One shear zone sample (4R-3, 73-76) was found to be large enough to be analyzed by X-ray diffraction (XRD), in order to compare the nature of its clay fraction relative to that of the host matrix. To do so, powders of material were collected by scrapping the shear zone structure and the host matrix, in the rock slab previously used for the preparation of the thin section. The powders were decarbonated and aqueous suspensions were prepared in a solution of 0.5 M NaCl in order to saturate clay minerals with Na. Oriented slides were prepared by drying at room temperature aqueous suspensions on monocrystalline silicon slides to obtain an air-dried (AD) preparation. Ethylene-glycol (EG) solvation of the samples was achieved by exposing them to EG vapor at 70 ∘C for a minimum of 12 h. XRD patterns were recorded on the AD and EG preparations using a Bruker D8 diffractometer equipped with an MHG Messtechnik humidity controller coupled to an Anton Paar CHC+ chamber. Intensities were measured with a SolXE Si(Li) solid-state detector (Baltic Scientific Instruments) for 10 s per 0.04∘2θ step over the 2–50∘2θ Cu Kα angular range. Divergence slits, the two Soller slits, the antiscatter, and resolution slits were 0.3, 2.3, 0.3, and 0.1∘, respectively. Samples were kept at 23 ∘C and a constant 40 % relative humidity (RH) in the CHC+ chamber during the whole data collection. RH was continuously monitored with a hygrometer (uncertainty of ∼ 2 % RH) located close to the sample.

Major- and minor-element analyses were carried out on some of the polished thin sections. Element maps were performed on two samples with an X-ray fluorescence (XRF) spectrometer at Joseph Fourier University in Grenoble (France) to map major elements on large sample surfaces, and with the SEM to map major and minor elements on smaller surfaces. On five shear zone and vein samples, quantitative analyses of eight elements were then performed using a CAMECA SX100 electron probe micro analyzer (EPMA) at Paris VI University. The EPMA was tuned at 15 kV and 10 nA, with a beam focused at 3 µm; a 5 s counting time for Na (measured first), Si, and K; and a 10 s counting time for other elements. The CAMECA set of standards (synthetic and natural minerals or oxides) was used for calibration. The correction methods of Bence and Albee (1968) were used to convert the raw intensity data to weight percent oxides. Analytical uncertainty is 1 %–2 % for Si, Al, Fe, and K; 5 %–15 % for Na, Mg, Ca, and Ti; and 50 %–100 % for Mn, P, and Cr. EPMA analyses were performed as profiles across planar deformation bands, at 3 to 14 µm intervals between each measurement depending on the thickness of the structure. To focus on the analysis of the clay-size fraction of the sediment, the locations of analyzed spots were then visually checked to exclude analyses in individual mineral grains. The remaining analyses were filtered to further exclude quartz or plagioclase (SiO2 + Al2O3 > 80 wt % and sum of volatile elements < 2 wt %), calcite (CaO > 40 wt %), sulfides or oxides (FeO or TiO2 > 30 wt %), and organic matter (sum of major elements < 30 wt %). This filtering resulted in the removal of 10 % to 21 % of the data depending on the profile.

After major-element analyses, one of the samples (21R-2, 82-85) was further selected for in situ measurements of trace-element concentrations. Those measurements were conducted on the rock slabs used for thin section preparation (VFC3 and VFC4). In situ trace-element concentrations were determined at Montpellier 2 University on a Thermo Finnigan Element 2 high-resolution inductively coupled plasma mass spectrometer (HR-ICPMS) using a single-collector double-focusing sector field Element XR (eXtended Range) coupled with laser ablation (LA) system, a Geolas (Microlas) automated platform housing a 193 nm Compex 102 laser from LambdaPhysik. The spot size was set to 102 µm and therefore included multiple mineral grains. Oxide level, measured using the ThO / Th ratio, was below 0.7 %. Silicium 29 was used as internal standard. For each zone (matrix, shear zones, and veins), 29Si was calibrated from the mean value of microprobe analyses in the same zone. Concentrations were calibrated against the NIST 612 rhyolitic glass using the values given in Pearce et al. (1997). Data were subsequently reduced using the GLITTER software using the linear-fit-to-ratio method (Van Achterberg et al., 2001). This typically resulted in a 1 % to 15 % precision (1sigma) for most analyses, evaluated by repeated analyses of USGS reference basalt BIR-1 run as an unknown before and after sample analysis (Table A1 in Appendix). Detection limits were between < 1 and 50 ppb for most trace elements; between 0.04 and 0.5 ppm for Li, B, Cr, Ti, Zn, and As; and between 1 and 15 ppm for Ni, Ca, and Si.

The petrographic inspection of samples revealed the presence of authigenic pyrite and barite in the sediment. Pyrite is by far the dominant authigenic mineral in all the samples. Pyrite displays two crystal morphologies, cubic microcrysts, and blocky macrocrysts. Pyrite microcrysts (≤ 0.5 µm in diameter) occur as isolated crystals or as “framboids”, i.e., circular aggregates, 2–20 µm in diameter (Fig. 3a–b). Framboids in deformation bands sometimes develop tails filled with pyrite microcrysts and resembling pressure shadows (Fig. 3b–d), which may therefore be assimilated to porphyroclastic (grown before the deformation) or porphyroblastic (grown during the deformation) microstructures (Passchier and Trouw, 1998). Blocky pyrite occurs as macrocrysts (> 50 µm in diameter) with euhedral (but never cubic) or anhedral shapes (Fig. 3d). This blocky morphology is scarce compared to the microcrystic and framboidal morphology and occurs both in the matrix and in deformation bands without a particular spatial distribution. Some blocky macrocrysts cut or seal the alignment of phyllosilicates in deformation bands (Fig. 3d) or grow at the expense of framboidal aggregates of microcrysts, which indicates a crystallization stage posterior to the deformation and also posterior to the microcrystic or framboidal morphology. As a general rule, pyrite is more abundant in deformation bands (up to 1.4 % in volume) than in their host matrix (< 0.5 vol %), the largest modal proportions being found in veins (Table 2; Fig. 4). This pyrite enrichment is in fact caused by a greater abundance of microcrysts and framboids in the deformation bands than in the sediment matrix (Figs. 3e–g, 5).

Barite has been observed only in one sample (21R-2, 82-85), as patches of ∼ 5 µm long flakes filling a vein, with pyrite microcrysts and framboids growing on their faces (Fig. 3h).

Ethylene-glycol-solvated XRD spectra of the shear zone and its host matrix (sample 4R-3, 73-76) are presented in Fig. 6. The different peaks reveal the presence of chlorite, kaolinite, quartz, feldspar, and possibly vermiculite in both the shear zone and the matrix. A broad peak at 17 Å (5∘2θ) indicates the presence of smectite or illite–smectite (I/S) mixed layers in the matrix, whereas this peak is absent in the shear zone. Illite is also present in both spectra as evidenced by the sharp peak at 10 Å (9∘2θ). However, the width at half-height of the 10 Å peak is smaller in the shear zone (0.14∘2θ) than in the matrix (0.12∘2θ), indicating a greater crystallinity of illite in the shear zone than in its host sediment (Kübler, 1968).

Two examples of XRF element maps, one in a shear zone and one in a vein, are provided in Fig. 7. An example of SEM element maps, displaying a shear zone cutting a vein, is also provided in Fig. 5. In addition, two examples of EPMA profiles, one in a shear zone and one in a vein, are presented in Fig. 8. Averaged values of major-element analyses are reported in Table 3 and presented in Fig. 9.

A common feature of all the maps and analyses is that major-element concentrations are generally higher in the deformation bands than in the host matrix, indicating a higher density of the material in them. This is in particular true for Al, Si, Fe, K, and S (Figs. 5, 7, 9). Within a given sample, the sum of elements is 11 %–16 % higher in shear zones and 10 %–28 % higher in veins than in the matrix, the highest values being reached in veins (Fig. 8, Table 3). The averaged sum of major elements is in the range 54.95 wt %–63.49 wt % for the clay-size fraction in the sediment matrix, whereas it is in the range 65.07 wt %–71.45 wt % for the clay-size fraction in shear zones and 66.68 wt %–72.91 wt % for the clay-size fraction in veins (Table 3). The comparison of XRF maps (Fig. 7) and EPMA analyses (Fig. 9) shows that the enrichment of Al, Si, and K observed in deformation bands is matched by an enrichment of these elements in the clay-size fraction of deformation bands. There are, however, noticeable exceptions to this general increase in major-element concentrations: the Fe enrichment of deformation bands observed in element maps (Figs. 5, 7) is not observed in the clay-size fraction (Fig. 9). This, and the correlation between Fe and S SEM maps (Fig. 5), shows that the Fe enrichment is due to the preferential growth of authigenic pyrite in deformation bands and does not occur in the clay-size fraction. Calcium is heterogeneously distributed in shear zones and is depleted in veins relative to the matrix (Figs. 7, 9). Sodium, Ti, and Mg do not show any obvious variation between the matrix and deformation bands (Fig. 9).

Trace-element analyses are reported in Table 4 and represented as averaged values in Fig. 10. As for the major elements, a general increase in trace-element concentrations is observed in shear zones and in veins relative to the sediment matrix. For some elements, this enrichment is due to the use of Si as an internal standard, hence with an increased concentration in deformation bands as measured by EPMA (Table 3). Again, however, there are significant anomalies, among which are a strong As enrichment in shear zones (+90 %) and to a lesser extent in veins (+40 %), a depletion of Ba in all the deformation bands (-31 % to -33 %), and an enrichment of Li and B in veins (+25 % to +38 %).

The first result of our study is that deformation bands are more compacted than the host sediment. This greater compaction is indicated by the general increases in element concentrations observed in SEM and XRF maps. It is also apparent in the EPMA data, as the concentration of volatile elements (approximated by 100 minus the sum of elements) is 36.51 wt %–45.05 wt % in the matrix whereas it is only 28.55 wt %–34.93 wt % in shear zones and 27.09 wt %–33.32 wt % in veins (Table 3), suggesting smaller interfoliar spaces in clays from deformation bands than from the matrix, and thus a smaller porosity. Shear zones and veins have therefore resulted in a greater volatile loss than the matrix, which may be interpreted as a greater fluid expulsion (Fig. 8). This result supports the conclusion of Milliken and Reed (2010) that deformation microstructures in mud core samples from Site C0008 (IODP Exp. 316) formed primarily by mechanical compaction of the unconsolidated mud and thus are indeed “dewatering structures”. This strain-induced reduction of pore spaces and density increase is likely responsible for the observed enrichment of many major and trace elements in deformation bands compared to those of the matrix. However, mechanical compaction should raise all the analyzed elements by the same proportion. This hypothesis may be tested using a correction for compaction by normalizing the sum of element concentrations to 100 %. Even after such correction, some elements like As and to a lesser extent Cu, Sb, and Pb remain enriched in deformation bands relative to the matrix and positively correlated to each other, while some like Ba are depleted in deformation bands (Figs. 10, 11). Veins also appear enriched in B, Li, and perhaps K2O but slightly depleted in CaO and perhaps Na2O relative to the matrix. This indicates that some chemical reactions occurred in the deformation bands, in addition to mechanical compaction. In the following, we explore the processes that may be responsible for the chemical differences observed in shear zones and veins relative to their host sediment matrix.

The dark color of deformation bands (Fig. 2), their increased content in authigenic pyrite microcrysts (Figs. 3e–g, 4, 5), their enrichment in chalcophile elements (Cu; As; Sn; Pb, Figs. 10, 11), and the positive correlations among these elements (Fig. 11) all concur to indicate that sulfide precipitation was enhanced in deformation bands compared to the matrix. Sulfide mineralization in low-temperature (< 60 ∘C) sediments is essentially a byproduct of the anaerobic degradation of organic matter by microorganisms (see the review by Megonigal et al., 2004). For this reason, authigenic pyrite is generally observed in anoxic, organic matter-rich sediments such as shales. The framboidal morphology of pyrite itself is generally, though not exclusively, taken as an indicator of microbial proliferation (Barbieri and Cavalazzi, 2005; Cavalazzi et al., 2012; Chen et al., 2006; Merinero et al., 2009; Peckmann et al., 2001). Organic carbon decomposition by heterotrophic microorganisms has the effect of releasing reduced dissolved iron and hydrogen sulfide, through Fe3+ and sulfate reduction reactions such as the following: R1CH2O+4Fe(OH)3+7CO2→8HCO3-+4Fe2++3H2O, R22CH2O+SO42-→2HCO3-+H2, where CH2O represents organic matter. CO2 in Reaction (R1) may be supplied by methanogenesis: R32CH2O→CO2+CH4. The products of these reactions are then involved in a suite of secondary redox reactions, among which pyrite (FeS2) precipitates via the formation of temporary sulfide phases (Hunger and Benning, 2007; Wilkin and Barnes, 1997): R4Fe2++2HS-→FeS2+H2. The precipitation of multiple pyrite microcrysts forming framboids requires a nucleation rate significantly greater than the crystal growth rate (i.e., surface-controlled growth), a condition achieved as long as the supply of reactants is not limited and the thermodynamic conditions are far from equilibrium (Ohfuji and Rickard, 2005; Wilkin and Barnes, 1997). This suggests that sulfate and reduced iron supplies were not limited at the time of deformation, and thus that the deformation bands occurred within the sulfate reduction zone. The Ba depletion in deformation bands (Figs. 10, 11) and the growth of pyrite microcrysts and framboids on barite needles in a vein (Fig. 3h) also support sulfate bio-reduction via metabolic Reactions (R1) and (R2). All these lines of evidence lead to the conclusion that the preferential crystallization of microcrystic and framboidal pyrite in deformation bands is a result of enhanced bacterial activity. According to syn-deformation microstructures of pyrite growth (Fig. 3b, c), this enhanced bacterial activity was coeval with the development of deformation bands and the associated reduction of porosity.

On the contrary, the blocky morphology of large and rare pyrite crystals indicates a transport-controlled crystal growth (Ohfuji and Rickard, 2005; Wilkin and Barnes, 1997), meaning that the demand of sulfate and iron exceeded their supply. These large blocky pyrite crystals grow at the expense of framboids and seal the CPO of deformation bands (Fig. 3d). This suggests that blocky pyrite is a recrystallization form of the framboid morphology, occurring without deformation, below the sulfate–methane transition zone, and hence not necessarily via metabolic reactions.

In tandem with pyrite diagenesis, XRD spectra and chemical analyses suggest that the development of deformation bands is accompanied by modifications of the clay mineralogy. In the only sample studied by XRD (4R-3, 73-76), the shear zone displays a disappearance of smectite or I/S mixed layers, and an increased crystallinity of illite, relative to its host matrix. In the other sample studied for trace elements (21R-2, 82-85), the correlated B and Li enrichments of the 12 shear zone and vein analyses relative to the matrix, particularly noticeable in veins (Fig. 11), are two additional arguments suggesting that deformation bands localize smectite transformation into illite. Indeed, illitization is known to result in an uptake of B and Li, as the former element substitutes in tetrahedral sites of illite and the later in octahedral sites (Williams et al., 2013).

Because of its importance as a reaction releasing freshwater, the destabilization of smectite and its transformation into illite has received much attention for the fluid budget of accretionary prisms. Heat and the availability of K+ are considered as the primary factors governing illitization, and time, pressure, or shear stress as secondary factors (Casciello et al., 2011; Ransom and Helgeson, 1995). Given the small size of the microstructures studied here, however, it is quite unlikely that any of these factors varied much between deformation bands and their matrix just a few millimeters apart. Biotic alteration may offer an alternative, more plausible explanation for the observed changes in the mineralogy of clays. Indeed, smectite and I/S mixed layers are mineral structures that are very sensitive to biotic alteration, because they are needed for anaerobic bacteria to reach structural Fe3+ (Dong et al., 2009; Esnault et al., 2013; Zhang et al., 2012). There are strong suspicions that microbial reduction of smectite and I/S mixed layers can trigger illitization (Esnault et al., 2013; Koo et al., 2014, 2016). Accordingly, we propose that enhanced anaerobic bacterial proliferation in deformation bands, a process already suggested by the localized crystallization of framboidal pyrite, is also the cause of the disappearance of smectite or I/S mixed layers, and the increased crystallinity of illite. This conclusion is speculative for the time being given the small corpus of data presented here (one shear zone sample analyzed for XRD and 12 trace-element analyses of shear zones and veins in another sample). Future work will have to test the reproducibility of these findings and their applicability at larger scale in accretionary prisms.

A first question that arises from the above discussion is why did the deforming sediment provide a more favorable ground to the development of anaerobic microorganisms than the undeformed matrix? A first possible explanation is that the greater compaction of deformation bands increased the availability of compounds necessary for metabolic Reactions (R1) and (R2) by increasing the concentration of these compounds compared to the matrix. As shown above, the majority of element concentrations are raised by ∼ 20 % due to the increased compaction of deformation bands (Figs. 8, 10), and thus the same enrichment is expected for organic matter, sulfate, and Fe3+. This concentration increase may have favored the proliferation of anaerobic microorganisms. Alternatively or additionally, it has been shown that the deformation of silicate minerals and the delamination of clay layers in particular generate H2 by chemical reactions between water and mechanoradicals created by the rupture of Si–O–Si or Al–O–Si bonds (Hirose et al., 2011; Kameda et al., 2004; Kita et al., 1982; Wakita et al., 1980). This H2 may have served as a terminal electron donor in heterotrophic metabolic reactions, acting in conjunction with mechanical compaction to favor anaerobic microbial proliferation during deformation of the unconsolidated sediment.

Another important question concerns the timing of deformation bands and their bacterial proliferation. Given the need of nutrients for metabolic reactions, it is tempting to interpret these structures as formed at shallow depth below the sea floor, in proximity of seawater sulfate supply. However, shear zones as well as veins were almost exclusively found in the accretionary prism (Unit II) and not in the slope sediment (Unit I) above the unconformity (Fig. 1c). This fact implies that most of the deformation bands studied here are not burial-related but are rather associated with the tectonics of the accretionary prism. A way to reconcile the two inferences is to suggest that deformation bands, and biological diagenesis in them, developed in the upper portion of the accretionary prism during thrusting, and before the deposition of slope sediments. Whether deformation bands are mechanically compatible with thrusting is unfortunately unknown because no kinematics could be assigned to the majority of them. Nevertheless, we note that this proposed timing coincides with the activity of the megasplay fault thrust uphill of C0001 (Fig. 1b). It is thus possible that deformation bands may represent early stages of strain localization, and fluid expulsion, in the context of megasplay fault development.

Reports of all the drilling expeditions across the Nankai accretionary prism are consistent in describing the ubiquitous existence of dark shear zones and veins in mud sediments (Ashi et al., 2008; Henry et al., 2012; Kinoshita et al., 2009a, b, c, d, e, f; Saffer et al., 2010; Ujiie et al., 2004). Dark deformation bands have also been described on land in mudstones of the Boso peninsula paleo-accretionary prism (Ohsumi and Ogawa, 2008), and worldwide in active continental margins (e.g., Behrmann et al., 1988), sometimes in conjunction with an enrichment in authigenic iron sulfides (Lindsley-Griffin et al., 1990). Dark deformation bands may thus be considered as an intrinsic feature of mudstone sediments in accretionary prisms. Our results suggest that these small structures localize pyrite crystallization and smectite–illite transformations, two diagenetic reactions probably mediated by the proliferation of anaerobic microorganisms, and both boosted by deformation. A possible implication of our study is that such increased diagenesis in deformation bands may be a source of freshwater, firstly because H2O is a product of metabolic reactions leading to pyritization (see Eq. 1 for instance), and secondly because H2O is also released by the illitization of smectite. In pervasively deformed mudstones with abundant deformation bands, these two effects may combine to potentially explain the local deficits of modeled versus observed pore water freshening found in previous studies (e.g., Saffer and McKiernan, 2009). Given the temporal consistency between megasplay faulting and deformation bands, the dewatering of these many microstructures could be supplied to major faults, which might explain some freshwater fluxes observed in accretionary prisms (e.g., Kastner et al., 1993; Vrolijk et al., 1991). More work is obviously necessary to quantify the contribution of microbial diagenesis to the pore water freshening of deformed sediments.

Another important implication of our study is that the microbial diagenesis might be a general feature of deformation structures whatever their size, from the small deformation bands described here to major thrust faults. This is particularly the case of the megasplay fault of the Nankai prism drilled at 270 mbsf at site C0004 during the NanTroSEIZE expedition 316 (Kinoshita et al., 2009c). The analysis of the core zone of this megasplay revealed a 2 cm thick dark gouge, interpreted as the principal slip zone and the subject of a vigorous scientific debate. Indeed, this dark gouge was found to combine an increased vitrinite reflectance, a smectite depletion, and an increased illite crystallinity compared to the surrounding breccia (Sakaguchi et al., 2011; Yamaguchi et al., 2011). These differences were interpreted as thermal maturation of organic matter and illitization in the dark gouge, both due to coseismic shear heating above 380 ∘C in the principal slip zone of the megasplay, and thus as evidence of seismic rupture propagation to the sea floor. However, other studies using trace-element analyses as well as thermal modeling contradicted this interpretation, by concluding that the coseismic temperature rise in the megasplay did not exceed 300 ∘C and could not be sufficient to activate the kinetics of illitization (Hirono et al., 2009, 2014). How can these two divergent views of the fault zone chemistry and mineralogy be reconciled?

The paradox might be solved by considering bacterial activity as a diagenetic process boosted by deformation. It is important to note that a peak bacterial concentration of 3.6×109 cells/cm3 was found in the damaged zone of the megasplay drilled at site C0004 (Kinoshita et al., 2009c). This cell abundance, noticeably in the upper range of sediments worldwide (typically of 106–109 cells/cm3), was also nearly as high as that of the sea floor at this site. Moreover, the lower part of the dark gouge in the megasplay fault drilled at site C0004 was also enriched in authigenic pyrite (Destrigneville et al., 2013), which makes diagenetic reactions in the megasplay very similar to those found in the tiny deformation bands studied here. Such an analogy strongly suggests that temperature rise is not the only factor capable of explaining the observed anomalies in the megasplay, and that many mineral reactions in the principal slip zone might be biologically mediated rather than thermally activated. Recent studies have shown that anaerobic microbial activity associated with framboidal pyrite crystallization could locally yield anomalously high reflectance values of amorphous organic matter compared to thermal maturation (e.g., Synnott et al., 2016). Microbial activity could thus explain the elevated vitrinite reflectance temperature obtained for the principal slip zone (Sakaguchi et al., 2011), as it does explain pyrite crystallization, smectite or I/S mixed layer destabilization, and illitization in deformation structures whatever their size. Anaerobic bacterial bloom, boosted by deformation, would thus solve the paradox of mineralogical and chemical anomalies in the principal slip zone of the Nankai megasplay. We therefore come to the conclusion that more caution is required before interpreting illitization in fault gouges as an evidence of coseismic slip, or more generally as a temperature rise, because this mineral reaction, like others, might be triggered or enhanced by microbial activity.