# Impact of Strain on Carbonaceous Matter Crystallinity: Insights from Raman Spectroscopy and microstructural analysis of strain gradients from exhumed accretionary complexes

Hugues Raimbourg, Benjamin Moris-Muttoni, Romain Augier, Kristijan Rajič, Ismay Vénice Akker, Aurélien Canizarès, Emmanuel Le Trong

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## ► To cite this version:

Hugues Raimbourg, Benjamin Moris-Muttoni, Romain Augier, Kristijan Rajič, Ismay Vénice Akker, et al.. Impact of Strain on Carbonaceous Matter Crystallinity: Insights from Raman Spectroscopy and microstructural analysis of strain gradients from exhumed accretionary complexes. Journal of Structural Geology, 2026, 203, pp.105560. 10.1016/j.jsg.2025.105560. insu-05373439v2

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Submitted on 3 Dec 2025

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# Impact of strain on carbonaceous matter crystallinity: Insights from Raman spectroscopy and microstructural analysis of strain gradients from exhumed accretionary complexes

Hugues Raimbourg<sup>a, *, †</sup>, Benjamin Moris-Muttoni<sup>a, †</sup>, Romain Augier<sup>a</sup>, Kristijan Rajić<sup>a, b †</sup>, Ismay Vénice Akker<sup>c, †</sup>, Aurélien Canizarès<sup>d</sup>, Emmanuel Le Trong<sup>a</sup>

## ARTICLE INFO

### Keywords:

- Carbonaceous matter maturity

- Raman spectroscopy

- Deformation

- Microstructures

- Carbonization

## ABSTRACT

Carbonaceous matter is a common component of metasedimentary rocks, mainly deriving from the diagenetic and then metamorphic evolution of the organic matter originally present in the sedimentary protolith. During burial, the maturity of carbonaceous particles increases by rearrangements of the aromatic carbon sheets and expulsion of heteroatoms, mostly as a result of temperature increase. However, additional external factors could enhance the maturity such as the pressure, fluid-rock interactions and/or deformation. In this work we explored the effect of strain at low metamorphic temperatures (in the range ~200–320 °C) along natural strain gradients of variable scales, in three accretionary or collisional complexes (Shimanto and Kodiak accretionary complexes, Infrahelvetic Flysch Units), using Raman spectra of carbonaceous matter. In these examples, both ductile and brittle zones of localized deformation were observed, in the form of shear and breccia zones, respectively. Carbonaceous particles observed by TEM are ~50–100 nm in size and are distributed throughout the rocks, enabling automated Raman spectroscopy profiles and maps to be carried out. Irrespective of the deformation process, the intensity ratio between D and G band (R1 ratio) of Raman spectra is systematically increased by up to ~60 % in high strain zones compared to the least deformed, reference zones. From the comparison with a series of undeformed metamorphic rocks spanning the same temperature range, the R1 increase reflects the progressive organization of the carbonaceous matter towards higher maturity. In mm-scale shear bands, the increase in R1 can be unambiguously ascribed to localized strain, whereas in brecciated domains, localized temperature increase may also have contributed to R1 anomaly.

## 1. Introduction

The Raman Spectroscopy of Carbonaceous Matter (RSCM) allows quantifying the degree of maturity of carbonaceous matter (CM) that derives from the transformation of organic matter originally present in the sediment. Considered as a reliable indicator of the metamorphic grade (Pasteris and Wopenka, 1991; Wada et al., 1994), the quantitative evolution of maturity of CM has been used to propose series of geo-thermometers calibrated with natural data between ca. 100 and more than 650 °C (Aoy a et al., 2010; Beyssac et al., 2002a; Kouketsu et al., 2014; Lahfid et al., 2010; Rahl et al., 2005). The irreversible character of

CM aromatization reactions allows retrieving peak-metamorphic temperature conditions that are consistent with classical thermometric estimates (Augier et al., 2005; Delchini et al., 2016). As a result, this method is widely used in metamorphic petrology as a robust and non-destructive method to retrieve peak temperatures for contact metamorphism (Aoy a et al., 2010; Hilchie and Jamieson, 2014; Süs senberger et al., 2020), or for regional metamorphism (Beaudoin et al., 2015; Bellanger et al., 2015; Clerc et al., 2015; Nibourel et al., 2018).

In the latter case, in addition to long-term temperature increase, rocks experience an amount of deformation that is, in many instances of collision belts and subduction zones, penetrative over large distances.

<footnote><sup>a</sup> Institut des Sciences de la Terre d'Orléans (ISTO), Université d'Orléans, CNRS, BRGM, ISTO, UMR 7327, F-45071, Orléans, France</footnote>

<footnote><sup>b</sup> Department of Earth Sciences, Durham University, Science Laboratories, South Road, Durham, DH1 3LE, United Kingdom</footnote>

<footnote><sup>c</sup> University of Padua, Department of Geosciences, Via G.Gradenigo 6, 35131, Padova, Italy</footnote>

<footnote><sup>d</sup> CEMHTI (Conditions extrêmes et matériaux: haute température et irradiation), UPR3079 CNRS, Orléans, France</footnote>

<footnote>* Corresponding author. ISTO, Université d'Orléans, Orléans, France.</footnote>

<footnote>E-mail address: hugues.raimbourg@univ-orleans.fr (H. Raimbourg).</footnote>
This is the case in the Schistes Lustrés Complex, France (Agard et al., 2001), and the Sambagawa Belt, Japan (Osozawa and Wakabayashi, 2015; Toriumi and Noda, 1986), which were used in calibrating a geo-thermometer based on CM maturity (Beyssac et al., 2002a). These

examples of pervasively deformed metamorphic terranes in convergent settings raise the question of the effect of strain on CM maturity.

The influence of deformation on Raman spectra of CM has been principally studied in structures of brittle deformation. For example, the

<figure><img src="image_3.png" /><figcaption>**Fig. 1.** Geological maps of the three studied areas. a) Geological map of the Shimanto Belt (Japan) b) Geological map of the Kodiak Accretionary Complex (Alaska, USA) c) Geological map of the Glarus Thrust area (Switzerland), modified from (Akker et al., 2018; Dielforder et al., 2016; Murata, 1997; Wilson et al., 2005); Stars localize the sampled outcrops or samples. Sample sites in bold writing correspond to the strain gradient samples and classic writing correspond to undeformed samples used for the regional calibration of the intensity ratio R1.</figcaption></figure>
intensity ratio (i.e. R1 in Beyssac et al., 2002a) has been used to detect drastic CM maturity variations in fault gouges and breccia (Furuichi et al., 2015; Ito et al., 2017; Kouketsu et al., 2017; Kuo et al., 2018; Ujiie et al., 2021). However, these applications implicitly assumes that CM spectra reflect only the thermal record, irrespectively of the potential impact of strain on CM maturity (Bonijoly et al., 1982; Bustin et al., 1995; Ross and Bustin, 1990; Tagiri and Tsuboi, 1979). Aseismic deformation has also been explored in order to understand the effect of strain on CM maturity in high-temperature rocks (Bonijoly et al., 1982; Bustin et al., 1995; Ross and Bustin, 1990; Ross et al., 1991; Suchy et al., 1997). These studies have shown the increase in CM maturity using RSCM, Vitrinite Reflectance and Transmission Electron Microscopy (TEM) in highly strained rocks. They have focused on the final graphitization transformation at high temperatures through natural examples (Barzoi, 2015; Barzoi and Guy, 2002; Stokes et al., 2024; Suchy et al., 1997) or slow creep experiments (Bonijoly et al., 1982; Bustin et al., 1995; Ross and Bustin, 1990; Ross et al., 1991). In contrast, the effect of deformation on CM maturity during the carbonization stage and early stages of graphitization (from 200 °C to 320 °C) remains poorly studied. This temperature range corresponds to a transition in deformation mechanisms and the structures that form are diverse, from brittle structures formed during earthquakes to low strain-rates shear zones (Meneghini et al., 2009, 2010), making the effect of deformation particularly complex to interpret (Nakamura et al., 2015). In recent deformation experiments performed in such metamorphic conditions, it has been shown that strain energy has a catalytic effect to enhance CM maturity (as apparent in R1 increase), independently from temperature (Moris-Muttoni et al., 2023), calling for a detailed examination of this effect in naturally strained rocks.

In this study, we present new results on the effect of strain on the Raman signal of carbonaceous particles during the carbonization and early graphitization stages along natural strain gradients from the Shimanto Belt, the Kodiak Accretionary Complex and the Infrahelvetic Flysch Units (IFU). In these examples, we combined evaluations of CM maturity with microstructures, using Raman Spectroscopy on CM particles that were localized within well-characterized natural strain gradients at scale ranging from mm to ~10m. In addition, a larger set of poorly deformed samples, including a series of rocks from the Variscan belt in the southern Massif Central area, France (Montmartin, 2021), was analyzed to connect RSCM parameters with prograde metamorphic evolution and the increase in crystalline ordering of CM. Finally, combining these datasets, we discuss the effect of deformation on CM maturity, including both aseismic deformation recorded within shear bands as well as seismic slip deformation experienced along faults.

## 2. Geological settings and sampling strategy

To address the possible effects of deformation on the maturity of CM, we selected different sampling areas sharing similar peak-temperatures, in the range of 200–320 °C, and a well-understood tectono-metamorphic evolution: the Shimanto Belt, Japan, the Kodiak Accretionary Complex, Alaska, and the IFU, Switzerland (Fig. 1 and Table 1). In each studied area, we carried out sampling of strongly deformed rocks, where strain is

either distributed in a network of shear zones or strongly localized within major tectonic structures. In addition to deformed samples, we selected poorly deformed or undeformed samples, which served (i) as local ‘reference samples’ to compare with adjacent, strongly deformed material, and (ii) as a sequence of samples to analyze the evolution of Raman parameters with increasing metamorphic grade. Please note that the metamorphic temperature estimates below are provided with an error bar that corresponds to the standard deviation of the individual CM analyses, to which a systematic error of ± 30 and ± 50 °C should be added, due to the calibration of the low T (Lahfid et al., 2010) or high T (Beyssac et al., 2002a) geothermometer, respectively.

### 2.1. The Shimanto Belt

The Shimanto Belt is an exhumed portion of a large-scale, Mesozoic to Cenozoic subduction complex that extends over 800 km in Japan (Charvet and Fabbri, 1987; Taira, 1981). The subduction complex is characterized by the accretion of NW-dipping sedimentary units, and shows a southeastward younging trend of the lithostratigraphic formations toward the current location of the Nankai Trough.

In Kyushu Island, two geological formations were selected, the Hyuga Group (including the Coherent Hyuga unit and the Hyuga Tectonic Mélange unit) and the Foliated Morotsuka Formation, from bottom to top in structural order (Saito et al., 1996; Saito et al., 1997; Fig. 1a). The Coherent Hyuga unit is mostly composed of undeformed to low-strain turbidites made of alternating shales and sandstone layers, with a thickness from 5 cm to a few meters. The Hyuga Tectonic Mélange is in general composed of blocks of metasandstones and metabasalts wrapped by a foliated shaly matrix. In the sampling sites of this study, the blocks are mm-to dm-sized elongated sandstone lenses wrapped in a finely foliated shaly matrix (Fig. 2). Deformation is distributed along an anastomosing network of top-to-the-SE shear zones deflecting the foliation and observable at all scales (Fig. 2). The Foliated Morotsuka Formation is composed of phyllosilicates-rich lithologies and marked by a penetrative metamorphic foliation bearing a strong lineation defined by white micas and chlorite (Raimbourg et al., 2014a). Vertical thinning is apparent through conjugate sets of small-scale shear zones at low angle with the metamorphic foliation (Raimbourg et al., 2014a).

Results from RSCM indicate that peak-temperatures conditions reached are around 215 ± 20 °C in the Coherent Hyuga, 260 ± 10 °C in the Hyuga Tectonic Mélange and 345 ± 15 °C in the Morotsuka Group (Palazzin, 2016; Raimbourg et al., 2014a, 2014b, 2017). For the latter formation, temperatures derived from illite-chlorite assemblages formed at peak metamorphic conditions yielded 320 ± 14 °C and 3.9 ± 0.4 kbar (Rajić et al., 2023).

The Hyuga Tectonic Mélange was sampled to establish the effect of strain localized in the network of small-scale shear zones, whereas poorly deformed samples in the Coherent Hyuga, Hyuga Tectonic Mélange and Morotsuka Group were selected to analyze the evolution of RSCM parameters with increasing metamorphic grade.

<caption>**Table 1**</caption>

Outcrop locations, temperature and structures from the Shimanto Belt, the Kodiak Accretionary Complex and the Infrahelvetic Flysch Units area.

<table><thead><tr><th rowspan="2">Geological setting</th><th rowspan="2">Unit</th><th rowspan="2">Outcrop</th><th colspan="2">Location</th><th rowspan="2">Strain gradient</th><th rowspan="2">Deformation microstructure/ structure</th><th rowspan="2">Regional temperature</th></tr><tr><th>Latitude</th><th>Longitude</th></tr></thead><tbody><tr><td rowspan="3">Shimanto Belt</td><td rowspan="3">Hyuga mélange</td><td>18NOB07</td><td>32.521972</td><td>131.384222</td><td>mm-scale</td><td>shear zone</td><td rowspan="3">260 ± 10 °C</td></tr><tr><td>18NOB08</td><td>32.490444</td><td>131.381472</td><td>mm-scale</td><td>shear zone</td></tr><tr><td>HN64</td><td>32.594419</td><td>131.462973</td><td>mm-scale</td><td>shear zone</td></tr><tr><td>Kodiak Accretionary Complex</td><td>Ghost Rock Formation</td><td>19KO14</td><td>57.422027</td><td>-152.480361</td><td>m-scale</td><td>Fault zone</td><td>220 ± 10 °C</td></tr><tr><td rowspan="2">Glarus Thrust</td><td rowspan="2">Infra-helvetic Complex</td><td>20GT03A</td><td>46.996830</td><td>9.091400</td><td>mm-scale</td><td>Breccia in contact with a fault</td><td rowspan="2">270 ± 10 °C</td></tr><tr><td>20GT19</td><td>46.954310</td><td>9.237210</td><td>m-scale</td><td>Glarus Thrust Zone</td></tr></tbody></table>
<figure><img src="image_4.png" /><figcaption>**Fig. 2.** Structures of deformation in Shimanto Belt in Japan (a and b) and Kodiak Complex in Alaska (c). In a) and b), the foliation S1 is defined by the alternation of lenses of quartz-rich and phyllosilicate-rich layers. The deformation is visible as a dense network of shear zones (shown by red arrows), at all scales, deflecting the lenses of sandstones, as well as the abundance of precipitated quartz filling mode I cracks in the more rigid sandstone lenses (yellow arrows). In c, a multilayered, fine-grained fault zone (i.e., the Black Fault Rock, or BFR) cutting across cataclastic rocks from the Ghost Rock Mélange, Pasagshak area, in the Kodiak Complex. (For interpretation of the references to colour in this figure legend, the reader is referred to the Web version of this article.)</figcaption></figure>

## 2.2. *The Kodiak Accretionary Complex*

The Kodiak Accretionary Complex (Fig. 1b), in southeast Alaska, is an exhumed portion of a Jurassic to Cenozoic NE-SW trending accretionary wedge. It is composed of northwest-dipping thrust-bounded stacked sedimentary units, showing a younging trend of the lithostatigraphic formations toward the southeast, i.e., towards the current location of the Aleutian Trench (Byrne and Fisher, 1987; Connelly, 1978; Moore et al., 1983; Sample and Moore, 1987). The Ghost Rock tectonic mélange lies in the southwest of Kodiak Island, and comprises blocks of sandstones and basalts embedded in an deformed argillaceous matrix (Moore et al., 1983). The metamorphic conditions of deformation, estimated from RSCM thermometry, are of the order of ~220 ± 10 °C (Rajić et al., 2023), and the kinematics derived from fault zones is NW-SE shortening. Particular fault zones cutting across Ghost Rock mélange have been described in Pasagshak Peninsula (Rajić et al., 2023; Rowe et al., 2005, 2011): the fault core contains several juxtaposed layers of ultrafine-grained material (hence they are called Black Fault

Rocks, or BFR) and have been interpreted as pseudotachylyte. The faults containing the BFR are considered as “minor”, by contrast with the Glarus Thrust examined hereafter, in the sense that they are enclosed within the Ghost Rock Formation, and are not associated with any significant gap in metamorphic temperature (Rajić et al., 2023).

## 2.3. *The Glarus Thrust and Infrahelvetic Complex flysches*

The Glarus Thrust is an Oligocene to Miocene out-of-sequence thrust exposed in the eastern part of the Swiss Alps and covers a surface of a 600 km² in the Kanton Glarus, southeast of Zürich, Switzerland (Fig. 1c). The thrust juxtaposes, by a displacement of about 40–50 km (Pfiffner, 1985), the allochthonous units of the Helvetic Complex over the Infrahelvetic Complex (Pfiffner, 1981, 1993). In this area, the lowermost unit of the Helvetic Complex, i.e., the hanging wall of the Glarus Thrust, is composed of a Permian-Triassic volcano-sedimentary unit, the Verrucano formation. In the footwall of the Thrust, the Infrahelvetic Complex is a result of the imbrication of several allochthonous Flysch units as well
as a para-autochthonous series composed of a crystalline basement covered by Mesozoic carbonate units and Cenozoic flysch units (Akker et al., 2018; Burkhard et al., 1992; Pfiffner, 1977). The different autochthonous and allochthonous flysch units form the Infrahelvetic Flysch Units (IFU).

The folded and imbricated IFU (Pfiffner, 1977) shows a gradual, north to south, increase in metamorphic conditions, from the anchizone to the lower greenschist-facies metamorphism (i.e. 200 °C–320 °C; (Ebert et al., 2007; Lahfid et al., 2010; Rahn et al., 1995), marked by variations in the fabric of the slates making up the IFU (Akker et al., 2021a, 2021b).

Two zones of very concentrated deformation were sampled in the IFU (Fig. 3). The first one (20GT19) corresponds to the Glarus Thrust itself, on the SE side of Foostock, marked by a 20 cm-to 5m-thick (in rare case) layer of strongly deformed calc-mylonites (Badertscher and Burkhard, 2000; Schmid, 1975). These mylonites are composed of a complex imbrication of para-autochthonous carbonates and flysch slivers from the IFU (Badertscher and Burkhard, 2000). There, the IFU was sampled at increasing distance from the mylonites of the core of the Glarus Thrust, down to less deformed rocks selected a few meters away to the contact.

The second studied area corresponds to a second order strain gradient developed within the IFU (20GT03) in the vicinity of a subordinate faults contributing to the imbrication of the flysch units, showing the progressive development of foliation and crenulation cleavage (Akker et al., 2018). There, rocks were selected along the fault plane (20GT03A).

## 2.4. Montagne Noire area

In order to extend the database of undeformed, reference samples where Raman spectra parameters are correlated to the

paleotemperatures of metamorphism, we used a series of samples collected in the metasedimentary formation located in the southern flank of the Montagne Noire, a migmatite dome formed during the Variscan orogenesis, in France (Montmartin et al., 2021). The metamorphic temperatures inferred from RSCM in these samples correlate well with other independent paleothermometers, e.g. conodont alteration index method (Wiederer et al., 2002). The temperature gradient related to the distance to the dome cuts across all tectonic structures and is therefore considered as postdating the tectonic nappe stacking (Montmartin et al., 2021). Hence, in such a metamorphic gradient, the influence of deformation on RSCM can be excluded.

# 3. Material and methods

## 3.1. Rock samples preparation

All rock samples (Table 1 and Supplementary Material Table S1) were cut in the XZ structural plane, parallel to the direction of the stretching lineation and orthogonal to the main foliation plane. Both standard thin (30 µm) and thick (200 µm) sections were prepared either for microstructural analyses or for RSCM analyses, respectively.

## 3.2. RSCM

All the particles of carbonaceous particles contained in the meta-sediments of this study have a size below the micrometer, preventing to analyze their geometry with optical microscopy. Some of the particles can still be identified as they appear as bright spots in reflected light. A series of Raman spectroscopic analyses was carried out on these individual CM particles. In addition, some of the samples contain, in addition to the visible ones, many more distributed carbonaceous particles, too small to be identified in optical microscopy but yielding a

<figure><img src="image_5.png" /><figcaption>**Fig. 3.** Structures of deformation in the footwall of the Glarus Thrust. a) Panoramic view on the Glarus Thrust, from the Segnes Pass. b) Strongly deformed calcschists from the Infra-Helvetic Flysch, in Segnes Pass area, with boudinaged lenses of carbonates crosscut by precipitated calcite filling in mode I cracks. c) Fault network within the Infra-Helvethic Flysch, below the Glarus Thrust, in the Kärpferbrücke area, with boudinaged, meter-scale lenses of limestone, and shales developing a penetrative crenulation cleavage. (c) corresponds to sampling site 20GT03.</figcaption></figure>
Raman signal. These distributed carbonaceous particles allowed the automated acquisition of Raman spectra along line profiles or maps.

Oriented thick sections were analyzed with ReniShaw InVIA micro-spectroscope, at ISTO-BRGM, Orléans, using an argon laser beam of 514 nm focused on the sample by a DM2500 Leica microscope equipped with 100× objective. This apparatus is operated using software Wire 4.0 (Renishaw). To avoid heating of CM during analysis, the laser power was set and maintained at 0.5 mW (1 %), yielding a power at the surface of the sample measured as 55 µW with the 100× objective.

### 3.2.1. Point analysis

This first method of acquisition consists in point analysis at locations chosen by the user as coinciding with bright spots of CM particles. The spectra resulted from the accumulation of repeated, 6 to 40 times, acquisitions of 5s. CM was systematically analyzed below transparent adjacent minerals, particularly quartz and calcite depending on the mineralogy of the metasediment (Ammar and Rouzaud, 2012; Katagiri et al., 1988; Pasteris, 1989).

Analyses within any given sample were either performed randomly (to get an average spectrum) or along series of cross-sections all oriented perpendicular to the orientation of the shear zones (i.e., in mm-scale strain gradients). Spacing between analyses were drastically reduced and concentrated, from few tens of micrometers to less than 10 µm, when approaching the core of the deformation structures. X-Y coordinates of point analyses were all frame-referenced. For average measurements (i.e., “reference samples” and Glarus Thrust sampling profile), on each rock sample, at least 20 point analyses were acquired randomly distributed over the entire thick section.

Raman spectra were then processed using PeakFit v4.12. To gather the best comparison between CM spectra from the sheared-domain and the poorly deformed-domain, only spectra presenting a planar baseline were considered and linear baseline subtraction function was applied. To determine the baseline, two bands of 100 cm⁻¹ of width were selected at ca. 900 and ca. 1900 cm⁻¹ as local references. The main Raman spectra parameter used in this work was the intensity ratio R1 (Beyssac et al., 2002a) following equation (R1) = I<sub>D</sub>/I<sub>G s.l.</sub>, where I<sub>D</sub> and I<sub>G s.l.</sub> represent the intensity of the D band and G<sub>sensu lato</sub> (G s.l.) band, respectively (Fig. 4). In other words, D and G s.l. are the two maxima, around 1350 cm⁻¹ and 1600 cm⁻¹, of the Raman spectrum.

In addition, the deconvolution procedure by Lahfid et al. (2010) was applied. This approach appears particularly well adapted to low grade metamorphic rocks consisting in 5 bands (D1, D2, D3, D4 and G) deconvolution using Lorentzian function (Fig. 4). D4+D1 bands form

the Defect massif (D band) centered around 1350 cm⁻¹ while G + D2 bands form the raw Graphite massif (G s.l. band) around 1600 cm⁻¹. The D3 band forms a very wide band of low intensity located between the two main massifs at ca. 1500 cm⁻¹ (Beny-Bassez and Rouzaud, 1985). By using the area ratio RA1 = (D1+D4)/(D1+D2+D3+D4+G) (Lahfid et al., 2010), it is possible to obtain a reliable temperature estimate in the range of 200–320 °C with ±30 °C calibration-related uncertainty (RA1 = 0.0008 × T<sub>max</sub> + 0.3758).

After treatment of the spectra, results of each analysis were precisely relocated on microstructural pictures using the GIS software ArcMap 10.6. Such high-resolution mapping allowed ensuring the best structural link between microstructures and the degree of maturity of the CM across deformation structures using intensity ratio R1. All Raman spectra and localization on cross-sections, are available to download.

### 3.2.2. Automated analysis along line and maps

Automated acquisition of Raman spectra was carried out along lines (1D) and maps (2D) in four samples, using the objective 50×. As laser autofocus is necessary for the analyses to be carried out throughout the entire investigated domain, we had to focus the laser at the surface, and not a few microns below, as for point analysis. We checked and confirmed that the spectrum shape was not different between the surface and below (see Stokes et al. (2024) and discussion therein). For each point, exposure time was set to 10s.

Each Raman spectrum was processed according to the following steps:

1. **Baseline Correction:** A baseline is estimated and subtracted from the raw signal using the arPLS method (Baek et al., 2015), which is a penalized least-square smoothing method where the weights are uneven, with large values in purely noisy regions and closer to zero under the peaks, with parameters set to $\lambda = 10^8$ and the asymmetry ratio to $10^{-3}$. The fluorescence drift is quantified as the difference in baseline intensity between the spectrum endpoints (i.e., the minimal and maximal wavelengths).

2. **Smoothing:** The baseline-corrected signal is smoothed using the Whittaker filter (Eilers, 2003), a smoother that minimizes deviations from the data and maximizes smoothness at the same time, the balance between the two being controlled by the parameter $\lambda = 10^2$.

3. **Peak and Trough Detection:** Two maxima are identified as the highest intensity values near the expected Raman shift positions (~1350 cm⁻¹ and ~1600 cm⁻¹). A minimum is located at the lowest intensity between these peaks.

<figure><img src="image_6.png" /><figcaption>**Fig. 4.** Typical spectrum of CM obtained using micro-Raman spectroscopy. The deconvolution processes into five Lorentzian bands is described in Lahfid et al. (2010). The intensity ratio R1 is defined as the ratio between the two maxima in the spectrum (the D band and the G s.l. band).</figcaption></figure>
4. **Quality Filtering:** The signal-to-noise ratio (SNR), defined here as the mean square of the smoothed signal divided by the mean square of the noise (i.e., the difference between the smoothed and raw spectra), is calculated for each spectrum. On any given map or line profile, a threshold is defined by the user, and all spectra with SNR lower than the threshold are discarded.

5. **Manual Curation:** The remaining spectra are visually inspected to exclude anomalous cases.

As a result of this processing, each spectrum yields eight parameters (two maxima, D and G s.l., one minimum, their respective positions, the width of D and G s.l. bands as full width at half maximum (FWHM)) characteristic of the analyzed CM and two parameters (fluorescence drift and SNR) assessing the quality of the analysis.

### 3.3. Microstructural analyses

Characterization of the microstructures of the sheared samples were performed using optical and scanning electron microscopy (OM and SEM). SEM observations were carried out using a SEM Merlin Compact Zeiss, equipped with backscatter electron (BSE), secondary electron (SE) and electron dispersive spectroscopy (EDS) Quantax detectors, at 15 kV with a working distance of 10 mm and a TESCAN MIRA 3 XMU at 15 kV with a working distance of 15 mm at ISTO-BRGM, Orléans.

### 3.4. Transmitted electron microscopy

Transmission electron microscopy was used to characterize the nanostructure of Kodiak BFR 19K014A, including High Resolution TEM imaging mode and EFTEM (Energy Filtered Transmission Electron Microscopy) maps on the carbon K-edge. These experiments were performed on a ARM200 Cold FEG microscope (JEOL Ltd) operating at 200 kV with a 12 µA emission, equipped with a double spherical aberration corrector and fitted with a GIF Quantum detector (Gatan Imaging Filter). The sample was first prepared by mechanical polishing with a tripod and inlaid diamond discs to reach a thickness of 40 µm and then thin foils were obtained by argon ion milling (Gatan PIPS).

## 4. Results

### 4.1. Carbonaceous particles

In clay-rich layers of undeformed metasediments, the Raman signal of carbonaceous matter is present over a large proportion of the surface (Fig. 5 and Supplementary Material S4), while the particles are too small to be observable in optical reflected light. TEM observations of one of the samples showed, in agreement with the ubiquitous presence of the RSCM signal, particles of size 50–100 nm disseminated between the rock-forming minerals, in this case albite and chlorite (Figs. 6 and 7). The distance between two adjacent particles is smaller than the spatial resolution of Raman analysis (~1 µm), so that the Raman spectrum integrates the signal of several individual carbonaceous particles.

### 4.2. Microstructures of deformation

**Ductile deformation:** The strained rocks from Japan and Switzerland considered in this study show features of distributed deformation, such as a pervasive foliation, defined by preferential orientation of phyllosilicates (principally mica and chlorite) alternating with elongated quartz/calcite lenses or ribbons (Fig. 8a–c and 9a, b). In addition to this distributed deformation, in the Shimanto Belt, zones of localized deformation are present in form of shear zones, characterized by (i) a much stronger shape preferred orientation of the phyllosilicates and (ii) a lower porosity compared to the surrounding domains (Fig. 8c).

**Brittle deformation:** In addition to ductile deformation microstructures, in the IFU below the Glarus thrust, we observed lenses or layers of breccia embedded or intercalated within foliated rocks (Fig. 9c–e). The breccia domains are characterized by randomly orientated phyllosilicates and a larger porosity compared to surrounding foliated domains. In the Ghost Rock Formation of Kodiak accretionary prism (Fig. 6), the Black Fault Rocks also consist of breccia, characterized principally by a large grain size decrease with respect to host rock, abundant pores with a sub-micrometer size, and a relatively weak, or absent, preferred orientation of the phyllosilicate flakes.

### 4.3. Raman spectra evolution with increasing temperature conditions in no-to low-strain rocks

Considering the thermal effect of regional metamorphism on the CM

<figure><img src="image_7.png" /></figure>

**Fig. 5.** RSCM map on undeformed sample HN204 (Shimanto Belt), showing a rather uniform field of R1. In this sample, only sedimentary banding is visible and shear bands or precipitated quartz lenses are absent. (a) and (b): optical microscopy in transmitted (a) and cross-polarized light (b), showing the location of the Raman map as a star, in the clay-rich layers of the sample. (c) RSCM mapping of R1 distribution. The average is 0.494, standard deviation is 0.0157, over 1511 analyses. White pixels correspond to Raman spectra where carbon signal was absent or too weak to be analyzed.
<figure><img src="image_8.png" /><figcaption>**Fig. 6.** Microstructures of Black Fault Rocks from the Kodiak Complex. a) The Black Fault Rock comprises numerous adjacent layers made of clasts + matrix, with variations in grain size of the clasts and the matrix, as well as in the proportion of the clasts. b) Boundary between the Black Fault Rock and the host rock cataclasite. c) The matrix of the Black Fault Rocks contains grains of albite, submicron chlorite flakes and pores, without any preferred orientation of the texture. The layer analyzed by TEM is shown as the yellow star in (a). a) reflected light photograph, b) optical microscope in transmitted light and c) BSE SEM image. b and c): sample 22KO14A. Please note that with respect to outcrop orientation (Fig. 2c), all figures here are rotated by 90°. (For interpretation of the references to colour in this figure legend, the reader is referred to the Web version of this article.)</figcaption></figure>

For all samples, the correlation shows a two-staged evolution. From 180 °C up to ~250 °C, R1 remains relatively stable around 0.5. For higher temperatures of regional metamorphism, in the range 250 °C–320 °C, R1 sharply increases, and even beyond the upper limit of the geothermometer by Lahfid et al. (2010), the correlated increase in R1 and RA1 continues, up to reach values as high as 1.6 for R1 (Fig. 10b).

evolution, RSCM analyses have been carried out on undeformed to poorly deformed samples from the Shimanto Belt and the Montagne Noire area to correlate R1 signal to RSCM paleotemperatures derived from RA1 ratio with the calibration by Lahfid et al. (2010) (Fig. 10). Both areas include indeed large-scale metamorphic gradients encompassing a large range of temperatures and RSCM signals (Montmartin et al., 2021; Raimbourg et al., 2014a, 2017).

### 4.4. Raman spectra record of strain in ductile shear zones: Shimanto case study

In all studied shear zones, there is a variable but systematic increase in intensity ratio R1 in the samples selected within the deformed structures compared to the samples selected in the surrounding poorly deformed domain (Figs. 11 and 12; Supplementary Material S3). In sample HN64, while host rock R1 is ~0.46, it increases up to 0.57 in the shear zone (i.e., a maximum difference of +23% with respect to the host rock; Fig. 11). In sample 18NOB7B-2, while host rock R1 is ~0.48, it increases up to 0.76 in the shear zone (i.e., +59% with respect to the host rock; Fig. 12).

A shear zone cutting through sample 18NOB7B was analyzed using RSCM mapping (Fig. 13 and Supplementary Material S4). In the host rock, R1 values are in average 0.495 ± 0.016 (where the standard

deviation is calculated from the distribution of R1 values across the dataset), whereas in the shear zone, R1 is increased in average up to 0.558 ± 0.039 (i.e., +13% with respect to the host rock, Table 2, Fig. 14). In addition, the width of D peak (D FWMH) is shifted from 156.9 ± 8.3 cm⁻¹ in the host rock to 155.5 ± 10.8 cm⁻¹ in the shear band. While the difference in R1 between the host rock and the shear band is larger than the standard deviation, the difference in D peak width is much smaller than the standard deviation.

### 4.5. Raman spectra record in minor fault zones: the BFR from Kodiak accretionary complex

As in the Shimanto Belt sample shown in Fig. 13, the abundance of carbonaceous particles present in the host and the fault rock enables a continuous RSCM profile to be carried out (Fig. 15). The fault rock presents several distinct breccia layers, visible by contrasts in the clast size/proportion. The boundaries between the different breccia layers are reflected by gaps in R1 profile. Along this profile, in the host rocks, R1 is 0.473 ± 0.015, while in the fault rock, R1 is systematically higher than in the host rock. Fault rock domain V shows the highest R1 values: 0.532 ± 0.012 (i.e., +12% with respect to the host rock). In addition, a systematic correlation is observed between the R1 values and the width of the D peak (D FWMH), so that the lower D FWMH are observed in the fault rocks, with variations that are larger than the standard deviation (Fig. 14).

### 4.6. Raman spectra record in the vicinity of major fault zones: the IFU below the Glarus thrust

In the highly deformed IFU, directly below the Glarus thrust,
<figure><img src="image_9.png" /><figcaption>**Fig. 7.** TEM imaging of the carbonaceous particles present in sample 19KO14A (Black Fault Rock from Kodiak Complex). The particles have a size typically 50–100 nm, and are distributed throughout the sample in between chlorite and plagioclase grains. (a): TEM micrograph, (b) carbon K-edge EFTEM compositional map, (c) HRTEM of one of the carbonaceous particles, (d) FFT of the particle shown in (c).</figcaption></figure>

deformation structures and microstructures consistently show an increase in R1 in domains with higher strain. In the Kärpfbrücke area, the sample 20GT03-A is in contact with a fault plane cutting across the Flysch units. Two brecciated lenses were identified within the foliated rock (Fig. 9e), and they both show a significant increase in R1 (Fig. 16). While R1 ratio averages 0.617 in the host-rock, it reaches up to 0.908 in the brecciated domains (+49 %).

In addition, we analyzed the evolution in Raman spectra within the Flysch below the Glarus thrust, from the contact to the thrust to ~6m away from it (Fig. 17), in the Foostock area. While the average R1 value in the host-rock is 0.59, there is minimal variation until reaching a ~0.15m thick zone in contact with the thrust. There, R1 drastically increases until reaching average value of 0.9 (i.e. +53 % increase; sample 19GT19A, in contact with the thrust). The latter sample shows a strong foliation with abundant ribbons of calcite (Fig. 9a–b) without any evidence for brittle deformation.

# 5. Discussion

## 5.1. *Nature and origin of CM*

Metapelites forming most of accretionary complexes contain

abundant particles of CM originating from debris of plants, algae or bacteria scattered within the sediments (Aoya et al., 2010; Beyssac et al., 2002a; Kouketsu et al., 2014; Lahfid et al., 2010). Amongst CM particles, those identified as vitrinite can be used to derive paleotemperature conditions, based on their optical reflectance (Sakaguchi, 1999; Sakaguchi et al., 2007). In such a case, vitrinite particles must have a size of the order or larger than 10 µm, and their morphology is visible under the microscope (Gonçalves et al., 2024). On the other hand, Raman spectroscopy enables to retrieve the properties of carbonaceous matter on particles with a much smaller size, often below the micrometer scale, in which case the morphology of the particles is inaccessible in optical microscopy. In metapelites from various terranes and conditions of low to middle-grade metamorphism, carbonaceous particles have a size of the order of tens to hundreds of nanometers (Buseck and Bo-Jun 1985). Similarly, in the series of metamorphic rocks from the Schistes Lustrés of the Alps that they used to derive their geothermometer (Beyssac et al., 2002a), Beyssac et al. (2002b) describe CM particles with an onion-like geometry and size that are of the order of tens of nanometers, i.e. far below what can be observed with microscope techniques other than TEM. Because of such submicrometer size, the carbonaceous particles used in RSCM studies are not described, neither in the reference studies calibrating the geothermometers (Kouketsu et al., 2014; Lahfid et al.,
<figure><img src="image_10.png" /><figcaption>**Fig. 8.** Microstructures of ductile deformation in the Shimanto Belt. a) Foliation is defined by the preferential orientation of the ribbons of quartz and the variations in the proportions of phyllosilicates. Shear bands, deflecting quartz lenses, are shown by red arrows. Yellow arrows point to tension cracks perpendicular to extension direction X. b) and c): Differences in the orientation of phyllosilicates, from random orientation far from shear zones (b) to strongly preferentially orientated parallel to the macroscopic foliation in the shear zones (c). (a) optical microscope and (b, c) SEM BSE pictures in sample HN64. Wm, Qtz and Chl stand for white mica, quartz and chlorite, respectively. (For interpretation of the references to colour in this figure legend, the reader is referred to the Web version of this article.)</figcaption></figure>

2010), nor in regional metamorphic studies (e.g., (Nibourel et al., 2021; Raimbourg et al., 2017; Rajić et al., 2023), for the three metamorphic Belt considered here). The present work follows the same pattern: CM particles are distributed throughout the rocks (Fig. 5), even in the absence of deformation, and where they are observed that have a size of the order of ~50–100 nm (Fig. 7), well below the resolution of optical microscopy. As a consequence, their change in morphology or size as a result of strain was not studied here and only RSCM variations are considered.

The present study relies also on the hypothesis that the CM present in deformed zone is the same as the one present elsewhere in the meta-sediment. The justification for this assumption is that CM is present everywhere in some undeformed samples, so that deformed domain did contain originally CM particles before deformation. Nonetheless, another origin is possible, related to precipitation from a metamorphic fluid. In some instances, high-T fluids have been associated to C deposition, although usually with higher grade conditions than the present work, i.e., with the T conditions of the order or above 400–500 °C (Evans

et al., 2002; Luque et al., 2009; Pasteris, 1999; Pasteris and Chou, 1998). In subduction zone contexts, Galvez et al. (2013) have shown that well-crystallized graphite precipitated in a reaction zone in Alpine Corsica at the contact between sediments and serpentine, as a result of redox reactions. Such a case study, based on isotopic evidence for abiotic C, is relevant for conditions of blueschist-facies conditions, of the order or above ~10 kbars and ~430 °C, i.e., much higher than the examples reported in here. Thermodynamical calculations have shown that a large range of C-bearing species, especially anionic species, might be present in the fluid and might result in the precipitation of abiotic reduced carbon (Sverjensky et al., 2014), but these conclusions apply only to high-pressure, i.e. typically eclogite-facies, conditions, whereas for the lower conditions relevant for the present work, C-bearing, mobile species are restricted to CH₄. Below ~300 °C, examples of precipitated carbon are much rarer. For example, in a fault zone cutting across granitic rocks devoid of carbon, Oohashi et al. (2014) interpreted the carbonaceous matter present as fluid-deposited carbon. This deposited CM has a very peculiar Raman signature, with a high-amplitude D3 (i.e.
<figure><img src="image_11.png" /><figcaption>**Fig. 9.** Microstructures of high-strain rocks in the footwall of the Glarus Thrust. a) and b): Foliation is defined by ribbons of calcite or calcite + albite and layers of white mica and chlorite. Phyllosilicate grains are themselves also parallel to the foliation. c) to e): Foliation consists in alternation of phyllosilicate-rich layers, where porosity is low and phyllosilicate grains are preferentially orientated parallel to the foliation (d) and layers of breccia, with a larger porosity, a higher concentration in oxydes and phyllosilicates randomly orientated (e). All images are SEM BSE pictures. a) and b): sample 20GT19A; c), d) and e): sample 20GT03A. Cc and Por stand for calcite and porosity, respectively.</figcaption></figure>

<figure><img src="image_12.png" /><figcaption>**Fig. 10.** a) Evolution of Raman spectra of CM particles with increasing metamorphic grade, with the concomitant increase of the Raman peak area ratio RA1 and the peak amplitude ratio R1. b) The correlative evolution of R1 and RA1 (and the T derived from RA1) follow a two-step evolution, with very little R1 evolution up to T~250 °C, then a sharp increase for higher temperatures. Temperatures were derived from RA1 using Lahfid et al. (2010), and are strictly calibrated only up to 320 °C, but the correlated increase in RA1 and R1 extends beyond the range of calibration. Note that the peak amplitudes are measured on the maxima of the raw spectrum, while the areas are determined on Lorentzian peaks obtained from deconvoluting the spectrum, following Lahfid et al. (2010). Samples from Montagne Noire area are from Montmartin et al. (2020), while the other samples, all from the Shimanto Belt in Japan, are from this study (see Table S1 in Supplementary Material).</figcaption></figure>

the saddle point between D and G bands), which is also reported in their
high-velocity friction experiments, but which is strongly different from
the Raman spectra shown in the present work in deformed microstruc-
tures. In summary, even if the CM particles analyzed here most likely are

inherited from the ones present in the undeformed host rock, one cannot
discard the possibility of precipitation from a fluid, and further carbon
isotopic studies (e.g. Galvez et al. (2013)) would be necessary to settle
this issue.
<figure><img src="image_13.png" /><figcaption>**Fig. 11.** Raman profiles along shear-zones of the sample HN64 (Shimanto Belt). a) and c): scan of the whole sample and reflected light photograph of the central shear zone, respectively. b) Two examples of endmember Raman spectra across the two profiles. d) RSCM R1 ratio across the profiles. The red domains in the sample photographs and in the RSCM profile correspond to the shear zones. (For interpretation of the references to colour in this figure legend, the reader is referred to the Web version of this article.)</figcaption></figure>

## 5.2. Tracking the maturity increase using the intensity ratio R1

The evolution of CM from immature kerogens to fully crystallized graphite is a complex and multi-stage process, whose study has benefited from the combination of techniques such as X-ray diffraction (XRD), High-resolution TEM (HRTEM) and Raman spectroscopy (Oberlin, 1984; Tuinstra and Koenig, 1970) as well as programmed temperature pyrolysis (Espitalié et al., 1977). During the first stages, heteroatoms (H, O, N, S) are expulsed and the C atoms arrange into Basic Structural Units (BSU), i.e. poly-aromatic structures less than 12 rings in diameters (Beny-Bassez and Rouzaud, 1985; Tissot and Welte, 1978). With further thermal maturation, the BSU coalesce and reorient to form stacks that progressively extend laterally, until finally reaching three-dimensional lattice of graphite (Oberlin, 1984).

Graphite, which corresponds to perfectly ordered crystallites in three dimensions, exhibits a single Raman peak at 1580 $cm^{-1}$ (G band) (Beny-Bassez and Rouzaud, 1985). In contrast to the graphite structure, the presence of defects in the 3D-structure of C atoms in CM give rise to an additional band around 1355 $cm^{-1}$ (D band; (Beny-Bassez and Rouzaud, 1985; Reich and Thomsen, 2004; Tuinstra and Koenig, 1970; Wang et al., 1990), which decreases in intensity and area (R1 and R1 ratios of the Raman spectrum, respectively) when CM tend towards graphite structure (Beny-Bassez and Rouzaud, 1985; Tuinstra and Koenig, 1970; Wopenka and Pasteris, 1993). This continuous evolution of the Raman spectra is documented in a series of CM-bearing metasedimentary rock collected across metamorphic gradients from lower greenschist-facies to amphibolite-facies or granulite-facies conditions

(Beaudoin et al., 2015; Pasteris and Wopenka, 1991; Wopenka and Pasteris, 1993). The calibration of this evolution with independently constrained paleotemperatures using conventional petrography resulted in a geothermometer based on the Raman spectra of CM, applicable in the range 330–650 °C of temperatures of metamorphism (Beyssac et al., 2002a).

The evolution in CM Raman spectra for lower conditions of metamorphism is much more complex and does not follow the trend observed for T larger than 330 °C. At temperatures below 330 °C, the R1 ratio increases with the metamorphic grade (Kouketsu et al., 2014; Lahfid et al., 2010). As a result, across the evolution from poorly organized CM to perfectly ordered graphite, the highest values in R1 are observed for T~300–350 °C, i.e. lower greenschist facies conditions of metamorphism (Muirhead et al., 2021). Considering the low metamorphic conditions in the series analyzed here (i.e. < 350 °C), we observed a similar pattern of an increase in R1 associated with an increase in metamorphic temperatures in the range ~250–320 °C (Fig. 10b). Similarly, in experiments of heating of CM-bearing metapelites under pressure, R1 increases variably, as a function of heating duration and applied temperature (Moris-Muttoni et al., 2023). Finally, in one of the two samples where a wealth of data enables statistically robust parametric correlations of be made (Fig. 14), R1 increase correlates with D width decrease, which is another proxy for increased maturity (Kouketsu et al., 2014; Stokes et al., 2024).

As a consequence, we chose in this work R1 as the parameter to track variations, across strain gradients, in the maturity of CM. If other candidates exist, such as band area ratio RA1, using R1 has the advantages
<figure><img src="image_14.png" /><figcaption>**Fig. 12.** Raman profiles along shear-zones of the sample 18NOB7B-2 (Shimanto Belt). a) and c): scan of the whole sample and reflected light photograph of the central shear zone, respectively. b) Two examples of endmember Raman spectra across the two profiles. d) RSCM R1 ratio across the profiles. The red domains in the sample photographs and in the RSCM profiles correspond to the shear zones. (For interpretation of the references to colour in this figure legend, the reader is referred to the Web version of this article.)</figcaption></figure>

<figure><img src="image_15.png" /><figcaption>**Fig. 13.** RSCM map on one shear zone cutting across sample 18NOB7B-A (Shimanto Belt), showing higher R1 values in the shear zone compared to outside of it. The shear zone is shown by red arrows in (a). (a and b) optical microscopy image in transmitted (a) and reflected (b) light. (c) RSCM-derived R1 map of the area shown in (b). White pixels correspond to Raman spectra where carbon signal was absent or too weak to be analyzed. (For interpretation of the references to colour in this figure legend, the reader is referred to the Web version of this article.)</figcaption></figure>
<table><caption>**Table 2**
RSCM parameters from line profile and maps shown in Figs. 15 and 13, respectively.</caption><thead><tr><th>19KO14A</th><th>D/G</th><th>standard deviation</th><th>D peak position (cm-1)</th><th>standard deviation (cm-1)</th><th>D width (cm-1)</th><th>standard deviation (cm-1)</th><th>G peak position (cm-1)</th><th>Standard deviation (cm-1)</th><th>G width (cm-1)</th><th>Standard deviation (cm-1)</th><th>number of analyses</th></tr></thead><tbody><tr><td>Fault rock - zone V</td><td>0.532</td><td>0.012</td><td>1337.8</td><td>2.2</td><td>142.1</td><td>5.3</td><td>1605.6</td><td>1.1</td><td>46.1</td><td>1.3</td><td>4981</td></tr><tr><td>Fault rock - zone IV</td><td>0.526</td><td>0.014</td><td>1338.4</td><td>2.3</td><td>145.3</td><td>6.5</td><td>1605.7</td><td>0.8</td><td>46.4</td><td>1.1</td><td>544</td></tr><tr><td>Fault rock - zone III</td><td>0.517</td><td>0.012</td><td>1338.2</td><td>2.1</td><td>148.3</td><td>5.6</td><td>1605.7</td><td>0.7</td><td>46.1</td><td>1.2</td><td>2744</td></tr><tr><td>Fault rock - zone II</td><td>0.523</td><td>0.015</td><td>1339.3</td><td>2.7</td><td>146.6</td><td>7.3</td><td>1605.7</td><td>0.8</td><td>46.7</td><td>1.2</td><td>3525</td></tr><tr><td>Fault rock - zone I</td><td>0.494</td><td>0.017</td><td>1341.1</td><td>2.8</td><td>160.1</td><td>7.1</td><td>1605.9</td><td>0.5</td><td>46.9</td><td>1.1</td><td>554</td></tr><tr><td>host rock</td><td>0.473</td><td>0.015</td><td>1343.4</td><td>4.5</td><td>170.2</td><td>10.5</td><td>1606.1</td><td>0.7</td><td>46.7</td><td>1.5</td><td>5637</td></tr><tr><td><strong>18NOB07B-A</strong></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Shear band</td><td>0.558</td><td>0.039</td><td>1342.3</td><td>4.8</td><td>155.5</td><td>10.8</td><td>1606.1</td><td>0.5</td><td>51.8</td><td>4.0</td><td>114</td></tr><tr><td>Host rock</td><td>0.495</td><td>0.016</td><td>1339.5</td><td>2.5</td><td>156.9</td><td>8.3</td><td>1606.1</td><td>0.2</td><td>46.7</td><td>1.1</td><td>308</td></tr></tbody></table>

<figure><img src="image_16.png" /><figcaption>**Fig. 14.** Distribution of RSCM parameters (D FWMH vs. R1) as a function of microstructural domains, derived from the map in Fig. 13 (sample 18NOB07B-A) and the line profile in Fig. 15 (19KO14A).</figcaption></figure>

that (i) it can be computed from the raw spectrum without peak deconvolution and (ii) it was widely used in previous works, including our own (see Moris-Muttoni et al. (2023) and references therein), to study the correlations between deformation and RSCM parameters.

The D band increase upon maturity increase in the metamorphic T range from ~250 to ~350 °C is rather counter-intuitive, as the D band is associated with disorder, compared to the perfect graphite structure (Tuinstra and Koenig, 1970; Wang et al., 1990). Besides, the effect of grain-size of CM particles on R1 is also unclear. Pyrolysis or combustion experiments on crushed and particle-size-sorted coals have shown some effect of the particle size on CM devolatilization (Inan et al., 1998; Yu et al., 2005). But in those experiments the characteristic grain size is of the order of tens of micrometers, i.e., not comparable with the material considered here where CM grains are submicron-sized, even in undeformed, reference material. Deformation experiments also hinted at a relationship between CM particle size and Raman properties (Furuichi et al., 2015), but the size variations are correlated with strain and potential slip-induced T rise, so that the sole effect of size cannot be isolated. As a result, as we lack a comprehensive dataset about CM grain size across the sample series, an effect of CM particle size onto R1 signal, and a connection with strain, cannot be excluded.

### 5.3. Deformation microstructures and their interpretation in terms of strain rate and temperature rise

The anomalies in CM Raman signal in samples from the Shimanto Belt are observed within shear bands that concentrate strain, as can be observed from the deflection in quartz lenses (Figs. 8, 11–13). These shear bands are filled with large-sized phyllosilicates (white mica and chlorite), preferentially orientated parallel to the shear band boundaries (Fig. 8c, see also Raimbourg et al. (2019)).

In similar tectonic mélanges from the Shimanto Belt in Japan, the overall rock fabrics was interpreted in terms of brittle deformation microstructures, i.e. Y-P-R features (Y = macroscopic shear plane, P = shortening/foliation plane and R=Riedel shear planes (Logan et al., 1979; Marone and Scholz, 1989)), in which strain is localized within Riedel shear bands that disrupt the competent sandstone lenses embedded within the clay-rich matrix (Kimura et al., 2012; Kitamura and Kimura, 2012; Kondo et al., 2005). In Alaska, Fisher and Byrne (1987) describe the concentration of strain within shear bands accommodated by particulate flow and mechanical rotation of phyllosilicates, aligning them parallel to the shear zone boundary. Analyzing different tectonic mélanges worldwide, Meneghini et al. (2009) denominate the localized deformation microstructures as brittle S-C structures, where C bands are geometrically and kinematically equivalent to the Riedel shear bands. The authors emphasize the syn-kinematic crystallization of chlorite and other low-grade metamorphic minerals in C bands, so that deformation occurred for peak metamorphic conditions.

The diversity of terminology used to account for the microstructures of localized deformation in the Shimanto brings somehow some confusion as to the processes of deformation, and to the strain rate that prevailed during the formation and the development of the deformation structures. Riedel shear zones often show cataclasis and grain commutation for a large range of experimental conditions of pressure and temperature (Logan, 2007; Marone and Scholz, 1989; Marti et al., 2017). No such decrease in grain-size is observed in the shear bands examined in this study (compare Fig. 8b and c), which localize strain without undergoing damaging or cataclasis, so that they would be more akin to ductile S-C-C' structures (Volpe et al., 2022). A central observation for the behavior of the shear bands described in this work is the systematic presence of metamorphic phyllosilicates formed and grown within the shear bands (Fig. 5c-e-f in Rajić et al. (2023)). Their syn-tectonic character is attested by their size, larger than phyllosilicates in the matrix (Fig. 13 in Raimbourg et al. (2019); Fig. 9 in Meneghini et al. (2009)). Syn-kinematic growth of minerals within the shear zones, and not passive rotation and realignment (as proposed in Fisher and Byrne (1987)) suggest rather slow slip rates, enabling concomitant crystal growth during shearing. Such slow rates are further supported by growth increments of syn-kinematic quartz crystals filling the shear
<figure><img src="image_17.png" /><figcaption>**Fig. 15.** RSCM line profile in sample 19KO14A (Black Fault Rock from Kodiak Complex), across different microstructural domains in the fault zone and the boundary with the host rock. A sharp contrast in R1 is visible across the latter boundary. Bottom: Optical microscopy in reflected light.</figcaption></figure>

bands, limiting each episode of slip to less than a few micrometers (Fig. 13 in Raimbourg et al. (2019)).

In contrast to the Shimanto shear bands, in the IFU below the Glarus thrust and in the Kodiak Black Fault Rock, local anomalies in CM Raman spectra are observed in coincidence with domains of breccia, characterized by small and angular clasts, abundant submicron pores and randomly orientated phyllosilicates particles (Figs. 6 and 9e). Similar microstructures, attesting cataclasis and grain comminution, have been observed in zones of localized deformation for a wide range of strain rates in experiments, from very low (~$10^{-2}$ µm/s; Moris-Muttoni et al. (2023)), to low (~$10$ µm/s; Haines et al. (2013)), up to high-slip velocity (~m/s; French et al. (2014)). In the latter type of experiments, variably high temperatures (from ~350 °C in French et al. (2014) to more than 1000 °C in Ito et al. (2017)) are achieved in the core of the clay-rich slip zone (Ito et al., 2017; Savage and Polissar, 2019; Savage et al., 2018). Peak temperature estimates are subject to a large uncertainty, but the temperature rise in itself is attested by both thermal models and various mineralogical and chemical features, such as the variations in methylphenanthrene isomers (Savage et al., 2018) or illite dehydroxylation (Ito et al., 2017).

Therefore, on the basis of their microstructures, shear bands in tectonic mélanges are unlikely to have hosted a sharp and transient rise in temperature due to rapid and localized slip, and RSCM parameters reflect solely the effect of strain. In contrast, the temperature history of breccia zones, such as the ones observed in samples in the IFU below the Glarus thrust (Fig. 9e) or in the Kodiak Black Fault Rock (Fig. 6), is less

clear and one cannot exclude the possibility of past, rapid frictional slip with transient temperature rise.

### 5.4. Anomalous Raman signature of CM in shear bands

In order to gain insights into the parameters of deformation, in particular the strain rate, several studies have focused on organic markers, including the Raman signature of carbonaceous particles (Hirono et al., 2013, 2015; Moris-Muttoni et al., 2022, 2023; Savage and Polissar, 2019; Savage et al., 2018). As highlighted in Moris-Muttoni et al. (2023), difficulty arises in cases where both temperature increase and strain localization are possibly coinciding, as they both bear a potential effect on CM maturity that is recorded by the Raman spectra.

Based on their microstructural features, the shear bands described in this work have not hosted deformation at large strain rate. Therefore, excluding the effect of local temperature increase, we observed that localized strain during deformation at a low strain rate results in an increase in the R1 ratio, for low-grade metamorphic conditions of deformation (200–320 °C). Such a strain-related increase in R1 was reproduced in laboratory experiments of deformation (Bustin et al., 1995; Moris-Muttoni et al., 2023), and is reported in several natural case studies (Ito et al., 2017; Kaneki et al., 2016; Ujiie et al., 2021). It is consistent with evidence for the effect of shear to induce CM structural ordering using vitrinite reflectance (Suchy et al., 1997) or RSCM parameters different from R1 (Stokes et al., 2024). Finally, it confirms a trend observed for higher conditions of metamorphism, where strain is
<figure><img src="image_18.png" /><figcaption>Fig. 16. RSCM R1 profiles across breccia zones and surrounding areas in sample 20GT03A (Glarus Thrust area). a) Overview of the sample. b) Two examples of endmember Raman spectra across the two profiles. c) Close-up of two profiles showing the evolution of microstructures using optical microscopy in reflected light (right). d) RSCM R1 profiles, where grey areas represent the breccia zones.</figcaption></figure>

also shown to increase CM maturity (Barzoi, 2015; Barzoi and Guy, 2002; Petroccia et al., 2022).

In contrast to this systematic trend, some natural examples of deformation zones exhibit the opposite behavior, where deformed rocks show a decrease in R1 compared to undeformed host-rocks. We propose to reappraise below such cases of R1 decrease in deformed rocks.

In Muirhead et al. (2021), two of those examples come from the Longmenshan Fault, China, which was actively studied and drilled (five drillholes WSFD-1 to 5) after the 2008 Ms 8.0 Wenchuan earthquake (Li et al., 2013). Kouketsu et al. (2017) studied surface outcrops of fault zones within the Triassic Xujiahe Formation located 2 km southwest of drillhole WFSD-1, whereas Kuo et al. (2017, 2018) studied two fault zones within the borehole WFSD-1. In both works by Kuo et al., CM particles present in the fault gouge and breccia were interpreted as originating from the Triassic Xujiahe Formation, similarly to Kouketsu et al. (2017). Difficulties arise, when analyzing penetratively deformed rocks, to define a reference, undeformed material, with the same original CM particles as the ones contained in corresponding deformed rocks. Kouketsu et al. (2017) analyzed undeformed Triassic conglomerates, but eventually discarded these rocks as a reliable reference to their deformed rocks, because (a) the conglomerates are too distant, (b) no rock fragments (i.e. clasts), were found in the fault rocks, and (c) they may contain inherited CM particles and show indeed a large range of values. Considering only variably deformed rocks, from breccia to gouge, there is a slight increase in average R1 (from 0.54 to 0.58, Table 2 in Kouketsu et al. (2017)). Furthermore, the most deformed rocks show the largest spread of R1 values, with a lower range similar to least deformed samples, but higher range that extends up to ~0.7 (Kouketsu et al., 2017).

In addition, in the two fault zones studied by Kuo et al. (2017, 2018) within the borehole WFSD-1, gouge and breccia are considered as originated from the Triassic Xujiahe Formation. But the Raman spectra from the breccia are much different, and of much higher grade than any other Raman spectra measured in the Triassic Xujiahe Formation (Airaghi et al., 2018; Kouketsu et al., 2017; Liu et al., 2016) or than vitrinite reflectance data in the same formation (Liu et al., 2018); see also a compilation of vitrinite reflectance data in Supplementary Material Table S2). On the other hand, higher-grade shales are present in the hanging wall of the main fault zone, the Yingxiu-Beichuan Fault, or in klippes of Paleozoic sediments overlying the Xujiahe Formation (Airaghi et al., 2018; Robert et al., 2010). Given these uncertainties on the origin of the CM present in the fault and in the breccia considered in Kuo et al. (2017, 2018), the conclusions about the effect of strain on CM maturity are rather disputable.

Finally, beside Longmenshan Fault samples, the other deformed samples that have lower R1 than undeformed ones in the work of Muirhead et al. (2021) come from Haut-Giffre, in the Alps. The vitrinite reflectance data for these samples are lower than 1 (Kübler et al., 1979), corresponding to the middle of the catagenesis stage of evolution (Tissot and Welte, 1978), well below 200 °C of regional metamorphic conditions (Burnham, 2019; Muirhead et al., 2020), hence out of the range of metamorphic conditions considered here.

Last, Kedar et al. (2020) also found some decrease in R1 ratio in the overturned limb of an asymmetric fold, but their data show a very large intra-sample scattering, i.e. dispersion within the population of the CM particles contained within a given sample, which limits the impact of the
<figure><img src="image_19.png" /><figcaption>**Fig. 17.** Outcrop from the Foostock area exposing the Glarus Thrust zone (20GT19). a. Field photograph and interpretation of the meter-thick Glarus Thrust deformation zone, delimited by the dotted line, indicating the location of the analyzed samples. Zoom in the contact with mylonitized material just above the contact. b. R1 Intensity ratio measurements from the 6 samples of the gradient. Grey numbers indicate the number of analyses per sample.</figcaption></figure>

average values that led to their conclusions. It appears also that the overall Raman signal is considerably low and characterized by low signal/noise ratio (Fig. 4 in Kedar et al. (2020) in some samples, which introduces some uncertainty to the evolution with strain reported.

To summarize, after carefully analyzing the case studies available in the literature, it appears that the general rule is that naturally deformed rocks within the metamorphic temperatures range of 200-320 °C show an increase in the R1 ratio compared to undeformed rocks. In some of these examples, such as the ones described in this work, the increase in R1 can be solely attributed to strain, and not to temperature increase. It should be emphasized here that the effect of strain to increase in R1 is restricted to the aforementioned range of initial maturity of CM and conditions of metamorphic temperature, anchizone to epizone (Lahfid et al., 2010; Schito et al., 2023). Samples that would have experienced lower temperature than this range, or conversely, undergoing graphitization at higher temperatures, may be influenced differently by strain (see also the discussion in Moris-Muttoni et al. (2023).

## 5.5. Anomalous Raman signature of CM in fault zones

We examined here two field examples of the relationship between strain and RSCM in fault zones, one in the BFR of Ghost Rock Formation from Kodiak, the other one in the footwall of the Glarus Thrust, Switzerland. The latter example includes the immediate vicinity of the main Glarus Thrust fault plane (Fig. 3a), as well as a secondary fault

from the damage zone in the footwall of the Glarus Thrust and few meters away from it (Fig. 3c).

Both samples from the BFR in Ghost Rock (Fig. 15) and from the secondary fault below the Glarus Thrust (Fig. 16) contain domains of breccia. In both cases, breccia zones have a R1 ratio in average higher than the host rock: In the case of the Ghost Rock BFR, R1 average 0.532 in the breccia vs 0.473 in the host rock. In the secondary fault below the Glarus Thrust, R1 values cluster around ~0.6-0.65 in the domains outside of the breccia, whereas R1 values are relatively scattered but rise up to 0.97 within the breccia.

The R1 anomaly in brecciated domains point to the effect of localized brittle deformation on RSCM, an effect that was reported both in natural and experimental rocks (e.g., Ito et al. (2017)). The interpretation of such an anomaly is not straightforward, as discussed in detail in Moris-Muttoni et al. (2022) and Moris-Muttoni et al. (2023), as the effect of localized shear heating and strain on RSCM are difficult to disentangle. Indeed, while high velocity friction experiments have demonstrated that shear heating at rapid seismic slip increases the R1 ratio (Furuichi et al., 2015; Kaneki and Hirono, 2019), slow strain-rate experiments have also shown that strain itself at constant temperature has a similar effect (Moris-Muttoni et al., 2023).

One argument raised in Moris-Muttoni et al. (2022) to discriminate the effects of localized heat and strain relied on the shape of RSCM profile. In the Ghost Rock BFR profile, the transition between host and fault rocks is sharp and occurs through two steps in R1 ratio, coinciding
with the limits between microstructural domains (points d and e in Fig. 15). The automated RSCM profile acquired here, with a 1 µm resolution, is similar to the one obtained with much fewer data points in (Moris-Muttoni et al., 2022). The stepwise evolution of R1, matching microstructural boundaries, was the main piece of evidence suggesting that localized strain, and not local temperature rise due to friction, was the driver for RSCM evolution in the fault zone. Thermal models of frictional slip coupled with kinetic model of RSCM evolution show indeed that only continuous, and not stepwise, RSCM profile, can result from slip-induced transient temperature anomalies (Moris-Muttoni et al., 2022).

On the other hand, in the sample from a secondary fault below the Glarus Thrust, the breccia does not constitute layers, but lenses, suggesting that they were reworked as a result of posterior deformation. Consequently, the boundary between the breccia and the surrounding foliated rocks (Fig. 9d and e) is not a preserved slip boundary, unlike the BFR rocks from Ghost Rock (Fig. 15), and the interpretation of the increased R1 in the breccia zone as a consequence of either localized strain or temperature anomaly, is not possible.

A different type of anomaly, not correlated with brittle microstructures such as breccia, was observed below the Glarus Thrust (Fig. 17). There, R1 increases towards the Glarus thrust, most significantly at very short distance (i.e., less than 15 cm) from the fault plane. A possible explanation for such RSCM anomaly towards the Glarus Thrust plane involves the influence of shear heating due to repeated slip events along the fault. Shear heating critically depends on the amount of heat dissipated on the fault plane during seismic slip, which is not very well constrained. Geological constraints, such as the absence of heat flux anomaly along large-scale faults (Lachenbruch et al., 1995), plead for very efficient weakening mechanisms during slip such as flash heating (Rubino et al., 2017), which reduce the heat-dissipating shear stress to the order of ~10 MPa (Noda et al., 2009). Considering the 2008 Wenchua earthquake as an extreme endmember of collisional thrust events sets upper bounds on the total slip during the Glarus thrust earthquakes as ~10m (Feng et al., 2010; Tong et al., 2010). Thus, the total heat dissipated during a single earthquake is of the order of $10^8$ J/m$^2$ of fault plane. The Raman spectra anomaly observed in Fig. 17 is only significant over a 15 cm-thick layer along the fault. Considering a rock heat capacity of 1180 J/kg/K (Cermak and Rybach, 1982) and a rock density of ~2600 kg/m$^3$, such a dissipated heat converts into an average temperature increase of the order of 200 °C. This temperature increase over a duration of at most tens of seconds (Moris-Muttoni et al., 2022) is unlikely to have any influence on CM Raman spectra, in the light of laboratory heating experiments (Moris-Muttoni et al., 2023). Furthermore, the temperature increase due to seismic slip is very heterogeneously distributed in space, i.e., maximum along the fault plane and decreasing away from it. Thermal modelling of seismic slip with similar slip amplitude (~10m) and coupled with kinetic models of CM Raman spectra evolution shows that the influence of shear heating is restricted to less than a 1 mm from the core of the slip zone (Moris-Muttoni et al., 2022). Hence, both average calculations and more refined thermal and kinetic models of seismic slip tend to show that the heat dissipated during very large earthquakes is unlikely to have a pronounced effect on CM Raman spectra of the particles contained in a 15 cm-thick zone along the fault slip zone, as observed.

Alternatively, the dm-wide band with a R1 anomaly as we observed in the footwall of the Glarus Thrust is equivalent to a very thin zone of inverted metamorphism, whose most famous large-scale example is the Main Central Thrust (MCT) of Himalaya. Although the origin of the inverted metamorphism within the MCT is disputed (Chakraborty et al., 2016; Stephenson et al., 2000), historical models of shear heating (Arita, 1983) or of conductive heating by a "hot" hanging wall (Le Fort, 1975) are not in agreement with more recent metamorphic studies (Chakraborty et al., 2016; Dasgupta et al., 2004; Stephenson et al., 2000). In the case of the Glarus thrust, there is small gap in metamorphic temperature recorded across the fault zone, as the Verrucano formation

in the hanging wall of the Thrust is characterized by slightly higher grade conditions as evidenced by illite crystallinity data (Frey, 1988; Wang et al., 1996). Frey (1988) emphasizes the fact that the metamorphism is discontinuous across the thrust, so that the motion on the fault plane postdated the record of peak-metamorphic temperature in the hanging wall. Nonetheless, burial and metamorphism in the IFU that constitute its footwall culminated during the stage of nappe stacking and of out-of-sequence thrusting such as along the Glarus thrust, around 20 Ma (Burkhard et al., 1992; Dielforder et al., 2015, 2016), before deformation stepped down into the structurally lower crystalline Aar Massif (Herwegh et al., 2020). It is therefore possible that the increase in CM maturity observed in a few-dms wide band along the thrust itself is the result of the influence of the overthrusting of slightly hotter Helvetic Nappes over the IFU. Last, samples analyzed in close vicinity to the Glarus thrust are strongly foliated (Fig. 9a) and may have localized more strain than a few meters away from the Thrust zone, although such possibility is difficult to ascertain in the absence of quantitative strain markers. In summary, the origin of the RSCM anomaly in the footwall of the Glarus thrust, if not accounted by shear heating, remains relatively ambiguous.

## 6. Conclusions

Accretionary complexes in subduction zones (e.g., the Shimanto Belt or the Kodiak Accretionary Complex) or fold-and-thrust belts in collisional zones (e.g., the IFU below the Glarus thrust) involve contrasted structures of deformation such as shear bands and fault zones, usually connected to ductile or brittle style of deformation, respectively. In this study, we focused on the potential effects of deformation on the maturity of CM for low-grade conditions of metamorphism (i.e., 200–320 °C). To describe quantitatively changes in CM maturity, we used the intensity ratio R1, one RSCM parameter, which appears as a reliable indicator of the CM lattice organization. In all types of mm-scale microstructures of deformation, shear bands and breccia, R1 is increased compared to undeformed rocks.

The interpretation of the increase in R1 in high-strain zones was discussed in the light of two potential factors, i.e., strain or temperature. While in shear bands the sole effect of strain is demonstrated by the increase in R1 in shear bands adjacent to unsheared regions at the thin section scale, in brecciated rocks the conclusions remain unclear as both strain and heat production can be localized. RSCM profiles, with stepwise variations in R1 ratio coinciding with microstructures (e.g. the BFR in Ghost Rock Formation, Fig. 15), suggest that strain is the main driver for R1 increase in deformed zone, but further work is necessary to fully support this conclusion.

A careful examination of several experimental and natural examples points to the same conclusion of an increase in R1 due to strain for CM particles deformed in the range of 200–320 °C. Given the complexity of the CM evolution from immature matter to fully crystalline graphite, we emphasize here that the strain-induced increase in R1 is only valid for the range of thermal conditions considered. Still, this temperature range corresponds to the seismogenic zone conditions, where large earthquakes coexist with slow creep. Beyond its demonstrated potential as geothermometer, RSCM can therefore prove to be a powerful tool to study conditions of slip and deformation within the seismogenic zone.

### CRediT authorship contribution statement

**Hugues Raimbourg:** Writing – review & editing, Writing – original draft, Supervision, Project administration, Investigation, Data curation, Conceptualization.
**Benjamin Moris-Muttoni:** Writing – review & editing, Writing – original draft, Methodology, Investigation, Funding acquisition, Conceptualization.
**Romain Augier:** Writing – original draft, Supervision, Investigation, Funding acquisition, Conceptualization.
**Kristijan Rajic:** Writing – original draft, Investigation.
**Ismay Vénice Akker:** Writing – review & editing, Writing – original draft.
**Aurélien Canizarès:** Investigation. **Emmanuel Le Trong:** Investigation.

Declaration of competing interest

The authors declare the following financial interests/personal re-
lationships which may be considered as potential competing interests:
Benjamin Morris-Muttoni reports financial support was provided by
National Centre for Scientific Research. If there are other authors, they
declare that they have no known competing financial interests or per-
sonal relationships that could have appeared to influence the work re-
ported in this paper.

Acknowledgments

We thank Y. Chen for advising and helping during this work. Sylvain Janiec is thanked for the preparation of the thin and thick sections, Ida Di Carlo and Patricia Benoist for the SEM imaging. We thank also Marco Herwegh and Alfons Berger for their help in preparing the field survey in Switzerland.

This work was supported by Institut National des Sciences de
l'Univers financial program through the CARMADEF (Pi Benjamin
Morris-Muttoni and Augier Romain) project funded in 2020 and 2021.
The authors also acknowledge support from both LabEx VOLTAIRE
(ANR-10-LABX-100-01) and EquipEx PLANEX (ANR-11-EQPX-0036)
projects.

**Appendix A. Supplementary data**

Supplementary data to this article can be found online at https://doi.org/10.1016/j.jsg.2025.105560.

Data availability

All the data are available in the repository recherche.data.gouv.fr
(https://doi.org/10.57745/KP23WI; https://doi.org/10.57745/IEAQF
W; https://doi.org/10.57745/KG1CMW; https://doi.org/10.57745/40
NSHG)

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