus on the inter-field comparison. Thus this chapter does not discuss secondary chemical modification associated with either low-temperature hydrothermal fluid discharges through the sedimentary environment (Gamo et al. 1991; Kawagucci et al. 2013b; Yokoyama et al. Chap. 31) or the liquid CO2 emersion and the relevant microbial ecosystems (Inagaki et al. 2006; Yanagawa et al. 2013). In addition, trace elements (Mo, W, REE, etc.) measured in some OT hydrothermal fluids (Kishida et al. 2004; Hongo et al. 2007; Kawagucci et al. 2013b) are also not included in this chapter because the limited available dataset prohibits global comparison.

2.2 Inter-Field (Field-to-Field) Variation and the OT Characteristics

Figure 30.1 shows the physical properties and the Mg-corrected endmember chemistries of high-temperature hydrothermal fluids in (1) the six OT fields (Table 30.1), (2) four sediment-associated MOR fields (Lilley et al. 1993, 2003; Proskurowski et al. 2004, 2006; von Damm et al. 2005; McCollom 2008; Ishibashi et al. 2002; Pearson et al. 2005; Gieskes et al. 2002; Butterfield et al. 1994; and references therein), (3) ten sediment-starved ABA fields (Tsunogai et al. 1994, 2005; Toki et al. 2008; Kishida et al. 2004; Ishibashi et al. 1994a, 1994b; 2002b; Grimaud et al. 1991; Reeves et al. 2011; Mottl et al. 2011; de Ronde et al. 2011; and references therein), and (4) twelve sediment-starved MOR fields (Gamo et al. 2001; Kumagai et al. 2008; Gallant and von Damm 2006; Fouquet et al. 2010; Charlou et al. 2010; Von Damm et al. 1985; Merlivat et al. 1987; and references therein). The hydrothermal fields shown in Fig. 30.1 represent global variations in host rock chemistry (silicic, mafic, and ultramafic) and geographical distribution (the Indian Ocean, the Pacific Ocean, the Atlantic Ocean, and the adjacent seas). The endmember Cl concentrations are deviated from seawater Cl levels in almost all the hydrothermal fields (Fig. 30.1), demonstrating alteration in fluid chemistry due to the phase separation. For ion species, element/Cl ratios are used instead of concentrations to eliminate the effect of phase separation (Fig. 30.1).

The highest measured temperatures of hydrothermal fluids (Tmax) in the OT hydrothermal fields range between 280 and 326 °C, which are lower than typical Tmax of the MOR fields (approximately 350 °C). Fluid temperature is very likely constrained by pressure, which determines the fluid boiling temperature (Fig. 30.2). Because the water depths of the OT hydrothermal fields range between 700 and 1,600 m and are generally shallower than other hydrothermal systems (Fig. 30.1), the pressure condition (the potential of Tmax) is lower. In fact, the fluid pressure-temperature conditions of the OT fields are very close to the two-phase boundary (Bischoff and Rosenbauer 1984) (Fig. 30.2).

The hydrothermal fluid pH of the OT fields, measured at 25 °C in an onboard laboratory, range from 4.65 to 5.7. This pH range is generally comparable with that in the sediment-associated MOR systems (pH = 4.5–5.9) but higher than the pH ranges of sediment-starved ABA and MOR systems (pH = 2–5) (Fig. 30.1). In addition, a low-temperature hydrothermal fluid with an extremely low pH (≤2.1) due to direct emission of magmatic SO2 to the seafloor has been observed in the DESMOS field in the eastern Manus Basin, Western Pacific (Gamo et al. 1997). The moderately acidic pH of the OT fluids despite the ABA tectonic background suggests that some components from sedimentary organic matter decomposition, such as millimolar levels of NH3/NH4 + (Table 30.1), buffer the pH. The presence of NH3/NH4 + in high-temperature hydrothermal fluids is known to yield a high pH at the 25 °C measurement but a lower pH at in situ hydrothermal conditions (Tivey et al. 1999). The OT hydrothermal fluids contain abundant CO2 (>40 mM) of magmatic origin, based on the carbon isotope ratios of the CO2 (Ishibashi et al. 1995; Kawagucci et al. 2011, 2013a). In contrast, the concentrations of H2S, another major component of the magma-derived volatiles, are similar between the OT and other fields, regardless of the geological/tectonic backgrounds (Fig. 30.1). The helium isotope ratios (3/4He) of the OT hydrothermal fluids are ≤7.1 and are similar to those in the sediment-associated MOR fluids (Fig. 30.1). This 3/4He range, which is slightly lower than the typical sediment-starved MOR value of 8 (Sano and Fischer 2013), implies a contribution of crustal 4He-enriched helium.

The lower Na/Cl ratios of the OT fluids (0.68–0.78) compared to the seawater Na/Cl ratio (0.86) suggest Na uptake into the solid phase during subseafloor fluid circulation (Fig. 30.1). High K/Cl ratios are typical characteristics of hydrothermal fields in ABA systems, including the OT fields (>0.09) (Fig. 30.1). The K enrichment in the ABA fluids is attributed to probable K enrichment in surrounding rocks, with which fluid interacts during fluid circulation (Sakai et al. 1990a). The magnitudes of Na-loss and K-gain in each OT field seem consistent between each other (Fig. 30.1), implying quantitative exchange of Na for K during fluid-mineral interaction. In addition, a geographical trend in the K/Cl variation within the OT is also found: higher K/Cl ratios in southern OT fields (Daiyon-Yonaguni and Hatoma) than in northern OT fields (particularly Minami-Ensei). This trend may reflect a difference in the K content of the basement rocks, although there is little data on the host rock chemistry of the OT hydrothermal fields (Shinjo and Kato 2000). The Ca/Cl ratios are lower in the OT fields (<0.05) than in the others, including sediment-associated MOR fields (Fig. 30.1). The Li/Cl ratios are substantially higher in all of the OT fields (>0.002) than in the other fields. Because of the almost quantitative partitioning of alkali elements into the fluid phase during high-temperature fluid-rock (-sediment) interactions, the high Li/Cl ratios in the OT fluids suggest Li enrichment in the OT basement rocks and/or the OT-filling sediment. Low Fe/Cl ratios (<0.001) are typical in the sediment-associated fields while the low Fe/Cl ratios of the OT fluids may be associated with low Tmax. The Mn/Cl ratios are similar regardless of tectonic/geological background, excluding several ABA fields where the Mn/Cl ratios are extremely high. The Sr/Cl ratios so far observed in hydrothermal fluids vary among the fields regardless of tectonic/geological backgrounds. The hydrothermal fluid Sr/Cl range, higher or lower than the seawater Sr/Cl ratio in each field, indicates that both net gain and loss of Sr from the starting seawater during the hydrothermal fluid circulation is possible. Strontium isotope ratios (87/86Sr) are high (≥0.709) in the OT Minami-Ensei and Jade fields (Noguchi et al. 2011) and a sediment-covered Escanaba field but low (approximately 0.704) in sediment-starved MOR and ABA fields and sediment-covered Guaymas and Middle Valley fields. The high 87/86Sr ratios of the OT fluids are attributed to those in the OT-filling sedimentary component (Noguchi et al. 2011).

3 Gas Species Chemistry

3.1 Methane Sources in Hydrothermal System

The chemical origins of subseafloor CH4 are typically classified by a combination of the carbon source (inorganic or organic) and the generation process (chemical or biological, which approximately correspond to high-temperature or low-temperature environments) (Fig. 30.3). This classification distinguishes three types: so-called abiotic CH4 (e.g., McCollom 2013), thermogenic CH4 (e.g., Welhan and Lupton 1987), and biogenic CH4 (e.g., Valentine et al. 2004). Abiotic CH4 has been considered to be the predominant source of hydrothermal fluid CH4 in typical sediment-starved hydrothermal systems. However, in almost all the sediment-starved hydrothermal fields so far observed, the measured CH4 concentrations are several orders of magnitude greater (or smaller) than the CH4 concentrations that are thermodynamically predicted from the H2 and CO2 concentrations and fluid temperature of the venting fluid (McCollom 2008). While the excess could be explained by the loss of H2 and/or CO2 during fluid upwelling from the deep zone where abiotic methanogenesis has actually occurred, CH4 derived from other sources could also account for the excess. Thermogenic CH4 has been thought to be the predominant source for abundant CH4 in hydrothermal systems occurring close to sedimentary organic-rich environments, such as the OT fields (Ishibashi et al. 1995). Even if the seafloor has a sediment-starved appearance, organic matter buried beneath the seafloor is assumed to be a source for thermogenic CH4 (Lilley et al. 1993). Biogenic CH4 includes methane generated by both hydrogenotrophic and organotrophic methanogenesis. The microbial syntrophic relationship between fermentative hydrogenogenesis and hydrogenotrophic methanogenesis is expected in the sedimentary environment (Fig. 30.3). Biogenic CH4 could be incorporated into hydrothermal fluid only in the recharge stage of the whole fluid circulation process because the reaction zone and the discharge stage are too hot for any microbes to be active (Takai et al. 2008). The recharge stage environment, where the penetrating seawater becomes reduced and warm, is very likely suitable for anaerobic and/or thermophilic microbes.

Fig. 30.3
figure 3

Schematic illustrating methanogenic pathways in subseafloor geofluid systems. Ac, C2+, and OM represent acetate, hydrocarbons with carbon numbers >2, and organic matter, respectively. After Kawagucci et al. (2013a) with minor modification

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The MMR model, which we have proposed in recent papers (Kawagucci et al. 2011, 2013a), describes the origin of hydrothermal fluid CH4 in not only geochemical terms but also geographical aspects. This model claims that CH4 in high-temperature hydrothermal fluid can originate from CH4 that is generated biologically and incorporated into the fluid in the low-temperature recharge stage of hydrothermal fluid circulation. The MMR model is emphasized particularly in the sediment-associated hydrothermal systems where the recharge stage of hydrothermal fluid circulation is expected to include an organic-rich sedimentary environment. Even if the sedimentary recharge zone is absent, methanogenic ecosystems in the crustal recharge zone might be fueled by abiotic H2 generated from serpentinization (McCollom and Bach 2009), water radiolysis (Lin et al. 2005), and/or fault activity (Hirose et al. 2011). Moreover, abiotic methanogenesis associated with serpentinization of ultramafic rocks in the recharge stage cannot be ruled out because of its kinetic and thermodynamic favorability at temperatures of 200–315 °C (McCollom and Bach 2009). Consequently, all three types of CH4 generation and their occurrence through each stage of the whole fluid circulation system should be considered to deduce the origins of CH4 in venting hydrothermal fluid.

3.2 Geochemical Tracers to Deduce the Origin of Methane

The origin of hydrothermal fluid CH4 can be deduced by using geochemical tracers, and the multiple tracers approach allows us to more accurately deduce the origin of hydrothermal fluid CH4. Although several tracers have been proposed and indeed utilized, the indications of the tracers include uncertainties in some circumstances and should be assessed carefully. This issue has already been discussed in previous papers on stable hydrogen isotope ratios (δDCH4) (Kawagucci et al. 2011), stable carbon isotope ratios (δ13CCH4), relative abundances of CH4 to non-methane hydrocarbons (C1/C2+), and H2 concentrations (Kawagucci et al. 2013a). Here, a summary of the previous discussion with some additional information is presented.

3.2.1 Methane Concentration and CH4/3He Ratio

The concentration of CH4 in the venting fluid is a fundamental piece of information in deducing the origin of the CH4. The global dataset of high-temperature hydrothermal fluid compositions illustrates two approximate relationships between CH4 concentrations and geological settings. First, CH4 concentrations are significantly higher in hydrothermal fields located in sedimentary locations (Figs. 30.1 and 30.3). This type of CH4 enrichment typically accompanies enrichment in certain other components, such as ammonium, suggesting anoxic organic matter decomposition for its origin. Second, CH4 concentrations are approximately related to the redox state of the rocks exposed around the vents: higher CH4 concentrations are associated with more reduced rocks (Figs. 30.1 and 30.4). Although, again, measured CH4 concentrations are not consistent with thermodynamic predictions (McCollom 2008), more reducing conditions in the reaction zone are advantageous for abiotic methanogenesis.

Fig. 30.4
figure 4

Plots of H2 concentration versus CH4 concentration in hydrothermal fluids. Open and gray squares represent sediment-starved and sediment-associated fields, respectively. Vertical bars indicate typical H2 concentrations in reaction zone fluids for silicic, mafic, and ultra-mafic basement systems. Labels GB and MV are Guaymas Basin and Middle Valley, respectively. After Kawagucci et al. (2013a) with minor modification

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The relative abundance of CH4 with 3He (CH4/3He) differentiates mantle-derived CH4 from crust- and sediment-derived CH4 because juvenile 3He should be derived from the mantle. CH4/3He ratios higher than 107 (mol/mol) are typically found in sediment-associated hydrothermal fields (Lilley et al. 1993; Ishibashi et al. 1995) and sediment-starved slow-spreading MOR fields (Charlou et al. 2000), while ratios of 106 are reported for moderate- to fast-spreading MOR hydrothermal fields (Kawagucci et al. 2008; and references therein).

3.2.2 Carbon and Hydrogen Isotope Ratios of Methane

The δ13CCH4 value is the geochemical indicator most widely used to deduce the origins of environmental CH4. Expected ranges of δ13CCH4 values for each of the three types have been investigated in laboratory experiments (Fig. 30.5). The δ13CCH4 values of abiotic CH4 in the experiments (−40 to −20 ‰) (e.g., Horita and Berndt 1999; McCollom and Seewald 2006; McCollom et al. 2010) are quite different from those frequently observed in the sediment-starved hydrothermal fields (−20 to −5 ‰). The inconsistency in abiotic δ13CCH4 values remains unresolved. I suspect the reason may be due to some experimental artifacts, such as unnatural catalysts of Fe-Ni alloy, for example. Thermogenic CH4 generated from sedimentary organic matter (δ13Corg = −25 to −20 ‰) likely results in δ13CCH4 values between −30 and −20 ‰ because of small isotope fractionation, based on experiments (Seewald et al. 1994). However, larger fractionations at lower temperature (Hoefs 2009) and intra-molecular 13C depletion in the methyl-carbon of long-chain hydrocarbons (Gilbert et al. 2013) imply possibly more 13C-depleted thermogenic CH4. Although a broad δ13CCH4 range (−50 to −20 ‰) was proposed for thermogenic CH4 based on the compilation of field observations (Whiticar 1999), such a broad range may result from the mixing of ‘pure’ thermogenic CH4 described above with 13C-depleted biogenic CH4 (−100 to −40 ‰) (Kawagucci et al. 2013a and references therein). The δ13CCH4 values at the time of generation might be modified by certain processes occurring during subseafloor fluid migration (e.g., carbon isotope exchange between CO2-CH4), but the δ13CCH4 modifications are expected to be negligible in terms of both reaction kinetics and/or magnitudes of the fractionation (see Kawagucci et al. 2013a for details). Consequently, the δ13CCH4 value is a robust and useful indicator for deducing the origin of subseafloor CH4.

Fig. 30.5
figure 5

δ13C values of CH4 (black bars), CO2 (white boxes), and organic carbon (gray bars) considered typical for each process based on experimental and theoretical studies and observations in hydrothermal fluids. After Kawagucci et al. (2013a) with minor modification

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The δDCH4 value has been used as an indicator to deduce the origin of environmental CH4, in addition to the δ13CCH4 value (Fig. 30.6). Many incubation experiments using both microbial isolates and bulk ecosystems (e.g., soil) revealed that biogenic CH4 displayed a typically D-depleted isotope signature (−350 to −250 ‰) (Sugimoto and Wada 1995; Valentine et al. 2004) at the case that the δDH2O value is +0 ‰. These incubations also revealed that the empirical δDCH4/δDH2O ratio is approximately 0.75 for hydrogenotrophic methanogenesis and approximately 0.25 for aceticlastic methanogenesis. In contrast, past compilations of observations from sediment, oil fields, and natural gases (Nakai et al. 1974; Whiticar 1999) proposed a broader range of δDCH4 values (−450 to −150 ‰) as the “biogenic methane signature.” The proposed biogenic δDCH4 characteristics from the compilation should be followed with care because environmental CH4 may result from the mixing of multiple CH4 sources and may have been partly consumed that results in increase of δDCH4 value. Nevertheless, all hydrothermal δDCH4 values measured so far have fallen into a narrow range (−130 to −96 ‰), regardless of the types of hydrothermal systems (Proskurowski et al. 2006; Kawagucci et al. 2011, 2013a; Reeves et al. 2011). This narrow range is very likely caused by hydrogen isotopic equilibrium between CH4 and H2O in the high-temperature fluids (>250 °C and δDH2O = +0 ‰) that lead to δDCH4 values of approximately −130 ‰ (Proskurowski et al. 2006) although the certain reaction kinetics of hydrogen isotope exchange between CH4 and H2O toward isotopic equilibrium have never been constrained (Reeves et al. 2012). While no experiment has yet verified the δDCH4 value of thermogenic CH4, the δDCH4 values would be dominated by subsequent isotope exchange resulting in the isotopic equilibrium due to the obligatory high temperature. In addition, experimentally estimated δDCH4 values for abiotic CH4 (e.g., approximately −500 ‰) (McCollom et al. 2010) (Fig. 30.6) are quite different from the isotopically equilibrated δDCH4 values, despite high-temperature condition for the methanogenesis. The presence of both equilibrated δDCH4 values and non-equilibrated δ13CCH4 values in a CH4 reservoir would be reasonable because the molecular geometry of CH4 (carbon as a central atom and hydrogen as terminal atoms) suggests more rapid exchange of hydrogen. Consequently, the δDCH4 value is, in general, not useful in deducing the origin of hydrothermal fluid CH4 due to the predominance of secondary isotope exchange. However, co-variation of δ13C and δD (ΔδD/Δδ13C = 8–10) has been identified in microbial methanotrophic activity (Feisthauer et al. 2011). Therefore, combination measurements between δDCH4 and δ13C CH4 for a series of fluid samples provides supporting information for discussion of whether the δ13CCH4 values of the samples have been altered by microbial methanotrophy (Fig. 30.6).

Fig. 30.6
figure 6

Plots of δDCH4 values versus δ13CCH4 values. Open symbols represent the observations in the Okinawa Trough (ME: Minami-Ensei; Izena: Jade and Hakurei). Light gray areas represent typical ranges of sediment-starved and -associated fields as labeled. Dark gray areas represent the expected ranges for thermogenic, biogenic, and abiotic methane, where the values of δ13CDIC (−25 to +0 ‰), δ13Corg (−25 to −20 ‰), and δDH2O (+0 ‰) are assumed for a methanogenic environment (see literature for details). An arrow and a horizontal break line represent the direction of δDCH4 change on the hydrogen isotope exchange between H2O and CH4 and the value at the isotope equilibrium, respectively. Diagonal dotted lines represent typical co-variation of δ13C-δD in microbial methane consumption (ΔδD/Δδ13C = 8–10) (Feisthauer et al. 2011)

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The radioactive carbon in CH4 (Δ14CCH4) is a possible tracer. Because magma-derived carbon should be 14C-dead, detectable 14C in hydrothermal fluid CH4 indicates carbon input from seawater or young sediment. Therefore, detectable 14C in the hydrothermal fluid CH4, if present, implies microbial methanogenesis in the recharge stage of hydrothermal fluid circulation as a dominant CH4 source. As far as the author is aware, only one study, which used samples from the sediment-covered Guaymas hydrothermal field, has reported 14C measurements of high-temperature hydrothermal fluid CH4 (Pearson et al. 2005) that revealed 14C-dead CH4 in hydrothermal fluid.

3.2.3 Molecular Hydrogen and Non-methane Hydrocarbons

The concentration of H2 in hydrothermal fluid primarily depends on the redox conditions of the high-temperature hydrothermal reaction zone. In fact, a clear relationship between the host rock type (ultramafic to silicic) and the measured H2 concentration has been identified in sediment-starved fields (Fig. 30.4). The relationship is also found even in several sediment-associated fields, such as the OT Minami-Ensei and Jade fields (Fig. 30.4). Because H2 is likely buffered in the reaction zone by rapid reaction kinetics, additions of H2 in the recharge stage of hydrothermal fluid circulation would be diminished in the reaction zone and never detectable in the venting fluid. In turn, because sub-millimolar levels of H2 are expected from mafic host rocks, the measured millimolar levels of H2 in hydrothermal fluids in the fully sediment-covered mafic-hosted Guaymas and Middle Valley fields are attributed to thermogenic H2 input during the discharge stage of hydrothermal fluid circulation. Consequently, the H2 concentration of the venting fluid is a useful indicator in determining whether thermal fluid-sediment interaction occurred in the discharge stage.

The C1/C2+ ratio has been often utilized to deduce the origins of subseafloor CH4 (Fig. 30.7). Although hydrothermal experiments demonstrating abiotic CH4 generation have shown broad ranges of C1/C2+ ratios of >103 (Horita and Berndt 1999) and approximately 3 (McCollom and Seewald 2006; McCollom et al. 2010), the observed range in sediment-starved hydrothermal fields is typically >103 (McCollom 2008 and references therein). The difference may be due to experimental artifacts, such as the use of unnatural catalysts of native iron, for example. Microbial methanogenesis accompanies little C2+ hydrocarbon generation and is expected to feature C1/C2+ ratios of >104 (Kawagucci et al. 2013a; and references therein). While the C1/C2+ ratios of thermogenic hydrocarbons under high-temperature hydrothermal conditions have not yet been experimentally verified, a C1/C2+ ratio of <102 has been practically used as a thermogenic hydrocarbon signature (e.g., Whiticar 1999). It should be noted that thermal decomposition of C2+ hydrocarbons to CH4 would likely occur under deep-sea hydrothermal conditions due to both thermodynamic and kinetic aspects and could result in increasing C1/C2+ ratios as high as those of abiotic/biogenic CH4 (Kawagucci et al. 2013a). It should also be noted that a C1/C2+ ratio resulting from bimodal mixing between typical thermogenic hydrocarbons and biogenic CH4 is not linearly correlated with the mixing ratio of CH4 between these hydrocarbons (Fig. 30.7). Consequently, low C1/C2+ ratios (<103) are a definite indication of thermogenic hydrocarbons contributions to the fluid, whereas higher C1/C2+ ratios (>103) do not rule out involvement of thermogenic hydrocarbons.

Fig. 30.7
figure 7

A diagram of C1/C2+ ratios versus δ13CCH4 values. Gray areas represent the expected ranges of each hydrocarbon source, estimated by assuming δ13CDIC (−25 ‰ to +0 ‰) and δ13Corg (−25 to −20 ‰) (see Kawagucci et al. (2013a)). Abiotic methane range is not described. Five curves demonstrate bimodal mixing scenarios between biogenic and thermogenic hydrocarbons. The bold curve is considered the most likely scenario for subseafloor environments in sediment-associated hydrothermal systems at this time. Arrows with percentages present contributions of biogenic methane at the arrowed points of the bimodal mixing curves. Open symbols represent the measured values in each Okinawa Trough hydrothermal field. After Kawagucci et al. (2013a) with minor modification

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4 Estimation of Fluid Circulation Path: Implications of Methane Geochemistry

By using the multiple tracers discussed above, the geochemical origins of the hydrothermal fluid CH4 in each OT field and the relationships with the geological settings can be modeled in two-dimension schematic illustrations (Fig. 30.8). This modeling enables us to deduce the geographical distribution of hydrothermal fluid circulation (Kawagucci et al. 2013a). I attempt to discover the geographical distribution of fluid circulations for all six OT hydrothermal fields, with the addition of the OT bathymetry. It should be noted that almost all of the hydrothermal fluid vents of the OT fields are hosted in volcanic bodies that outcrop from the sediment that fills the OT.

Fig. 30.8
figure 8

Schematic illustrations of the 2D model representing methane origins based on the MMR model (after Kawagucci et al. (2011, 2013a) with minor modification). The geological setting and possible processes controlling fluid geochemistry, along with fluid circulation, are described (not to scale). White, dark-gray, and light-gray fields represent the seawater column, seafloor sediment, and rock basement, respectively. Arrows represent hydrothermal fluid migration paths. The two character designations represent the following: BR biological processes during the Recharge stage, TR thermal fluid-sediment interaction during the Recharge stage, MI magmatic volatile Inputs, RZ reaction zone, and TR thermal fluid-sediment interaction during the Discharge stage. Dashed lines are conceptual isotherms indicating the boundary between the microbial activity-dominated region and the thermochemical reaction-dominated region. Panels correspond to the models for describing (a) Iheya North and Hatoma, (b) Minami-Ensei, (c) Jade and Hakurei, and (d) Daiyon-Yonaguni

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4.1 Iheya North

The Iheya North hydrothermal fluid features low δ13CCH4 values (−58 to −51 ‰), high C1/C2+ ratios (>3,700), equilibrated δDCH4 values (−130 ‰), and low H2 concentrations (~0.2 mM) (Kawagucci et al. 2011, 2013b). The low δ13CCH4 values strongly suggest biogenic CH4 as the dominant geochemical origin. The high C1/C2+ ratios and the low H2 concentrations eliminate the possibility of thermal fluid-sediment interaction during the discharge stage. All these geochemical indications can be consistently explained by dominance of microbial methanogenesis in the recharge stage as the origin of hydrothermal fluid CH4 (Kawagucci et al. 2011) (Fig. 30.8a). Carbon mass balance calculations for the Iheya North hydrothermal system, as well as subseafloor hydrogeological structures revealed by geophysical study (Tsuji et al. 2012), suggest that the recharge stage of the Iheya North hydrothermal system includes the spatially abundant OT-filling sediments surrounding the Iheya North Knoll. This requires more than several kilometers of subseafloor hydrothermal fluid migration from the OT-filling sediment to the NBC vent (Fig. 30.9a).

Fig. 30.9
figure 9

Bathymetry maps of the six high-temperature hydrothermal fields at the same horizontal scale (coloring of depths is independent among the maps). Broken red circles on the maps represent the 2-miles radius from the main vent for each hydrothermal field

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4.2 Minami-Ensei

The Minami-Ensei hydrothermal fluid features high δ13CCH4 values (−25 ‰), high C1/C2+ ratios (>1,710), nearly equilibrated δDCH4 values (−105 ‰), and low H2 concentrations (0.04 mM) (Fig. 30.4) (Kawagucci et al. 2013a). The high δ13CCH4 values strongly suggest predominance of thermogenic CH4 as its geochemical origin while the low H2 concentrations rule out thermal fluid-sediment interaction during the discharge stage. These geochemical indications can be explained by dominance of thermal fluid-sediment interaction occurring in the recharge stage and reaction zone of hydrothermal fluid circulation as the origin of hydrothermal fluid CH4 (Kawagucci et al. 2013a) (Fig. 30.8b). Less biogenic CH4 contribution in the Minami-Ensei fluid implies a laterally shrunken hydrothermal fluid circulation that minimizes incorporation of biogenic CH4 from the OT-filling sediment. This model seems consistent with the large volcanic body of the Minami-Ensei Knoll (approximately 700 m in height and 5 km in radius) (Fig. 30.9b).

4.3 Izena Hall (Jade and Hakurei)

The Jade and Hakurei hydrothermal fluids are known to have identical fluid geochemistry, including moderate δ13C values (−32 ‰) and the nearly equilibrated δD values (−113 ‰) (Table 30.1) (Ishibashi et al. 2014). However, there are significant differences in C2H6 concentrations (C1/C2+ = 14,900 for Jade and C1/C2+ = 2,730 for Hakurei) and H2 concentrations (0.06 mM for Jade and 1.4 mM for Hakurei). The almost identical fluid geochemistry can be explained by the fact that the Jade and Hakurei fluids likely share a common fluid reservoir, which means that they share a common fluid circulation system for the recharge stage and reaction zones (Fig. 30.8c) (Ishibashi et al. 2014). This model does not conflict with the identical δ13CCH4 and δDCH4 values at the case that the hydrothermal fluid CH4 of these fields are mainly derived from the recharge stage of hydrothermal fluid circulation. The moderate δ13CCH4 value suggests concomitant biogenic and thermogenic CH4 in the recharge stage. In turn, the relatively low C1/C2+ ratios (Fig. 30.7) and high H2 concentrations of the Hakurei fluid (Fig. 30.4) suggest thermal fluid-sediment interaction during the discharge stage. Because the sedimentary settings of the discharge stages are quite different between the fields (thick sediment-covered Hakurei field and sediment-starved Jade field), it is reasonable that thermogenic hydrocarbons and molecular hydrogen entrained into the upwelling fluid occur only in the Hakurei field (Fig. 30.8c) (Kawagucci et al. 2010a; Ishibashi et al. 2014). Hydrothermal fluid reservoir underlying the southern part of the bottom and north-eastern inner wall of the Izena Hall imply that the fluid recharge occurs not only within the Izena Cauldron but also in the volcanic body of the Izena Hall and the OT-filling sediment surrounding the Izena Hall, >4 km from the vents (Fig. 30.9c).

4.4 Daiyon-Yonaguni

The Daiyon-Yonaguni hydrothermal fluids have high δ13C values (−26 ‰) and high H2 concentrations (0.8–5.5 mM), but no C1/C2+ ratios or δD values have been reported (Konno et al. 2006). The millimolar levels of H2 concentration, despite the most likely silicic rock-hosted hydrothermal system (Fig. 30.4), suggest thermal fluid-sediment interaction occurs in the discharge stage (Fig. 30.8d). This interpretation is consistent with the high δ13C values that indicate that thermogenic CH4 is the dominant CH4 source. The geochemical origin of the hydrothermal fluid CH4, particularly less biogenic CH4, implies that the outcropping knolls surrounding the fluid venting area serve as the recharge zone of hydrothermal fluid circulation: Such fluid recharge reduces incorporation of biogenic CH4 from the sediment surrounding the Daiyon-Yonaguni knoll (Fig. 30.8d). Seafloor sediment of the fluid venting area, a valley of the small knolls, hosts thermal fluid-sediment interaction during fluid upwelling, which increases H2 in the fluid (Fig. 30.8d). Assuming the above model, the lateral extent of the hydrothermal fluid circulation is <4 km in radius from the vents (Fig. 30.9d).

4.5 Hatoma

The Hatoma hydrothermal fluids feature low δ13C values (−50 ‰) (Saegusa et al. 2006) and low H2 concentrations (0.3 mM) (Kawagucci et al. 2010b). Although no C1/C2+ ratios or δD values have been reported, the geochemical origin of CH4 in the Hatoma fluid can be explained by a model similar to that of the Iheya North field: Biogenic CH4 is incorporated during the recharge stage of fluid circulation while there is little involvement of thermogenic compounds (Fig. 30.8a). Because the Hatoma Knoll is small (approximately 1 km in radius) compared with the other OT knolls hosting high-temperature hydrothermal fields (Fig. 30.9), the OT-filling sediments surrounding the Hatoma Knoll are likely the recharge stage for hydrothermal fluid circulation and the source of biogenic CH4 in the vent fluid (Fig. 30.9e).

5 Concluding Remarks

This chapter provides a compilation of high-temperature hydrothermal fluid geochemistry from the OT hydrothermal fields. The generation of the ‘TAIGA of Methane’ (in other words, how and where fluid-sediment interaction occurs within a hydrothermal system) is classified by the geochemical and geographical origins of the venting methane into three patterns: (1) microbial methanogenesis in the sedimentary environment during the recharge stage, (2) thermal degradation of sedimentary organic matter during the recharge stage, and (3) thermal degradation of sedimentary organic matter during high-temperature hydrothermal fluid upwelling through the sediment layer beneath the vent (Fig. 30.8). The discussion of the origins of CH4 provides implications for the lateral extents of hydrothermal fluid circulation in each OT field (Fig. 30.9), such as larger contribution of biogenic CH4 relating to more widespread recharge zone. However, this remains highly speculative. To strengthen the model, an understanding of subseafloor hydrogeological conditions via analysis of seismic reflection data and measurement of the permeability of cored rocks from seafloor drilling is essential. Nevertheless, the multiple geochemical tracer approach, when combined with interdisciplinary geological and microbiological investigations, will be a great breakthrough in clarifying the geochemical and geographical origins of CH4 and associated hydrocarbons not only in deep-sea hydrothermal systems but also in the other sub(sea)floor geofluid systems, including subseafloor methane hydrate.