Sources and physicochemical characteristics of fluids along a subduction-zone megathrust: A geochemical approach using syn-tectonic mineral veins in the Mugi mélange, Shimanto accretionary complex



[1] The Mugi mélange in the Shimanto accretionary complex, southwest Japan, records faulting and fluid flow patterns at the updip limit of the seismogenic region of the Nankai subduction zone. To characterize the origin and behavior of syn-tectonic fluids, we investigated the carbon, oxygen, and strontium isotopic compositions, and rare earth element (REE) patterns of syn-tectonic calcite within veins along fault zones in the mélange, as well as the Sr isotopic compositions and REE patterns of surrounding host rocks. With the exception of intra-basalt veins formed prior to subduction, the δ13C values of veins range from −10‰ to −19‰, suggesting a mixed carbon source (i.e., marine carbonate and organic matter). The vein-forming fluids have positive oxygen isotopic compositions (+2‰ to +9‰ (SMOW)) and high 87Sr/86Sr values (0.70794–0.70850), suggesting that the source was rock-buffered fluids affected both by terrigenous sediments and altered oceanic crust. The veins found in filling the fault zone associated with tectonic underplating have different REE patterns to those of the other veins, implying a difference in physicochemical processes affecting the fault zone near the subduction megathrust.

1. Introduction

[2] Subduction zones are characterized by both active rock deformation and high fluid fluxes, and ultimately by physicochemical interactions between the two. The fluids present are trapped both as pore water and within the crystal lattice of hydrous minerals in subducting sediments and oceanic crust. Most of the fluid is released early in the subduction process in response to porosity reduction and mineral dehydration associated with consolidation, diagenesis, and low-grade metamorphism [e.g., Kastner et al., 1991; Moore and Vrolijk, 1992; Saffer and Tobin, 2011], but some is carried to the forearc mantle and much deeper levels, ultimately causing mantle serpentinization and arc magmatism [e.g., Peacock, 1990]. Thus, the origin, volume, and composition of fluids in subduction zones are important aspects to study in the exploration of material cycling processes within the earth. At crustal levels, fluids are also thought to play important roles in the mechanics of faulting in both static and dynamic weakening processes during earthquake events [Hubbert and Rubey, 1959; Sibson, 1973; Mase and Smith, 1987; Sibson et al., 1988; Cox, 1995; Ujiie et al., 2007b]. This process should be essential along plate boundary thrust faults because these structural geological settings correspond to intensely deformed regions of crust with high fracture permeability.

[3] The origin and composition of fluids at surface to shallow depths of subduction zones have been investigated with results of ocean drillings and submersible dives [Moore, 1989; Kastner et al., 1991; Sample et al., 1993; Brown et al., 2001; Hensen et al., 2004; Teichert et al., 2005]. However, despite the importance of fluids in subduction zone processes, those occurring at seismogenic depths in subduction megathrusts are poorly understood, with a few exceptions [Magaritz and Taylor, 1976; Vannucchi et al., 2010; Yamaguchi et al., 2011]. In this paper, we focus on determining the source of fluids during deformation in a subduction zone, by analyzing syn-tectonic mineral veins in the Mugi mélange of the Shimanto accretionary complex of southwest Japan. The Mugi mélange is one of the most well studied on-land tectonic mélanges, of the variety that form near the traditional updip limit of seismogenic zone (temperature range of 100–150°C [Hyndman et al., 1997]) along the interface of a subducting plate [Kimura et al., 2007, 2012]. Previous studies of this mélange have constrained the P–T conditions during deformation [Matsumura et al., 2003], the nature of deformation features [Onishi and Kimura, 1995; Ikesawa et al., 2005; Kitamura et al., 2005; Hashimoto et al., 2006], clay mineralogy of basalt and fault rocks [Kameda et al., 2011a], the nature of fault rocks showing dynamic weakening mechanisms such as melt lubrication [Ujiie et al., 2007a, 2009] or thermal pressurization-induced fluidization [Ujiie et al., 2007b, 2008, 2010]. Here we provide field (outcrop) and petrographic descriptions, and the geological context of vein occurrences in the mélange, and present the results of a multidisciplinary geochemical and isotopic study of vein carbonates, including determination of carbon and oxygen isotope compositions, Sr isotope compositions, and rare earth element (REE) compositions. On the basis of this new data set for the Mugi mélange, we discuss the origin and evolution of syn-tectonic fluids occurring along a seismogenic plate boundary.

2. Geological Setting

[4] The Mugi mélange is part of the Late Cretaceous to Early Tertiary Shimanto accretionary complex in eastern Shikoku, Japan [Onishi and Kimura, 1995; Ikesawa et al., 2005; Kitamura et al., 2005] (Figure 1a). The mélange consists of a sheared black shale matrix containing sandstone blocks and slabs of basaltic rocks with N-MORB compositions [Kiminami et al., 1992]. Kinematic analyses suggest that the mélange was deformed during underthrusting—an idea that is consistent with relative plate motions in the region during the Late Cretaceous to Early Tertiary [Onishi and Kimura, 1995]. The imbrication of tectonostratigraphic units (comprising terrigenous sedimentary and basaltic rocks) along thrust faults is interpreted to reflect duplex underplating (Figures 1b, 1c, and 2b) [Ikesawa et al., 2005; Shibata et al., 2008]. Deformation associated with thrusting is characterized by cataclastic deformation of basaltic rocks, which followed earlier subduction-related pervasive pressure solution and the development of boudinage structures in terrigenous sandstones.

Figure 1.

Geological setting of the study area. (a) Distribution of the Shimanto accretionary complex in southwest Japan and location of the Mugi mélange. (b) Cross-section through the Mugi mélange and surrounding area. (c) Geological map of the Mugi mélange (modified from Shibata et al. [2008] and Kimura et al. [2012]). The upper and lower sections are bounded by the Mizoochi Fault, which is deduced from the thermal inversion associated with reverse faulting. The ghost ocean floor stratigraphy is structurally repeated by thrust faults, possibly representing a duplex structure.

Figure 2.

Schematic diagrams showing the tectonic setting of the Mugi mélange and the varied geological occurrences of syn-tectonic veins. (a) Schematic profile of the Nankai Trough, as derived from a seismic profile (modified from Park et al. [2002]). (b) Interpretation of the tectonic setting of the Mugi mélange (modified from Ikesawa et al. [2005]). (c) Schematic block diagram showing the geological occurrences of the four types of veins identified in this study, and their relationships to the mélange and ramp thrust of the duplex structure. Intra-basalt veins are widely developed in the basaltic rocks. Boudin-neck veins are observed at the neck regions of pinch-and-swell and boudinage structures in the mélange. Network veins are developed in damage zones along the ramp thrust, and fault-fill veins occur along dilational jogs in the thrust.

[5] The studied section (see Figure 1c), which contains the most well-exposed thrust sheet in the Mugi mélange, can be traced for ∼600 m along strike (ENE–WSW). In this area, the studied rocks dip steeply to the north. Previous analysis of contemporaneous water and methane fluid inclusions, combined with vitrinite reflectance data, reveals that these rocks were underplated at depths of 4–6 km and at temperatures of 130–150°C [Matsumura et al., 2003; Ikesawa et al., 2005], which corresponds to the traditional updip limit of the seismogenic zone in the modern Nankai Trough (Figure 2a) [Park et al., 2002; Kimura et al., 2007, 2012]. The temperatures of 130–150°C also coincide with regional trend of burial temperature in the Northern Shimanto Belt [Ohmori et al., 1997]. The age of underthrusting and underplating is poorly constrained, but U–Pb age data on detrital zircons (61–62 Ma [Shibata et al, 2008]) provide a reliable maximum age estimate for the timing of this deformation, which may have taken place relatively soon after 61 Ma.

3. Structure of the Fault Zone

[6] The Mugi mélange contains cataclastic fault zones measuring ∼20 m in thickness. These highly deformed zones are located at the bottom of each tectonostratigraphic unit (characterized by ocean floor stratigraphy). Such thrust faults cutting the mélange fabrics are commonly observed within the mélanges in the Shimanto accretionary complex, and are considered to represent ramp thrusts within a duplex structure related to underplating [Kimura and Mukai, 1991, Hashimoto and Kimura, 1999; Hashimoto et al., 2012]. In one of these fault zones located at the bottom of one particular tectonostratigraphic unit (Unit 2 of Ikesawa et al. [2005]), the distribution of fault rock (Figure 3a) indicates that localized slip took place at the upper structural levels of the fault zone [Ujiie et al., 2007b], where thin layers of ultracataclasite (with thicknesses up to several centimeters) are developed along the lithological boundary between shale-rich mélange and basaltic rocks. These ultracataclasite-bearing slip zones record evidence of the fluidization and injection of comminuted material, dilational brecciation in extensional jogs, the formation of an amorphous silica layer along the ultracataclasite [Ujiie et al., 2007b], and the stretching of fluid inclusions in calcite veins due to frictional heating [Ujiie et al., 2008]. Collectively, these features are interpreted to indicate thermal pressurization of ultracataclasite associated with seismic slip [Ujiie et al., 2010]. In addition, the existence of numerous mineral veins and intense alteration of basalt along the ultracataclasite-bearing zones suggests that intense water–rock interaction took place in connection with faulting (Figure 3a).

Figure 3.

(a) Tectonostratigraphic section of the fault zone, showing the distribution of protolith, deformation, and alteration patterns (modified from Ujiie et al. [2007b] and Kimura et al. [2012]). (b) Schematic showing the crosscutting relationships between ultracataclasite, fault-fill veins, and network veins.

4. Syn-tectonic Mineral Veins in the Mugi Mélange

[7] The Mugi mélange contains many syn-tectonic mineral veins suggestive of intense circulation of fluids during deformation. The veins are classified into four types based on their location (geological setting), crosscutting relationships, and mode of occurrence: intra-basalt, boudin-neck, network, and fault-fill veins (Figures 2c and 4).

Figure 4.

Occurrences of syn-tectonic veins in the Mugi mélange. Cal: calcite; Qz: quartz; Lau: laumontite. (a) Occurrence of intra-basalt veins filling cavities in pillow basalt. (b) Occurrence of intra-basalt veins filling extension cracks in hydrothermal cherts within pillow breccia. (c) Occurrence of boudin-neck veins in the field. (d) Photomicrograph of boudin-neck veins. (e) Occurrence of network veins in the field, most of which are truncated by basaltic ultracataclasite. (f) Photomicrograph of a network vein. (g) Occurrence of fault-fill veins in the field. (h) Photomicrograph of a fault-fill vein.

4.1. Intra-basalt Veins

[8] Intra-basalt veins are widely developed within basaltic rocks of the study area (Figure 2c), and can be found infilling cavities and cooling cracks within pillow basalts, pillow breccias, massive basalts, and hyaloclastites (Figure 4a), and also within blocks of hydrothermal chert (Figure 4b). These occurrences suggest that the veins formed during the normal movement of aging oceanic plate; i.e., after eruption of basaltic rocks at the mid-ocean ridge, and before entering the subduction zone. The mineralogy of these intra-basalt veins typically includes calcite, laumontite, quartz, chlorite, prehnite, and pumpellyite.

4.2. Boudin-Neck Veins

[9] Boudin-neck veins, distributed widely in the mélange, are observed to infill extensional cracks developed in the necked regions of pinch-and-swell and boudinage structures within sandstone (Figures 2c and 4c). Crosscutting relationships between veins, boudinaged sandstones, and surrounding mudstones indicate that this variety of veins developed synchronously with boudin formation, which took place during mélange formation [Hashimoto et al., 2006]. Quartz and calcite are the main mineralogical components of these veins (Figure 4d). These vein crystals show syntaxial growth, suggesting that the rate of crack opening was slow enough to balance the rate of mineral precipitation [Oliver and Bons, 2001].

4.3. Network Veins

[10] In contrast to boudin-neck veins, network veins and fault-fill veins are limitedly observed along the cataclastic fault zone at the bottom of each thrust sheet. Network veins are observed to infill extensional cracks that occur within wall rocks on both sides of the fault zone (Figures 2c and 4e). These veins post-date mélange formation because they are observed to crosscut the mélange fabric (Figure 4e), and interpreted that vein formation was related with underplating posterior to mélange formation [Kimura et al., 2012; Hashimoto et al., 2012]. Most of these network veins are truncated by fault rocks, although in some cases the opposite crosscutting relationship is observed. These relationships suggest repeated episodes of faulting and vein precipitation. The network veins range in thickness from tens of micrometers to several centimeters, and the vein-forming minerals consist of calcite, quartz, and laumontite. Blocky and dendritic textures (Figure 4f) are observed in thin sections observed by transmitted light microscopy, suggesting that network veins precipitated rapidly from an oversaturated fluid [Oliver and Bons, 2001].

4.4. Fault-Fill Veins

[11] Fault-fill veins occur within shear cracks in the fault zone (Figures 2c and 4g), and are restricted to only the gently dipping part of the zone. These veins include brecciated ultracataclasite (Figure 4h), suggesting rapid precipitation immediately after faulting. From these occurrences, we infer that the veins formed during an implosion process of dilational jogs within a thrust system (implosion breccias [Sibson, 1986]). Fragments of these fault-fill veins are also included in the ultracataclasite (Figures 3b and 4g), which suggests repeated episodes of faulting and vein precipitation as with the network veins. In terms of mineralogy, fault-fill veins consist of coarse-grained blocky calcite. Lack of laumontite in comparison with network veins suggests that CO2 pressure was relatively high during the precipitation of fault-fill veins [Thompson, 1971].

4.5. Order of Vein Formation and Vein Precipitation Temperature

[12] Figures 2c and 3b show the crosscutting relationships of the observed veins and faults. Both intra-basalt and boudin-neck veins are inferred to have formed prior to development of the fault zone, because they are crosscut by the fault-related deformation structures. Network veins, fault rocks (ultracataclasite), and fault-fill veins are mutually crosscutting (Figure 3b). Several previous studies reported similar sets of crosscutting extensional and shear cracks, interpreted to have formed in response to fluctuations in fluid pressure during seismic cycles (i.e., fault-valve behavior [e.g., Sibson et al., 1988; Cox, 1995]). Such a model is also consistent with the geometrical and crosscutting relationships observed between network and fault-fill veins in this study. The ultracataclasite contains injection structures associated with fluidization of comminuted material (Figure 4g), and fluid inclusions in calcite fragments were stretched by frictional heating [Ujiie et al., 2007b, 2008], which suggests that the formation of these fault-related veins was closely linked with earthquake faulting.

[13] Matsumura et al. [2003] analyzed coexisting H2O-CH4 inclusions in quartz veins, and estimated fluid trapping temperatures and pressures as 125–192°C and 92–144 MPa for boudin-neck vein (their vein I), and 135–245°C and 107–149 MPa for network vein (their vein II), respectively. Considerable uncertainty remains in their result because the estimation is based on limited number of measurement (total n = 77 for heating and n = 41 for cooling experiment), so here we attempt to evaluate more accurate temperature estimation based on additional data sets.

[14] The crosscutting relationships observed between mélange textures and veins indicates that the boudin-neck veins were precipitated at a maximum burial depth corresponding to temperatures of approximately 130–150°C, as constrained by vitrinite reflectance data (Ro = ∼1.0–1.5 [Ikesawa et al., 2005]). For the fault-related (network and fault-fill) veins, another previous study on fluid inclusion geothermometry was performed by Ujiie et al. [2008, 2010]. They focused inclusions in calcite as weak host mineral [Mullis, 1987], and found that the fluid inclusions within the fault-fill veins were thermally overprinted and re-equilibrated by frictional heating during seismic faulting. However, their abundant data (total n = 602) and the patterns observed on histograms of homogenization temperatures also suggest that the effect of the high-temperature overprint was limited, and that most of the Th values of the fluid inclusions appear to represent the original homogenization temperature (Figure 5). Although fluid inclusions within calcite are water-rich [Ujiie et al., 2008], coexisting quartz veins contains methane-rich inclusions [Matsumura et al., 2003]. This result suggests that the H2O-CH4 immiscible fluids might have coexisted at the time of vein precipitation. In the case of immiscible fluid, homogenization temperature is directly regarded as fluid trap temperature [Alderton and Bevins, 1996; Lewis et al., 2000; Hashimoto et al., 2002, 2003; Matsumura et al., 2003]. Assuming that the modes observed on frequency histograms of homogenization temperatures represent the original homogenization temperatures and fluids were H2O-CH4 immiscible fluid, the temperatures of precipitation of the network and fault-fill veins are therefore estimated to be ∼150–160°C and ∼160–170°C, respectively. These estimated temperatures are within the range of Matsumura et al.'s [2003] result.

Figure 5.

Histograms of the homogenization temperatures (Th) of fluid inclusions in calcite from the Mugi mélange (modified from Ujiie et al. [2010]). (a) Network veins (corresponding to “veins in the mélange” in Ujiie et al. [2010]). (b) Fault-fill veins (corresponding to “fragments in the ultracataclasite” in Ujiie et al. [2010]). Although the plots for fault-fill veins are characterized by frequency distributions with a wide range of Th, suggesting stretched fluid inclusions, the mode of the histograms (160–170°C) suggest that these inclusions retain the original Th.

5. Analysis of Carbon, Oxygen, and Strontium Isotopes, and Rare Earth Elements

5.1. Analytical Methods

5.1.1. Carbon and Oxygen Isotopes

[15] For analysis of carbon and oxygen isotopes, we collected 8 samples of intra-basalt veins, 27 samples of boudin-neck veins, 20 samples of network veins, and 19 samples of fault-fill veins within a 400 × 200 m area of the studied section. Samples of vein powder (10 mg of carbonate) were obtained by microdrilling from rock chips. CO2 was obtained by dissolving mineral powders in 100% phosphoric acid at 25°C overnight, following McCrea's techniques [McCrea, 1950]. The liberated CO2 gas from each sample was subsequently analyzed for carbon and oxygen isotopic ratios on a Finnigan MAT252 mass spectrometer at the University of Tokyo, Japan. Carbon and oxygen isotopic values are reported relative to the PeeDee Belemnite (PDB) and standard mean ocean water (SMOW), respectively, using the standard notation of ‰ (permil). The suitability of the standardization procedure employed for the analysis of oxygen and carbon isotopes in this study was verified by analyzing isotopic ratios of various laboratory and international carbonate standards, including NBS-19 (calcite). Uncertainties (expressed as 1σ) are ∼0.04‰ and ∼0.05‰ for δ13C and δ18O, respectively.

5.1.2. Strontium Isotopes and Rare Earth Elements

[16] For Sr isotope and REE analyses, nine vein samples (three samples each of boudin-neck veins, network veins, and fault-fill veins) and five host rock samples (three shale matrix of mélange and two altered basalts) were collected from the same region from where samples were collected for analysis of δ13C and δ18O in calcite. Sr isotope and REE analyses were performed at the Earthquake Research Institute, University of Tokyo.

[17] About 30 mg of each powdered vein sample was dissolved by HNO3 digestion in a tightly sealed 15- or 30-ml Teflon PFA screw-cap vial, while host rock samples were dissolved by HNO3–HF–HClO4 digestion. The remaining undissolved residue was then separated by centrifugation. Two thirds of each sample was used for Sr isotope analyses, and Sr was purified using Sr resin (Eichrom Technologies LLC, USA). Sr isotope ratios were determined using a multicollector–inductively coupled plasma–mass spectrometer (IsoProbe; GV Instruments Ltd., UK). Instrumental mass fractionation was corrected following an exponential law based on 86Sr/88Sr = 0.1194. The 87Sr/86Sr data were normalized to 87Sr/86Sr = 0.710258 for SRM987. The 87Sr/86Sr values obtained for the analyzed standards agree well with previously published data [Nishio et al., 2004].

[18] The concentrations of trace elements were determined without further chemical treatment using the remaining aliquots of sample solutions, and were analyzed on an inductively coupled plasma–mass spectrometer equipped with a quadrupole mass analyzer (PQ3; Thermo Elemental Winsford, UK). Overall, the samples of network veins contained very low abundances of REEs. For analysis of network veins, 10 mg of each sample was dissolved in HNO3. After centrifugation, Ca and Ba were separated using 2 M HCl and 2 M HNO3, respectively, with a cation exchange resin (AG50W-X8; Bio-Rad Laboratories Inc., USA), and the REEs were recovered by 6 M HCl prior to analysis. The precision of measured REE concentrations in rock samples is ∼10% of the quoted abundances (Table 2) [Fukuda et al., 2008].

5.2. Results

5.2.1. Carbon and Oxygen Isotopes

[19] The results of isotopic analyses of vein calcite are summarized in Table 1 and Figure 6. δ13C values of intra-basalt veins range from −0.3‰ to +1.6‰ (mean = +0.4‰), and δ18O values range from +14.6‰ to +18.5‰ (mean = +16.8‰). Boudin-neck veins yield δ13C values of −15.9‰ to −9.1‰ (mean = −13.1‰), and δ18O values of +16.0‰ to +18.0‰ (mean = +16.6‰). Network veins have δ13C values ranging from −17.2‰ to −7.4‰ (mean = −15.2‰), and δ18O values of +17.3‰ to +18.7‰ (mean = +17.6‰). The δ13C values of fault-fill veins range from −13.7‰ to −8.4‰ (mean = −11.6‰), and δ18O values range from +18.0‰ to +21.4‰ (mean = +19.2‰). Intra-basalt veins have δ13C values close to 0‰, while other veins have lower δ13C values. δ18O values of veins are lower than those of marine carbonates represented by PDB (δ18O(smow) = 1.03091 δ18O(PDB) + 30.91 [Coplen et al., 1983]), suggesting that little isotope fractionation (i.e., isotope exchange at high temperature) occurred during precipitation of veins.

Table 1. Carbon and Oxygen Isotopic Compositions of Intra-basalt, Boudin-Neck, Network, and Fault-Fill Veinsa
Sample NameStandard Deviation (δ13C)Standard Deviation (δ18O)δ13Cvein (PDB)δ18Ovein (SMOW)
  • a

    δ13Cvein (PDB): carbon isotopic composition of vein calcite; δ18Ovein (SMOW): oxygen isotopic composition of vein calcite.

Intra-basalt Veins
mean  0.4416.85
Boudin-neck Veins
mean  −13.1116.61
Network Veins
mean  −15.1717.61
Fault-Fill Veins
mean  −11.5819.23
Figure 6.

Carbon and oxygen isotopic values of vein calcite in the Mugi mélange. Measurement error is smaller than the symbol size.

5.2.2. Strontium Isotopes and Rare Earth Elements

[20] Table 2 lists the results of trace element and Sr isotope analysis of veins. 87Sr/86Sr values of vein calcite range from 0.70794 to 0.70844 for boudin-neck veins, 0.70816 to 0.70848 for network veins, and 0.70847 to 0.70850 for fault-fill veins. All of these values range between those of host rocks (shale and basalt) in the Mugi mélange (see below). Chondrite-normalized REE patterns are shown in Figure 7, and the values of (La/Yb)CN (i.e., chondrite-normalized La/Yb ratios indicating the slope of REE patterns) are also listed in Table 2. Boudin-neck veins and network veins have similar REE slopes ((La/Yb)CN values of 34–66 for boudin-neck veins, and 5.1–17 for network veins), while fault-fill veins yield relatively flat REE patterns ((La/Yb)CN values of 2.7–4.5) compared with the other vein types.

Table 2. Trace Element Concentrations, 87Sr/86Sr Ratios, and Chondrite-Normalized La/Yb Ratios of Boudin-Neck, Network, and Fault-Fill Veinsa
(ppm)Boudin-Neck VeinsNetwork VeinsFault-Fill Veins
  • a

    The bolded values show the results from Ca- and Ba-separated samples. N.D.: not detected.

85Rb0.593. D.0.48
133Cs0.040.400. D.0.063
159 Tb0.440.570.530.0610.0290.100.460.590.34
181Ta0.00300.0062N. D.0.00650.0190.00540.000730.00470.0033
Figure 7.

Chondrite-normalized REE patterns of boudin-neck, network, and fault-fill veins.

[21] The results of trace element and Sr isotope analyses of host rocks are listed in Table 3. 87Sr/86Sr values of host rocks range from 0.71128 to 0.71658 for shale, and from 0.70657 to 0.70682 for basalt. (La/Yb)CN values are 4.2–5.2 for shale and 0.55–0.56 for basalt. Chondrite-normalized REE patterns of host rocks are shown in Figure 8. REE patterns of shale show trends similar to those of typical terrigenous sediments with a negative Eu anomaly [McLennan, 1989; Plank and Langmuir, 1998]. The REE patterns of basalt host rocks are similar to those of typical ocean floor basalt, characterized by depletion of light REEs [Patchett, 1989].

Table 3. Trace Element Concentrations, and 87Sr/86Sr and Chondrite-Normalized La/Yb Ratios of Host Rocks
(ppm)Shale Matrix of MélangeAltered Basalt
Figure 8.

Chondrite-normalized REE patterns of host rocks.

6. Discussion

6.1. Fluid Source

[22] Intra-basalt veins have δ13C values close to 0‰, while other veins have lower δ13C values. These results might reflect a difference in fluid sources between intra-basalt veins and other vein types. The δ13C values determined for intra-basalt veins are within the range of marine carbonates, and of carbonate veins in modern oceanic crust and ophiolites described by Alt and Teagle [2003] and Miller et al. [2001]. This similarity suggests that intra-basalt veins were precipitated from seawater in a hydrothermal system within the oceanic crust, which is also consistent with the occurrence (geological context) of these veins. In contrast, the three other vein types have relatively light carbon isotopic compositions ranging from −17‰ to −7‰, possibly reflecting a mixed carbon source, comprising contributions from organic carbon with 13C depletion (δ13C = −25‰) and sedimentary carbonate carbon (δ13C = 0‰). These values are consistent with those of terrigenous input sediments in modern convergent plate margins (e.g., Central American margin [Li and Bebout, 2005]) as well as those of ancient subduction complexes (e.g., Franciscan Complex [Magaritz and Taylor, 1976; Sadofsky and Bebout, 2001]).

[23] The oxygen isotopic composition of the source fluids of the veins (δ18Ofluid) was calculated based on the isotopic composition of vein calcite (δ18Ovein) and the formation temperature of each vein, assuming isotopic equilibrium. Oxygen isotopic fractionation between coexisting calcite and H2O is expressed as follows [O'Neil et al., 1969]:

display math

where T is temperature and α is the isotope fractionation factor. Here, the formation temperature of each vein type was estimated to be 130–150°C for boudin-neck veins, 150–160°C for network veins, and 160–170°C for fault-fill veins (section 4.5 and Table 4). Using equation (1), the oxygen isotopic compositions of vein-forming fluids were estimated to be +1.8‰ to +4.4‰ for boudin-neck veins, +4.4‰ to +6.1‰ for network veins, and +6.6‰ to +8.7‰ for fault-fill veins (Table 4).

Table 4. Summary of Paleotemperatures and Isotopic and Geochemical Features of Syn-tectonic Fluids in the Mugi Mélangea
 Intra-basalt VeinsBoudin-Neck VeinsNetwork VeinsFault-Fill Veins
  • a

    Mean δ13Cvein (PDB): mean value of the carbon isotopic composition of vein calcite; mean δ18Ovein (SMOW): mean value of the oxygen isotopic composition of vein calcite; temperature: vein-forming temperature, as derived from vitrinite reflectance analyses and fluid inclusion geothermometry [Ikesawa et al., 2005; Ujiie et al., 2008, 2010]; δ18Ofluid (SMOW): minimum and maximum values of the oxygen isotopic composition of source fluid, as deduced from δ18Ovein values and vein-forming temperatures, assuming isotopic equilibrium; 87Sr/86Sr: Sr isotope ratios of vein calcite; (La/Yb)CN: Chondrite-normalized La/Yb ratios, representing the slopes of chondrite-normalized REE patterns.

mean δ13Cvein (PDB)0.4−13.1−15.2−11.6
mean δ18Ovein (SMOW)16.816.617.619.2
δ18Ofluid (SMOW)1.8–4.44.4–6.16.6–8.7

[24] The δ18O values of the vein-forming fluids (δ18Ofluid) are obviously higher than those of seawater (∼0‰) and meteoric water (<0‰ [Sheppard, 1986]), demonstrating that the boudin-neck, network, and fault-fill veins were strongly affected by rock-buffered fluids (Table 5 shows representative δ18O values of shales and altered oceanic crust). Dehydration reaction of clay minerals such as smectite-illite transition [Vrolijk, 1990; Bekins et al., 1994; Moore and Saffer, 2001] and/or saponite-chlorite transition [Kameda et al., 2011b] would be the candidates to control fluid composition in metasediments or metabasalts at these temperatures.

Table 5. Representative δ13C, δ18O, and 87Sr/86Sr Values of Possible End-Member Source Fluids
 60 Ma SeawaterAltered BasaltShale Matrix of Mélange
δ18O (SMOW)∼0+10.7–+12.7b+14.0–+19.0c

[25] We also observe that δ18Ofluid increases with both increasing temperature and vein type, showing a systematic increase from boudin-neck, to network, and finally to fault-fill veins. Relatively higher δ18O values of fault-fill-vein fluid would also suggest fluid-rock buffering at temperatures higher than inferred for vein formation, perhaps suggesting transport along the fault zone. This is consistent with the presence of 18O-enriched carbonate at the surfaces of modern accretionary prisms (Cascadia [Sample et al., 1993]; Nankai [Kawamura et al., 2009]) possibly reflecting channelized fluid migration from deep portions of the accretionary prisms.

[26] There is still considerable uncertainty in the estimation of temperatures of formation of both network and fault-fill veins because the fluid inclusions contained within them have been influenced and re-equilibrated by later frictional heating along the fault zone (Figure 5). However, if the formation temperature of these veins were higher than the vein formation temperatures estimated above, the corresponding δ18Ofluid values would probably be significantly larger, because isotopic fractionation becomes smaller at higher temperatures. Therefore, strongly rock-buffered fluids (δ18O ≫ 0‰) are considered to represent the likely geological sources of the vein-mineralizing fluids.

[27] The Sr isotope ratios of two representative end-member host rock compositions, including the shale matrix of the mélange (87Sr/86Sr = 0.7113−0.7166) and altered basalts (0.7066–0.7068), are consistent with previously reported typical isotopic compositions of terrigenous sediments and altered oceanic crust entering subduction zones (Table 5) [e.g., Plank and Langmuir, 1998; Kawahata et al., 2001]. The Sr isotope ratios obtained from boudin-neck, network, and fault-fill veins (87Sr/86Sr = 0.7079–0.7085) are intermediate between those of the two aforementioned end-member compositions determined on surrounding host rocks, and are slightly higher than the estimated Sr isotopic composition of 60 Ma seawater (87Sr/86Sr = 0.7078 [McArthur et al., 2001]). Sr isotopes do not generally fractionate during fluid volatilization or mineral precipitation. Assuming that rock-buffered fluids have similar Sr isotope ratios to surrounding rocks, the origin of the vein-forming fluids in this study can be explained by the mixing of 60 Ma seawater and rock-buffered fluid sources (i.e., a mixture of terrigenous sediments and oceanic crust).

[28] These relationships between isotopic systems, vein types, host rock compositions, and fluid sources are all displayed in Figure 9, which highlights a 87Sr/86Sr–δ18O plot of data for three types of vein-forming fluids and three end-member (source) compositions. These 87Sr/86Sr values, together with δ18O data, suggest that boudin-neck veins formed from fluids with relatively low 18O and low 87Sr/86Sr values, whereas the fault-fill veins formed from fluids enriched in 18O and high 87Sr/86Sr values (Figure 8). From Figure 9, the aforementioned rock-buffered fluid source composition is estimated to originate from approximately 20–60% input from terrigenous sediment and 40–80% input from oceanic crust.

Figure 9.

Sr and oxygen isotopic compositions for three types of vein-forming fluids and three end-member source compositions (shown in Table 5).

6.2. Physicochemical Features of Fluids

[29] As mentioned in section 6.1., all of the δ18O and 87Sr/86Sr data for the three types of vein-forming fluids lie along a mixing line between 60 Ma seawater and a rock-buffered end-member, suggesting that all of the source fluids for these veins were affected by the compositions of surrounding host rocks and were chemically altered by both terrigenous sediments and oceanic crust. The REE patterns of veins, however, indicate a different paragenesis than the δ18O and 87Sr/86Sr data imply. Specifically, the chondrite-normalized REE patterns obtained for fault-fill veins tend to show relatively flat patterns compared with the patterns obtained for the other two vein types, which yield LREE-enriched patterns.

[30] Therefore, the REE patterns of veins cannot simply be explained by mixing of sediment- and basalt-derived fluids. This discrepancy suggests that the REE patterns of veins reflect not only the source fluids of the veins, but also the physicochemical processes of fluid evolution. According to reviews on solution chemistry of REEs, the REE pattern of a fluid phase is controlled by both sorption and chemical complexation reactions [Brookins, 1989; Bau, 1991]. The composition of LREE-enriched ((La/Yb)CN > 1) fluid is controlled mainly by sorption processes operating under mildly acidic conditions, while the complexation reactions occurring with carbonate, fluoride, and hydroxide complexes lead to the formation of HREE-enriched ((La/Yb)CN < 1) fluid (Figure 10a). These complexation reactions are also enhanced under nearly neutral to mildly basic conditions.

Figure 10.

(a) Schematic illustration showing factors controlling the trend of REEs in vein-forming fluids. (b) REE patterns of appropriate rock-buffered end-member fluid compositions: ranges between shale = 60%, altered basalt = 40% (upper line); and shale = 20%, altered basalt = 80% (lower line). (c) REE patterns of three vein types.

[31] At this point, if we assume a mixing ratio of 20% shale and 80% altered basalt, or 60% shale and 40% altered basalt, the resultant rock-buffered end-member of the vein-forming fluids would have flat REE patterns with (La/Yb)CN = 0.8–1.8 (Figure 10b). The REE patterns of the boudin-neck and network veins are characterized by LREE-enriched patterns of (La/Yb)CN = 60–88 and 4.9–18, respectively (Figures 7 and 10c). Therefore, a sorption process is apparently necessary to produce the LREE-enriched patterns of boudin-neck and fault-fill veins. We cannot explain why this sorption process was enhanced in this fluid-rich structural geological system. We can postulate, however, that the pH values of boudin-neck and network vein source fluids were probably quite low, and that some other kind of water–rock interaction enhanced the sorption processes affecting the LREEs.

[32] Fault-fill veins, on the other hand, have relatively shallow REE profiles with (La/Yb)CN = 2.7–4.1, possibly because the sorption processes affecting the REEs was inhibited in the geological history of the source fluids for the fault-fill veins. The other possibility is that the fault-fill vein-forming fluid was affected by sorption processes, as with the other two vein-forming fluids, but in this case (i.e., for fault-fill veins) additional complexation processes occurred during the fluid evolution, selectively overprinting the REE composition of the fluid. In the latter possible scenario case, fluid pH is estimated to have been elevated.

[33] Of the two aforementioned possible explanations for shallowly sloping REE profiles, the mineralogy and structural features of the fault-fill veins seem to support the latter possibility (involving complexation). Given that the fault-fill veins are composed of calcite, the marked lack of quartz and laumontite in the veins is consistent with a fluid evolution involving enhanced complexation processes operating at neutral to mildly basic conditions. A basic pH during the complexation reactions affecting these fault-fill vein fluids is consistent with the tectonics literature. For instance, Kameda et al. [2003] and Saito and Tanaka [2007] performed crushing experiments of granite samples under fluid-saturated conditions using a ball mill, and reported an increase in fluid pH after the experiments. This pH increase may originate from the breakdown of feldspar, which causes the release of cations such as K+ or Na+ into solution, and which in turn causes pH to rise.

[34] Yamaguchi et al. [2011] reported positive Eu anomalies in fault-fill ankerite veins within the Nobeoka thrust in the Shimanto accretionary complex, and discussed the possibility of an origin by fluid reduction during faulting. In the present study, however, we observe no Eu anomalies within fault-fill veins of the Mugi mélange. This difference in physicochemical features of vein-forming fluids between the two thrust settings could reflect varying and distinct patterns of water–rock interaction in subduction zones, controlled by contrasting temperature–pressure conditions, tectonic settings (e.g., megasplay fault versus plate boundary décollement), host rock compositions, slip behaviors, and so on.

6.3. Fluid Evolution Along Subduction Plate Boundaries

[35] As discussed in section 4.5., precipitation temperatures for the boudin-neck, network, and fault-fill veins are estimated at ∼130–140, ∼150–160, and ∼160–170°C, respectively. The progressive enrichment in 18O and 87Sr with vein type (in the order of boudin-neck, to network, to fault-fill veins) is interpreted to reflect the dominant influence of a rock-buffered fluid component during the deeper stages of underplating, relative to the shallower stages of underthrusting (Figure 11) and also in comparison with the seawater fluid source component.

Figure 11.

Schematic diagram showing fluid sources and physicochemical features constrained from syn-tectonic veins in the Mugi mélange.

[36] As shown in Figure 9, the oxygen and Sr isotopic results from rocks in this study suggest that not only terrigenous sediments but altered oceanic crust contributed considerably to source fluid compositions of the vein-forming fluids during deformation. This finding is consistent with the tectonics literature on the Mugi mélange [Kimura et al., 2012] and recent discussions on the sources of fluids at seismogenic plate boundaries [Kameda et al., 2011b]. Although deformation exhibited by block-in-matrix textures of terrigenous sediments occurs widely throughout the Mugi mélange, duplexed thrust sheets identifiable by tectonostratigraphic repetitions of ocean floor stratigraphy are also recognized in the mélange [Kimura et al., 2012]. Basaltic rocks are located at the bottom of each thrust sheet, indicating that the composition of vein-forming fluids in the Mugi mélange was influenced by basalts. Recently, Kameda et al. [2011a, 2011b] investigated the clay mineralogy of basaltic rocks and pointed out that the dehydration which takes place during the saponite–chlorite transition within altered oceanic crust, should contribute significantly to fluid budgets in the seismogenic zone. In evaluating the origin of fluids in the seismogenic zone, the role of terrigenous sediment, especially the smectite–illite transition, has already been investigated to great extent [e.g., Vrolijk, 1990; Moore and Saffer, 2001; Vannucchi et al., 2010] and hydrological investigations have been performed [e.g., Bekins et al., 1994; Saffer et al., 2008; Saffer and Tobin, 2011], but the possible role of oceanic crust in controlling fluid compositions in the seismogenic zone has yet to be evaluated. The results presented in this study suggest that both of these components contributed to varying extents in controlling the composition of vein-forming fluids, although further geochemical studies are needed to better constrain the nature of fluid sources in seismogenic zones at subduction plate boundaries.

[37] The occurrence of two types of veins along the fault zone, namely the network and fault-fill vein types, represents evidence of cyclic vein formation linked with punctuated and episodic seismic cycles. Accordingly, the difference observed in REE patterns between network and fault-fill veins implies a temporal change in the physicochemical behavior of vein-forming fluids along the fault zone during its history of activation. Considering the similarity of REE patterns between network veins and boudin-neck veins that are distributed widely throughout the mélange, the inhibited sorption or selective complexation reactions that contributed to the composition of the fault-fill vein-forming fluids, would likely reflect faulting-related fluid–rock interactions along the fault plane.

[38] In the Nankai Trough off southwest Japan, the Integrated Ocean Drilling Program (IODP) Nankai Trough Seismogenic Zone Experiment (NanTroSEIZE) has been ongoing with the aim of targeting the plate boundary fault located at a depth of ∼6–7 km below seafloor, as well as the megasplay fault [Tobin and Kinoshita, 2006]. The Mugi mélange is regarded as an on-land analog of the IODP drill target of the plate boundary at depth [Kimura et al., 2012]. Constraints on fluid evolution and changes in the physicochemical features of vein formation demonstrated here could also be tested directly by studies of drill core and long-time borehole monitoring at the planned ultradeep hole.

7. Conclusions

[39] To characterize the origin and behavior of syn-tectonic fluids in the Mugi mélange, Shimanto accretionary complex, southwest Japan, we analyzed the C, O, and Sr isotopic compositions and REE patterns of calcite veins as well as the Sr isotopic compositions and REE patterns of host rocks. Intra-basalt veins are characterized by a carbon isotopic composition of ∼0‰, suggesting that they precipitated from seawater (i.e., formed via the circulation of hydrothermal fluids at a mid-ocean ridge). In contrast, the boudin-neck, network, and fault-fill veins have relatively light carbon isotopic compositions (−10‰ to −19‰), reflecting a mixed origin of carbon derived from both organic matter (∼−25‰) and marine carbonate (∼0‰). The high oxygen isotopic compositions of the vein-forming fluids (+2‰ to +9‰ (SMOW)) and the Sr isotope variations of the veins suggest that rock-buffered fluid was the source for vein minerals, and that such fluid was also affected by both terrigenous sediment and altered oceanic crust. LREE-enriched REE patterns of boudin-neck and network veins suggest that the vein-forming fluids from which they precipitated were affected by sorption processes. In contrast, the source fluid of fault-fill veins was not significantly affected by sorption, nor by complexation processes, suggesting that a temporal change in the physicochemical behavior of the fluids took place only along the fault plane. The details of fluid evolution and physicochemical processes of vein-forming fluids revealed in this study should be of great value in further investigations of faulting, fluid flow, and mass flux along seismogenic plate boundaries.


[40] This research was financially supported by Research Fellowships of the Japan Society for the Promotion of Science for Young Scientists, the Plate Dynamic Program of the Japan Agency for Marine-Earth Science & Technology, the 21st Century Center of Excellence Program of the University of Tokyo, and Grants-in-Aid for Scientific Research from MEXT. Discussions with H. Yamaguchi, M. Toriumi, T. Urabe, Y. Nishio, T. Ishikawa, T. Noguchi, J. Kameda, and Y. Hamada were of great benefit in clarifying our ideas. H. Raimbourg helped to improve an early version of the manuscript. K. Ohbai provided assistance during our time in the field. R. Matsumoto helped with analysis of carbon and oxygen isotopes. This paper was improved by constructive comments of two anonymous reviewers and the Associate Editor, D. M. Saffer.