# Fabric and anisotropy of slates: From classical studies to new results

H. -R. Wenk, R. Yu, V. Cárdenes, M. A. Lopez-Sanchez, M. Sintubin

► To cite this version:

H. -R. Wenk, R. Yu, V. Cárdenes, M. A. Lopez-Sanchez, M. Sintubin. Fabric and anisotropy of slates: From classical studies to new results. Journal of Structural Geology, 2020, 138, 10.1016/j.jsg.2020.104066. insu-03661299

HAL Id: insu-03661299

https://insu.hal.science/insu-03661299v1

Submitted on 10 Aug 2025

HAL is a multi-disciplinary open access
archive for the deposit and dissemination of sci-
entific research documents, whether they are pub-
lished or not. The documents may come from
teaching and research institutions in France or
abroad, or from public or private research centers.

L'archive ouverte pluridisciplinaire HAL, est
destinée au dépôt et à la diffusion de documents
scientifiques de niveau recherche, publiés ou non,
émanant des établissements d'enseignement et de
recherche français ou étrangers, des laboratoires
publics ou privés.

<figure><img src="image_1.png" /></figure>

Distributed under a Creative Commons Attribution 4.0 International License
4

5 ¹Department of Earth and Planetary Science, University of California, Berkeley CA 94720

6 ²Xi’an Polytechnic University, Xi’an, China

7 ³Department of Geology, Oviedo University, 33005 Oviedo, Spain

8 ⁴Geosciences, University of Montpellier, 34095 Montpellier, France

9 ⁵Department of Earth and Environmental Sciences, KU Leuven, B-3001 Leuven, Belgium

10

11 Abstract

12 Slates are low-grade metamorphic argillaceous rocks which have developed an intense tectonic
13 fabric, referred to as slaty cleavage. Due to their unique properties they have long been used as
14 building materials, especially tiles for roofs and floors. The fabric, characterized by strong
15 alignment of platy phyllosilicate crystals, produces extreme anisotropy that has been of great
16 interest in structural geology: What different fabric types are observed? How do they form? Here
17 we review crystal preferred orientation (CPO) and add new information to establish a
18 comprehensive base for future studies. Slates are composed of fine-grained phyllosilicates
19 (dominantly white mica and chlorite, sometimes chloritoid), quartz and accessories. Texture
20 analyses of these polyphase rocks with high energy synchrotron X-ray diffraction and high
21 resolution EBSD reveal strong alignment of mica and chlorite, in some cases exceeding any
22 previously reported texture strength, including metals. Interestingly, the CPO of quartz, with
23 highly flattened grains (shape preferred orientation or SPO), is practically random.

24 Recrystallization microstructures observed by scanning electron microscopy suggest that the
25 phyllosilicate preferred orientation is largely caused by crystal growth under stress, and that
26 quartz was subject to pressure solution and reprecipitation, producing an anisotropic shape.

27 While phyllosilicates preferred orientation in slates is much higher than reports for shale, schists
28 and gneiss, elastic seismic anisotropy for slates is similar to shales where a significant
29 component is due to aligned porosity.
30

31 Highlights

32 Slates are low-grade metamorphic argillaceous rocks which have developed an intense tectonic
33 fabric. The fabric is characterized by strong alignment of platy phyllosilicate crystals and
34 produces extreme anisotropy. Here we review the current state of slate crystal preferred
35 orientation research, comparing classical studies with X-ray goniometry with results from new
36 methods such as synchrotron X-ray diffraction and EBSD. With these methods much stronger
37 crystal orientation and anisotropy can be documented.
38 **Fabric and Anisotropy of Slates. From Classical Studies to New Experimental**
39 **Techniques. Outline**

40 **1. Introduction**

41 **2. Geological background of slate deposits**

42 **3. Slate as building material**

43 **4. Experimental methods applied to slate fabric analysis**

44 **4.1 Microstructural analysis with optical and electron microscopy**

45 **4.2 Early studies using X-ray diffraction to investigate slate preferred orientation**

46 **4.3 Synchrotron X-rays**

47 **4.4 SEM electron backscatter diffraction**

48 **4.5 Magnetic fabric analysis**

49 **5. Early studies of slate preferred orientation**

50 **6. Survey of slate preferred orientation analyzed with synchrotron X-rays: new records for**
51 **orientation strengths**

52 **6.1 Samples**

53 **6.2 Microstructures**

54 **6.3 Preferred orientation of phyllosilicates**

55 **6.4 Preferred orientation of quartz and accessories**

56 **7. Investigation of CPO with EBSD**

57 **8. Discussion**

58 **8.1 Extreme preferred orientation of phyllosilicates compared with earlier studies.**

59 **8.2 Absence of significant CPO of quartz.**

60 **8.3 Suggestion of mechanisms leading to such extraordinary fabrics: March Jeffery model.**

61 **8.4 Recrystallization and dissolution under stress**

62 **8.5 Significance of organic components**
63 8.6 Multiple stages of deformation

64 8.7 Comparison of CPO in slates with shales and gneiss.

65 8.8 Seismic anisotropy based on CPO.

66 8.9 Different methods to measure textures

67 **9. Conclusions**

68

69
## 1. Introduction

Slates are extraordinary metamorphic rocks characterized by their unique fabric and strength.

They have been a focus in structural geology (e.g. Twiss and Moores, 2006; Marshak, 2016;

Fossen, 2018) as well as engineering (e.g. Cardenes et al., 2014). There are no other rocks that

cleave over large planar surfaces at the millimeter scale. Because of the strength, thermal

stability and chemical inertness, slates have been widely used as roof and floor tiles, as well as

school boards (e.g. Richards, 2007).

The excellent cleavage properties and the large anisotropy of elastic properties (e.g. Christensen,

1965) are clearly correlated with the alignment of phyllosilicates such as muscovite and chlorite.

A previous review focused on industrial slate production (Cardenes et al., 2014). Here we

emphasize microstructures and preferred orientation, both of crystals (CPO) and grain shapes

(SPO), as well as elastic anisotropy. First we give a brief summary about the geologic occurrence

of slates and their origin and tectonic significance. Then we discuss slate production, from

ancient Rome to modern times.

Because of the unique properties there have been early studies to investigate microstructures

characterizing slate, first with optical microscopy in the 19th century (e.g. Sorby 1853, 1856) and

later with scanning electron microscopy (SEM) and transmission electron microscopy (TEM)

(e.g., Oertel and Phakey, 1972; Knipe and White, 1977; White and Knipe, 1978). Since these

rocks are very fine-grained, optical microscopy with a Universal stage first introduced for texture

analysis by Schmidt (1925) cannot be applied to quantify preferred orientation. Early also X-ray

diffraction has been used (Sander and Sachs, 1930). Later X-ray pole figure goniometry was

applied (e.g. Oertel and Phakey, 1972). More recently advanced methods such as synchrotron X-

ray diffraction revealed unique alignment of mica in slates from the Belgian Ardennes, with

stronger alignment than recrystallized metal sheets (Wenk et al., 2019). This stimulated this

project to revisit classical studies of slate CPO and discuss it in a broader context by comparing

results from older methods with investigations of a range of slates from different environments

with new techniques, to provide a basis for better understanding slate anisotropy. Results

confirm that the intense alignment of the slate from the Ardennes is not unique but advanced

methods and data analysis are required to document this exceptional preferred orientation.
In the discussion we review the significance of phyllosilicate content on anisotropy, mechanisms that may be responsible for creating the orientation patterns, compare fabric features of slates with shales and gneiss, and explain how CPO can be used to predict elastic anisotropy. Focus in this review is on CPO, not on the tectonic history or metamorphic conditions that produced these rocks. While this is a review, we hope we can stimulate other researchers to follow up with future systematic studies of slate preferred orientation, particularly from well-defined environments, applying high energy synchrotron X-ray diffraction and electron backscatter diffraction (EBSD). Many issues remain unresolved.

107

## 2. Geological background of slate deposits

Slates are argillaceous rocks composed mainly of muscovite, chlorite and quartz. They transformed during relatively low-grade metamorphism from sedimentary shales during tectonic events. Slates display an extraordinary cleavage with perfectly planar surfaces as observed for example in Mort Slate at Bull Point, North Devon (UK) (Fig. 1). The transformation of shales to slates occurs dominantly at greenschist facies metamorphism.

Slates occur worldwide and span a wide range of ages, ranging from Paleoprotorozoic in Hebei, China (e.g. Xue, 2002; Jinbiao, 2010), Cambrian (e.g. Newfoundland, Canada, e.g. Alam et al., 2008; Argentina, e.g. Chernicoff et al., 2008; Wales, U.K. such as Penrhyn, e.g. Gwyn, 2015; Hughes et al., 2016), Ordovician (e.g. Ratcliffe et al., 2011, Cárdenes et al., 2014), Devonian (e.g. Ardennes, Belgium and France, e.g. Sintubin, 1996; Hunsrück, Germany e.g. Wagner et al., 1997), to upper Cretaceous (Liguria, Italy, e.g. Bradolini, 1988; Del Soldato and Coli, 2013). The extraordinary cleavage in slates has been of long-standing interest (e.g. Sorby, 1853, 1856; Dieterich, 1969; Siddans, 1972; Wood, 1974; Means, 1981; Weber, 1981). Cleavage is the property of rocks to break along planar structures and, in the case of slates, schists and gneiss it is related to the preferred orientation of phyllosilicates with a planar morphology (Ramsay & Huber, 1987). Publications in structural geology discuss how cleavage in slates develops during low-grade metamorphism from shales (e.g. Tullis, 1976; Wood and Oertel, 1980; Lee et al., 1986; Ishii, 1988; Merriman et al., 1990; Ho et al., 2001).

In shales cleavage is parallel to the bedding plane formed during sedimentation and compaction. During the transformation of shales to slates the sedimentary bedding plane is sometimes
preserved. However, often the "slaty cleavage" deviates significantly from the bedding plane as in a sample of slate from the Valdeorras in Northern Spain (Fig. 2). The cleavage with perfectly planar surfaces may extend over many meters (Fig. 3). But slate deposits are often heterogeneous due to compositional variations and folding as illustrated for a quarry in Valdeorras (Fig. 4).

The slaty cleavage is a secondary, tectonic fabric developing in low-grade metamorphic conditions in argillaceous rocks. It reflects the finite strain that the argillaceous rock has undergone (see Siddans, 1972; Wood, 1974; Williams, 1977 for overviews). The slaty cleavage primarily represents a flattening strain, caused by the phyllosilicates preferred orientation, with the platelets oriented perpendicular to the principal shortening direction. Slaty cleavage development is often associated with folding and develops parallel to the axial plane of the folds.

In hand specimens slaty cleavage can be very homogeneous with no other microscopic fabric features (Fig. 5a). However, often the fabric is more complex with dominal fabrics (e.g. Powell, 1979). Due to the highly planar fabric at the micron scale, multiple strain episodes can lead to kinking and microfolds (Fig. 5b). Sometimes two cleavages are present and the slate displays a pencil-like morphology (Fig. 5c). Within the cleavage plane, a stretching lineation may develop parallel to the extension direction (Fig. 5d). In this example from Hunsrück, Germany, the cleavage coincides with the original bedding plane, illustrated by a fossil (Aspidosoma tischbeinianum). Besides the stretching lineation, different types of intersection lineations can often be recognized on cleavage planes. The most common one is the bedding-cleavage intersection lineation, often parallel to the fold hinge line of associated folds. Another type of intersection lineation is caused by the development of a secondary tectonic fabric, resulting in a crenulation of the primary cleavage fabric (Fig. 5e)..

Slate belts occupy a specific position in an orogenic system, in between the non-metamorphosed, foreland fold-and-thrust belt and the higher-grade, crystalline internal domains of the orogeny. They accommodate significant shortening (up to 60%) and drive the development of the foreland fold-and-thrust belt (e.g. Le Gall, 1992). Paleo-geographically slate belts represent the vestiges of rapid synrift of sedimentary basin fills, preceding a rift-drift transition, and leaving behind the aborted rift basins on the passive continental margins (e.g. Oncken et al., 2000).

### 3. Slate as building material
159 Slate is a classic building material in European historic and modern architecture (Fig. 6). In 160 2018, the volume of slate traded on the World market was over 10 million tons, valued at over 12 161 billion US$. Of this, roofing slate represents about 10% in volume and 5% in value of World 162 trade (data: United Nations Statistics Division UNSTATS, statistical codes 6802 for natural 163 building stone, i.e. marble, granite, etc., and 6803 for roofing slate).

164 The main characteristic of slate, the slaty cleavage, is most significant from a construction 165 standpoint. Slate is easy to split into regular blocks with many applications. However, not all 166 slates have the high degree of homogeneity needed to produce roofing slate shingles, 4 or 8 mm 167 in thickness and 20x30 cm in size. Several types of metamorphic rocks are used as roofing slates, 168 but the commonest lithologies (together with their estimated production volume) are low grade 169 slate (5%) – slate (92%) – phyllite (1.5%) – mica schist (1.5%). Roofing slate is classified using 170 two parameters: color and petrology (Cárdenes et al., 2019b). Color is the result of the 171 environmental genetic conditions, which determine the occurrence of small quantities of coloring 172 minerals (black: organic matter; green: chlorite; red: hematite), while petrology refers to both the 173 lithology and microscopic arrangement of the rock.

174 Slate architecture has evolved together with the development of the manufacturing process. In 175 the Bronze age slate sheets were extracted from outcrops with little formatting, just one or two 176 holes for wooden nails (Fig. 7a, b). First evidences of the use of slate in construction date back to 177 Prehistoric times. Slate is an important element in the tumuli from the Neolithic-Chalcolithic age 178 in some parts of the Iberian Peninsula (Iglesias Ponce de León, 1995). During the Bronze and 179 Iron ages, constructions of fortified settlements (hillforts, “castros” in Spanish) were common in 180 many regions of Europe. Some examples of these slate hillforts, in a reasonable good state of 181 conservation, can be found in Galicia and Asturias (Spain). The oldest recorded example of civil 182 architecture with slate is an aristocratic residence in the region of Extremadura (Spain) from the 183 6th century BC (Iglesias Ponce de Leon, 2012). Other examples of architecture with slate are the 184 thermal baths from Lugo (Spain), dating back to the 1st century AD.

185 During the Roman Empire slate was also an important building material. There were buildings 186 with slate walls and pavements, but only occasionally slate was used in roofs. Roman roofing 187 slate shingles have been recovered in places such as Wales (Lindsay, 1974; Gwyn, 2015),
188

Leicester, England (McWhirr, 1988), Flanders, Belgium (De Clercq, 2011) and Rheinland, Germany (Wagner, 2014).

189

190

After the Romans, following civilizations continued to use slate for roofing. During the Middle Ages roofing slate was considered as an exquisite material, reserved for Kings and the Church.

191

One of the first records regarding slate quarrying is a requirement from Bishop Licinus of Anjou (Angers, France) in the 6th century to use the slate production exclusively on religious buildings (Cnudde, 1998). Later, in 1570, an ode written by a Welsh poet mentions the historic quarry of Penrhyn, where “the bard Sion Tudor asks the Dean of Bangor for a load of slates from the Caehir quarry” (Davies, 1878). During the 16th and 17th centuries there was important trading of roofing slate between both sides of the English Channel (Wagner, 1995; Jope and Dunning, 2011). The slate industry became a significant source of wealth for the areas surrounding the quarries. By 1826, the first steam engine started operating at the Easdale quarry (Scotland), and by the second half of the 19th century, all Welsh quarries were mechanized. Almost every quarry had railways to transport roofing slate crates and Welsh slate was exported to the rest of the world (Gwyn, 2015, Price and Ronck, 2019).

192

193

194

195

196

197

198

199

200

201

202

The first slate quarry in the USA was opened in 1734 at the Peach Bottom District in Pennsylvania (Stearns et al., 2000). During the second half of the 19th century, Welsh, Italian, Irish and Slovak immigration played a major role in the development of the roofing slate industry in Slate Valley (Vermont). Roofing slate was mainly used in areas near slate outcrops in the states of Vermont, New York and Pennsylvania, mainly for singular and monumental buildings. Slate quarries were also opened in Quebec (Canada) and San Luis (Argentina).

203

204

205

206

207

208

In the 20th century, the global slate industry was severely affected by two World Wars. In the US most of the slate production was diverted to fabricate switchboards for electrical installations. Due to its low electric conductivity, slate was ideal for this application and the demand for this use greatly increased.

210

211

212

213

In the second half of the 20th century, the European slate market, the biggest in the world, was affected by the advancement of slate mining in NW Spain. So far, the Spanish roofing slate industry had been limited to the historical quarries from Bernardos, N Madrid. These quarries had been working since the 16th century, but their production was reserved for royal buildings and churches (Cárdenes et al., 2019a). However, in the 1960s, coinciding with a period of

214

215

216

217
strikes and social movements in the main French production areas (Angers), the Spanish industry skyrocketed. French and German engineers had opened quarries and brought the latest mining technologies to Spain, such as electric saws and diamond wire cutting, though the final cleaving of tiles is still done with chisel and hammer (Fig. 8). The Valdeorras region transformed to a blooming slate industry. In a few years, the center of global roofing slate industry moved to this Spanish remote region, which came to manufacture up to 90% of the World's roofing slate in 2000 (Cárdenes et al., 2014; Carcia-Guinea et al., 1997, 1998; Ward and Gomez-Fernandez, 2003).

Today, the Spanish hegemony is threatened by other countries. By the beginning of the 21st century, the rapid development of new trade routes has facilitated the export of slate products from Brazil (Chiodi, 2015), China (Xue, 2002; Jinbiao, 2010, Fig. 3) and India (Sharma et al., 2019). But the quality of these roofing slates does not always meet the requirements by clients, especially the demand from France and Germany for high quality black roofing slate (Cárdenes et al., 2019b). The relatively low-grade green and black slates produced in Brazil are not very suitable for roofing, while India and China are not yet able to provide a reliable supply of roofing slate of high quality. Figure 9 is an overview of slate production in 2010 with largest producers Spain, China and Brazil. In Figure 10 slate world production and consumption are highlighted.

235

## 4. Experimental methods applied to slate fabric analysis

### 4.1 Microstructural analysis with optical and electron microscopy

Early investigations of slate fabrics relied on optical microscopy (e.g. Sorby, 1853; Williams, 1972; Roy, 1978; Morris, 1981) but since slates are very fine-grained, application of both scanning (SEM) and transmission electron microscopy (TEM) provided important new information to characterize and understand microstructures (e.g. Gray, 1977; Knipe and White, 1979; Knipe, 1981; Weber, 1981). With the SEM, particularly in backscatter mode (BES), mineral morphology can be visualized (examples will be discussed in section 6.2). With additional energy-dispersive detectors chemical composition can be assessed and phases can be easily identified. With electron backscatter diffraction (EBSD) crystal orientation can be mapped (see sections 4.3 and 7).

247
## 4.2 Early studies using X-ray diffraction to investigate slate preferred orientation

Preferred orientation of minerals in rocks was traditionally measured in petrographic thin sections with an optical microscope equipped with a Universal stage, enabling rotation of the thin section (e.g. Schmidt, 1925). This is not applicable to opaque materials and also to very fine-grained rocks such as slates. Here X-ray diffraction provided a new opportunity and was originally applied to metals (Wever, 1924). Sander and Sachs (1930) applied the method to rocks and provided qualitative data for a Paleozoic slate from Tirol (Austria) in a photographic Laue diffraction image (Fig. 11, Sander, 1950, Fig 13, p. 53). Subsequently photographic diffraction images were used extensively for ore minerals (von Gehlen, 1960).

But it was only with the development of Geiger counters that quantitative X-ray texture analysis became possible (Decker et al., 1948). With a pole figure goniometer a sample is irradiated with monochromatic X-rays in reflection or transmission geometry. A detector set at the Bragg angle for a given lattice plane hkl records diffraction intensities while the sample is rotated to different orientations, covering the pole distribution of hkl relative to sample coordinates. With optical methods only quartz c-axes could be measured; with a pole figure goniometer orientations of any lattice planes could be recorded. There are complications: For example during sample tilt the irradiated area of the sample changes, peak broadening occurs and appropriate corrections need to be applied. Particularly for multiphase rocks such as slates diffraction peaks overlap and peak separation is limited. It is often difficult to accurately determine the scattering background which is very significant for strong crystal orientation patterns (some of these factors are discussed e.g. by Kocks et al., 1998, Chapter 4). The first application of pole figure texture measurements of phyllosilicate-rich rocks was reported by Oertel (1970) and Siddans (1976) gives some details. Section 5 describes results of some of these studies.

## 4.3 Synchrotron X-rays

Conventional X-ray tubes used for pole figure goniometry produce a broad beam (~1mm) of relatively low intensity and energy (e.g. wavelength 1.5 Å for Cu-radiation). In a synchrotron a highly focused intense beam of X-rays with monochromatic or continuous wavelengths can be produced of a wide range of energies. The use of two-dimensional digital detectors permits fast and reliable data acquisition. Figure 12a is a schematic display of a typical experiment. The X-
ray beam reaches the sample (in this case a ~5 mm diameter cylinder mounted on an aluminum rod, Fig. 12b) and causes diffraction. The recorded image displays Debye rings and intensity variations along the rings immediately indicate the presence of texture. Figure 13 illustrate four diffraction images of slates with different composition and fabric characteristics. In all of them there are strong intensities of basal (00l) reflections of phyllosilicates in the horizontal direction, suggesting strong alignment of (001) lattice planes perpendicular to this direction. Also all images display a more or less continuous ring for quartz 100 and 101, indicating random preferred orientation.

The use of high energy is advantageous because of good sample penetration and moderate absorption, as well as small 2θ angles. An example of a high energy beamline is 11-ID-C at the Advanced Photon Source (APS) of Argonne National Laboratory, with a monochromatic wavelength of about 0.1 Å (energy ~115 keV) which is ten times shorter than normal laboratory X-rays. Diffraction images at this beamline are recorded within seconds with a 2D image plate detector mounted about 2 m behind the sample (Fig. 12a). During data collection the sample can be translated to obtain a representative average. A single image covers a circle on the pole figure (Fig. 12c, #1). To obtain a better coverage the sample is rotated around the axis x in increments, e.g. from -90° to 90° in 15° increments of ω.

For quantitative texture information the digital X-ray diffraction images are then analyzed with the Rietveld method (Rietveld, 1969), which relies on a least-squares fit of measured diffraction information to minimize the difference between experimental diffraction data and a calculated model based on scattering background, crystal structures, volume fractions of components and crystal preferred orientation. Currently the software MAUD (Lutterotti et al., 1997, 2014; Wenk et al., 2014) is the most advanced software for Rietveld texture analysis and applications to some slate samples will be given in section 6.3.

The average over the whole azimuthal range produces a diffraction pattern similar to a powder diffraction pattern on which diffraction peaks and phases can be identified (Fig. SM 1). But for texture analysis we are interested in the azimuthal variation. For slates described in section 6 each image was integrated over 10° azimuthal sectors to obtain a total of 36 “diffraction patterns”. Stacks of all 36 patterns resulting from one image are shown in Figure 14 with grey shades representing intensities for two slate samples. The strong intensity variations of some
reflections (e.g. muscovite Mu 002) with azimuth on the diffraction image (vertical axis in Fig. 308
309 is representative of preferred orientation, which is also obvious in the diffraction images 310
(Fig. 13). The diffraction patterns are represented as function of Q. Q (Q=2π/d) is used instead of 311
2θ to make it independent of the rather arbitrary wavelength. Notice in Figure 14 that even for 312
muscovite 200 and chlorite 002 there are small intensities perpendicular to the main maximum, 313
indicating that some of the platy phyllosilicate crystals are orientated perpendicular to the slaty 314
cleavage. (This is consistent with SEM images (Fig. 18 and Fig. SM 2) and will be discussed in 315
section 6.2). ENT1 from Spain (Fig. 14a) shows very sharp diffraction peaks indicating high 316
crystallinity, EPS39A from the California foothills (Fig. 14b) displays diffuse peaks suggesting a 317
more disordered crystal structure and small grain size. While phyllosilicates are strongly 318
oriented, quartz has no significant intensity variations (e.g. Q 100), suggesting randomly oriented 319
crystal lattices. A comparison of the calculated Rietveld model (Fig. 14 top) with experimental 320
patterns (Fig. 14 bottom) indicates a close similarity, indicative of a good fit, both in intensities 321
as well as position of diffraction peaks.

322 MAUD calculates three-dimensional statistical crystallographic preferred orientation distribution
323 functions (ODF), relating crystal coordinates of mineral components to sample coordinates.
324 These were exported from MAUD and further processed with Beartex (Wenk et al., 1998) or
325 MTEX (Hielscher and Schaeben, 2008) to obtain, transform and plot pole figures which
326 represent the distribution of lattice plane poles and are represented in spherical projections
327 (section 6). The ODFs are also used to calculate elastic properties (section 8.8).

328

329 4.4. SEM Electron Backscatter Diffraction

330 Local crystal orientations can be measured with the scanning electron microscope (SEM) and
331 this technique has become very popular, both in materials and earth sciences. It was recently
332 applied to the slate variety phyllonite (Wallis et al. 2015). Interaction of the electron beam with
333 the uppermost surface layer of a polished sample produces single crystal electron backscatter
334 diffraction images (EBSD) (Fig. 15a) from which orientations at the sub-micron scale can be
335 determined (e.g. Dingley, 1981, Fig. 15b). By translating the sample or tilting the beam, the
336 surface can be scanned, usually covering a range of ~1 mm². At each position an EBSD image is
337 recorded, then the digitized image is automatically indexed and the orientation of the crystal is
338 extracted (e.g. Wright and Adams, 1992). On modern systems over 4000 measurements can be
339 performed in a second. EBSD maps contain much more information than bulk X-ray diffraction
340 images, particularly about orientation relationships between grains and misorientations, grain
341 shapes and sizes, domαιnal fabrics as well as evidence for grain bending. An example of EBSD
342 analysis on slate will be shown in section 7.

343 EBSD Orientation Imaging Microscopy (OIM) is a modern extension of earlier optical
344 measurements on geological materials (axis distribution analysis, "AVA", Sander, 1950).

345

346 **4.5. Magnetic fabric analysis**

347 Another method to identify the preferred orientation in slates is the use of the anisotropy of the
348 low field magnetic susceptibility (AMS) (Borradaile and Henry, 1997; Borradaile and Jackson,
349 2010). The AMS is related to the orientation distribution of iron-rich rock-forming minerals and
350 their intrinsic magnetic properties. In the case of slates, the magnetic fabric is dominated by
351 paramagnetic chlorite and chloritoid.

352 Already Wood et al. (1976) found in Cambrian slates of North Wales (UK) a good agreement
353 between the magnetic fabric and the phyllosilicate preferred orientation obtained by X-ray pole
354 figure goniometry. Typically the minimum principal magnetic susceptibility axis is oriented
355 perpendicular to the cleavage plane. On the other hand, the maximum principal magnetic
356 susceptibility axis is parallel to the fabric lineation (both stretching and intersection lineation).
357 However, there is often a mismatch between the magnetic susceptibility pattern and the
358 phyllosilicate fabric (e.g. Haerinck et al., 2013). This may be caused by different magnetic
359 carriers contributing to the magnetic fabric. To interpret such composite magnetic fabrics
360 correctly, it is imperative to understand the contributions of the different minerals to the bulk
361 magnetic fabric (Martin-Hernandez and Hirt, 2003, Biedermann et al. 2014).

362

363 **5. Early studies of slate preferred orientation**

364 Sander (1950) published a pole figure of (001) muscovite in phyllite from Zillertal (Austria),
365 measured with optical microscopy, and determined a maximum peak density of 11 multiples of
366 random distribution (mrd) (#109) (Fig. 16a). The unit to express pole densities in different
367 directions is to integrate the distribution over the whole sphere and normalize this to 1.0. In a 368 textured material there are directions which are larger than 1 (pole figure maxima) and regions 369 which are lower than 1. In a material with a random orientation distribution, all directions are 370 1.0. "Random distribution" is also sometimes called "uniform distribution". This normalization 371 applies both to ODF and to pole figures. However, there can be differences if individual crystals 372 are considered (as with Sander) in the integration, grain volumes (as with X-ray diffraction) or 373 grain surfaces (as with EBSD).

374 Early investigations of slate fabrics have been reviewed by Oertel (1983). It was Oertel (1970) 375 who first applied pole figure goniometry to study preferred orientation of muscovite in slates and 376 perfected the technique to provide quantitative information represented in (001) pole figures. 377 Emphasis was on slates from Wales and the Lake District in England. (001) pole figures of 378 muscovite in slaty andesite lapillar tuff of Ordovician age documented concentrations up to 8 379 mrd. Shape of lapilli revealed compaction and subsequent compressive deformation 380 perpendicular to the cleavage plane (Oertel, 1970). The pattern of preferred orientation of 381 muscovite was interpreted based on the March (1930) model that assumes alignment of rigid 382 platelets in a viscous medium (see Section 8.3).

383 Similar studies on a Cambrian slate from the Penrhyn quarry in N Wales illustrate maxima of 384 (001) poles perpendicular to the slaty cleavage, up to 16 mrd and an ellipsoidal distribution (Fig. 385 16b) (Oertel and Phakey, 1972). The slaty cleavage on which the pole figure is projected is quite 386 different from the bedding plane indicated by the dashed line. Oertel and Curtis (1972) studied 387 Carboniferous shale from Penistone, England, with (001) maxima for illite and kaolinite up to 9 388 mrd, attributed to progressive compaction in accordance with shape changes of concretions. The 389 comparison of strain observed in reduction bodies with preferred orientation of muscovite up to 390 15 mrd was the content of several studies (Tullis and Wood, 1975; Siddans, 1976; Wood et al., 391 1976; Wood and Oertel, 1980).

392 Oertel (1974) ingeniously used variations of chlorite preferred orientation in a fold with layers of 393 slate and quartzite from the Inyo Mountains in California to determine strain based on the March 394 concept (March, 1930) of rigid particles to unfold the fold. Texture patterns were recorded in 36 395 points with maxima ranging from 4.7 to 17.6 mrd.
Rumble and Oertel (1979) studied metaclastic rocks from New Hampshire where low grade
slates transform to higher grade schists. The strains suggested by shape of deformed pebbles
does not relate systematically with alignment of phyllosilicates in the matrix because of
extensive recrystallization. Another investigation on the effect of static metamorphism on slates
is fabric changes in Michigamme (N Michigan) samples (Ho et al. 2001) where preferred
orientation increases with increasing grade from 3.3 mrd to 9.8 mrd. However, with conventional
pole figure goniometry muscovite could not be separated from biotite due to peak overlaps and
biotite may dominate the high temperature fabric.

Ho et al. (1995) also explored fabric changes in the Lehigh Gap (Pennsylvania) as shale with
muscovite and chlorite transforms to slate, and slaty cleavage develops more or less
perpendicular to the original bedding plane. In shales the phyllosilicates have an (001) maximum
perpendicular to the bedding plane and in slates perpendicular to the cleavage.

Ishii (1988) describes grain growth and re-orientation of phyllosilicates during the development
of slaty cleavage in Northern Japan with strong girdle distributions and pole density maxima for
illite and chlorite perpendicular to the cleavage ranging from 2-8 mrd.

An interesting regional study uses mica preferred orientation in slates to interpret tectonic
processes in the Variscan orogeny of Central Europe (Fielitz, 1992). Oblique strike-slip faults
develop from pre-existing tectonic boundaries and are influenced by structural trends in the
basement. In the Stavelot-Venn basement inlier within the High-Ardenne slate belt in the Eifel
region (Germany), Cambro-Ordovician metasediments transformed to slates during tectonic
coaxial shortening with qualitative (001) mica pole figures displaying axisymmetric patterns
(Fig. 17). Later thrusting resulting in the Monschau shear zone, in which (001) pole figures
developed an orthorhombic symmetry, with the short axis of the pole figure pattern parallel to
the stretching lineation.

Sintubin (1994a) explored cleavage evolution in a Lower Palaeozoic basement inlier within the
High-Ardenne slate belt in Belgium, identifying high-strain zones based on the phyllosilicate
preferred orientation. Orthorhombic pole figure patterns superposed on axial symmetrical pole
figure patterns is considered indicative of a non-coaxial shear in high-strain zones affecting a
pre-existing strong slaty cleavage. The highest pole figure maximum for (001) muscovite is 19,5
mrd and for chlorite is 20 mrd.
426 A texture analysis of Paleozoic and Mesozoic shales from boreholes in the Campine basin
427 (Belgium) shows that the preferred orientation development in shales is independent of the burial
428 depth (Sintubin, 1994b), indicating that during the early stages of burial a stable shale fabric
429 develops, which will stay unchanged until a tectonic deformation at low-grade metamorphic
430 conditions leads to the development of a slate fabric.

431 Sintubin (1996) tracked the cleavage evolution across the frontal fold-and-thrust belt of the
432 Variscan orogen in the Belgian Ardennes by means of phyllosilicate textures, indicating that in
433 low-grade metamorphic deformation conditions rock composition and local strain conditions are
434 controlling the phyllosilicate preferred orientation. The highest pole figure maximum recorded in
435 this survey is 12 mrd for (001) muscovite.

436 Recently Haerinck et al. (2015) observed a very high degree of anisotropy of magnetic
437 susceptibility (AMS) (1.41) in Paleozoic muscovite-chloritoid slates from Brittany (France). This
438 was a first application of synchrotron diffraction to slates and documented a (001) pole figure
439 maximum of 21 mrd of chloritoid, explaining the high AMS anisotropy. Muscovite in this
440 sample had a (001) maximum of 39 mrd, two times larger than any previously recorded
441 phyllosilicate texture strength recorded with pole-figure goniometry. It lead to a follow-up study
442 of slate from the Ardennes that set a new record with 123 mrd (Wenk et al., 2019), and
443 encouraged us to use the method to survey crystal preferred orientation in a wide range of slates
444 and results will be reported in the next section.

445
446 6. Survey of slate preferred orientation analyzed with synchrotron X-rays: new records for
447 orientation strengths

448 6.1 *Samples*

449 In this section we provide results about slate fabrics investigated with synchrotron X-ray
diffraction to explore orientation patterns observed in a variety of geological settings. Samples
450 were collected more or less at random from different parts of the world (Table 1) and are rather
451 arbitrarily divided into three groups of six to represent pole figures on three pages. As mentioned
452 earlier in this context we emphasize preferred orientation and not details about the tectonic or
453 metamorphic histories.
454
A first set of slates is from Devonian slates in Belgium and Germany: the La Fortelle quarry in the Belgian Ardennes (Be-E1 and Be-E2; e.g. Asselberghs, 1924; Schavemaker et al., 2012), at Ruhrsee, Eifel, Germany (Be-A; e.g. Van Noten et al., 2012), a sample with fossils from the classical Sturtz collection from the lower Devonian Hunsrück formation at Bundenbach, Germany (Sturtz85; e.g. Briggs et al., 1996). One sample is a Silurian slate of the Plougastel Formation of Brittany, France (Sin2; Van Noorden et al., 2007) and one from a quarry of Moravian-Silesian slate deposits at Domasov in the Czech Republic (Cz-SLA1; e.g. Zimak et al., 2002).

In the second set of samples is a metasedimentary greenschist slate at the base of the Variscan Grevasalvas granite in the Austroalpine Err/Carungas nappe (Brg1806e; Cornelius, 1950; Handy et al., 1996). Two samples are from slate quarries in Northern Spain (ENT1, and RZ1, Cárdenes et al., 2014, 2015). Three samples are from China, one collected on floors in Pudong/Shanghai (Ch1) and two from Huangshan City (former Tunxi) in southern Anhui Province (Ch2a, Ch2b). They are probably of Ordovician age and were quarried in Jiangxi Province near Xingzi (e.g. Jinbiao, 2010; Xue, 2002).

The remaining six samples are from Northwestern North-America. Two samples are from the Jurassic Mariposa formation near Angels Camp in the foothills of the Sierra Nevada in California (EPS39A and EPS39B; e.g. Clark, 1964; Bogen, 1984;), one from the Paleozoic Shoo Fly Group near Iron Point, Placer Co. CA (DL1; e.g. Snow and Scherer, 2006), and two slates are from Spanish Creek, N of Quincy CA (NS1 and NS3, e.g. Day et al., 1985, 2004). A Paleozoic slate from the Blade Group is from Black Butte, Pima County, Arizona (DL3).

Most samples display a perfect macroscopic planar cleavage (e.g. Be-E1, Fig. 5a). But in the vicinity of sample Be-E1 are slates that have undergone substantial secondary deformation, resulting in kinking (Fig. 5b). An exceptional sample with two cleavages is from Ruhrsee, Germany, resulting in pencil-shaped morphology (Be-A, Fig. 5c), resembling samples from the Appalachians studied by Engelder and Geiser (1979). The sample from Bundenbach contains well-preserved Aspidosoma tischbeinianum fossils indicating that in this case the cleavage coincides more or less with the sedimentary bedding plane (Sturtz85, Fig. 5d). Samples from China have a pronounced crenulation lineation (e.g. Ch1, Fig. 5e). Lineations are also well
484
485

developed in slates from Bundenbach (Fig. 5d), the Czech Republic (Cz-SLA1) and Spain (RZ1).

486
487
488
489
490
491
492
493
494
495
496
497
498
499

The main phases in all samples are white mica, chlorite and quartz (Table 2). The volume fractions of major components were determined with the Rietveld refinement as described in section 4.2. The composition of individual phases, including accessories was identified with SEM-EDS. Quartz ranges from 25-79 Vol%. Muscovite (white mica) is the dominant phyllosilicate (22-47 Vol%) and chlorite ranges from 0-42%. One sample from China (Ch1) has a significant amount of calcite (7%) and some dolomite. Sample DL3 from Arizona contains 4% of hematite, enough to include it in the texture refinement. There are some accessories as determined by SEM-EDS: Many slate samples contain albite (0-18%). Be-E1 and Be-E2 have small amounts of calcite and two generations of chlorite. Sin2, RZ1, EPS39A and DL3 contain rutile, Ch2a and Ch2b have ilmenite, and Cz-SLA1 is rich in sphene. Be-A, Ch1 and Ch2b have apatite and ENT1 has monazite. Mica in EPS39A contains some iron and in Sin2 mica is enriched in Na. Chlorite is often interlayered with mica. Sin2 and Ch1 contain Al-Fe-rich chlorite. EPS39A, EPS39B contain ~11% of illite.

500
501

## 6.2 Microstructures

502
503
504
505
506
507
508

The microstructure is best viewed with backscatter SEM-BSE images (selected images are shown in Fig. 18 and images of all samples are summarized in supplementary material Fig. SM 2). The brightness in SEM-BES images depends largely on atomic numbers: bright for high atomic numbers such as iron-rich and dark for low atomic numbers such as Si and O (e.g. Reed, 2010). On SEM-BSE images dark is generally quartz, dark grey is white mica (muscovite) and lighter grey is chlorite. Chlorite in Sin1, RZ1 and Ch1 is very bright due to a high Fe-content. Very bright colors are ore minerals such as pyrite. Black colors are pores, some introduced during sample preparation.

509
510
511
512

In many samples phyllosilicates have a fairly euhedral geometry and are not bent (Fig. 18a, c, d, f; Fig. SM 2-a2, a5, b7, b8, b9, b12, c16, c17). The slates from the Sierra Nevada foothills in California (EPS39A, EPS39B and DL1) shows a crenulation cleavage, indicative of secondary deformation, with some alteration to illite-smectite (Fig. SM 2-c13, c14, c15). The pencil
513 cleavage (Fig. 6c), caused by the superposition of two cleavages (Be-A) is exceptional, with
514 different phyllosilicate orientations (Fig. 18h, Fig. SM 2-a3).

515 Occasionally there is interlayered growth of muscovite and chlorite (e.g. Fig. 18b,c,e, h and Fig. 516 SM 2-a1 Be-A, a4 Sturtz85, b9 RZ1, b10 Ch1, c14 EPS39B and b11 Ch2a). In sample Ch2a, 517 with a macroscopic intersection lineation, there is some evidence for kinkbands with bending, 518 creating a crenulation cleavage (Fig. 18g).

519 Quartz has in most samples a flattened shape (Fig. 18) except for the quartz band in Fig. 18f with 520 a polyhedral microstructure. In most samples quartz is surrounded by phyllosilicates. From SEM 521 images an average aspect ratio was derived which ranges for quartz from 9:1 to 2:1 (Table 3), 522 compared with 20:1 to 6:1 for phyllosilicates. Accessories such as calcite, monazite, rutile and 523 titanite are fairly equiaxed.

524

525 **6.3 Preferred orientation of phyllosilicates**

526 In this section we discuss crystallographic preferred orientation (CPO). (001) and (010) pole 527 figures in equal area projection on the cleavage plane for white mica and chlorite are shown in 528 Fig. 19a-c. This orientation is optimal for phyllosilicate pole figures because the distribution of 529 the most important element – the (001) distribution – is least distorted. Additional information is 530 summarized in Tables 2 and 3. Note that for texture analysis of monoclinic crystals such as 531 muscovite and monoclinic chlorite the “first-setting” needs to be applied in MAUD, i.e. (100) is 532 the cleavage plane, instead of a more familiar “second-setting” with (001) as the cleavage plane 533 (Matthies and Wenk, 2009). But for labels of lattice planes *hkl* in text, tables and figures we use 534 the more conventional second setting. For pole figures colors are used to represent pole densities 535 and the same pole density scales are applied to all samples which makes it easy to assess the 536 wide range of patterns. Logarithmic scales are used for phyllosilicates (Fig. 19) and linear scales 537 for quartz (Fig. 20). Samples with red maxima and white minima have strongly aligned 538 phyllosilicates.

539 For muscovite the (001) pole figure maxima vary from 4 mrd (Be-A) to 123 mrd (Be-E1) and for 540 chlorite from 4 mrd (Ch1) to 71 mrd (Be-E1). Plotting the (001) maxima of muscovite versus 541 chlorite, one can observe a rough correlation (Fig. SM 3). This is also confirmed by the 542 microstructures (Fig. 18).
In samples with high CPO for muscovite and chlorite (Be-E1, Be-E2, Sturtz85, Sin2, Brg1806e, ENT1, DL1, NS1, NS3, DL3), the orientation distribution of the (001) poles has more or less axial symmetry. Also the (010) pole figures show an axial symmetry, inferring that there is no a-axis CPO present in these samples and there is a rotational freedom about the (001) pole of phyllosilicate platelets in the cleavage plane (Fig. 19a-c).

In other samples (001) pole figures show an orthorhombic symmetry with the (001) maximum spread out, in some cases even defining a girdle (e.g. Be-A in Fig. 19a or Ch2a and Ch2b in Fig. 19b). For most samples muscovite and chlorite pole figures have (001) maxima at the same position. An exception is Ch1 where the muscovite maximum is perpendicular to the cleavage plane but the chlorite maximum is significantly inclined (Fig. 19b). However, the chlorite texture is much weaker than muscovite. This is the only sample that shows a significant deviation from an orthorhombic symmetry pattern of phyllosilicate pole figures.

The (001) texture maximum can be further quantified by determining the angular width at half-maximum of the texture peak (Table 3). This is a single number for axially symmetric textures and two numbers for girdle textures with an ellipsoidal distortion. For axially symmetric ones, the width ranges from 4° to 20° (Fig. SM 3). For the orthorhombic distributions the short width (at half-maximum) ranges from 14° to 43°. Muscovite and chlorite show very similar distributions (Fig. SM 4).

## 6.4 Preferred orientation of quartz and accessories

Perhaps most surprising has been the effective absence of significant preferred orientation for quartz (Fig. 20ab), in spite of the highly flattened shape (Fig. 18, Table 3). Variations between 0.6 (light blue) and 1.4 mrd (yellow) are most likely artifacts of the Rietveld analysis with extremely oriented phyllosilicates and diffraction peaks overlapping with quartz. A relatively high CPO was observed in Be-A (with the weakest phyllosilicate texture), with c-axes preferentially aligned in the lineation direction (1.5 mrd, Fig. 20a). The maxima of positive rhombs (10-11) perpendicular to the cleavage plane could be conceivably be attributed to mechanical Dauphiné twinning (e.g. Be-E2 with a 1.5 mrd maximum and ENT1 with a maximum of 2.3 mrd, Fig. 20a) but also this is not conclusive since diffraction from positive and
negative rhombs overlap in the same diffraction peak and may be difficult to separate, given the
superposition from strongly oriented phyllosilicates.

Only one sample, Ch1, has a significant amount of calcite (7%) and here calcite does display
preferred orientation, with a (0001) maximum perpendicular to the cleavage (Fig. SM 5a). In
DL3, with 6% hematite, the texture displays a fairly strong (0001) maximum perpendicular to the
foliation (Fig. SM 5b).

578

7. Investigation of CPO with EBSD

Electron backscatter diffraction (EBSD), as introduced in section 4.4 is a technique that allows
collecting the full crystal orientation of different mineral phases in areas as large as tens of mm²
with a scanning electron microscope (SEM). The main strength of EBSD is the ability to
characterize materials by directly linking microstructure and the full crystallographic texture. For
example, by characterizing and locating grain boundary types including epitaxial relationships
and establishing links between crystallographic orientation and the grain size, shape and/or
intracystalline plastic deformation. State-of-the-art EBSD systems have spatial resolutions down
to tens of nanometers, and angular resolution within the 0.1 to 0.5° (<0.01° using HR-EBSD)
(e.g. Wilkinson et al. 2006; Wallis et al. 2016). Many SEM-EBSD systems also allow the
acquisition of EDS (Energy-dispersive X-ray spectroscopy) maps simultaneously, allowing
complementary compositional and microstructural data to be obtained. The main limitation of
the EBSD method is that it operates on surfaces and therefore inherits all the limitations inherent
to section-dependent techniques (e.g. apparent grain sizes).

In the past phyllosilicates and clays suffered from severe misindexing problems limiting the
application of EBSD in slates (Valcke et al., 2006). Nowadays, the problem of misindexation in
mica-rich rocks mainly lies in the quality of sample preparation. Recent studies proved that with
careful preparation EBSD is a suitable technique for determining phyllonite textures and
microstructures with similar composition and grain size as slates (e.g. Wallis et al., 2015).

Here we choose a sample (BEI) of a green slate from a quarry in Beira do Rio, NW Spain, to
demonstrate the applicability of EBSD. The EBSD analysis was performed at Montpellier
University, using a Camscan Crystal Probe X500FE, with an accelerating voltage of 20 kV, and
a step size (pixel length) of 0.8 µm. Resulting data were processed with MTEX v.5.2 (Hielscher
and Schaeben, 2008, Klein et al., 2010). An SEM-BE image is shown in Figure 21a. Main components are flattened quartz (dark grey), fairly equiaxed albite (darker grey), muscovite (brighter grey) and chlorite (bright grey), similar to microstructures in Figure 18. The area displayed in Figure 21b was scanned at 0.8 µm steps and a typical diffraction image for chlorite has been shown in Figure 15. Based on diffraction images and EDS data 8 phases could be identified, dominantly quartz (26%, 216036 spots), muscovite (24%, 198501 spots), albite (21%, 175386 spots), and chlorite (14%, 116312 spots). The coverage of indexed diffraction patterns is 87% (109888 spots could not be indexed). We can immediately see the distribution of quartz (blue) with flattened grains (aspect ratio ~ 1:3) and also the distribution of muscovite (orange) clustered in domains.

Pole figures for (001) and (010) of muscovite, and (0001) and (10-11) of quartz are shown in Figure 22. Muscovite has an (001) maximum of 51 mrd, chlorite of 58 mrd, averaging all measured orientations over all grains. Pole figures for quartz display basically a random orientation distribution, similar to X-ray diffraction measurements. There is no apparent maximum for 10-11 perpendicular to the cleavage.

617

## 8. Discussion

### 8.1, Extreme preferred orientation of phyllosilicates compared with earlier studies.

Applying new quantitative methods of texture analysis based on high energy synchrotron X-ray diffraction developed for shales (e.g. Wenk et al., 2014) and SEM-EBSD for phyllonites (Wallis et al., 2015) revealed extreme preferred orientation of mica in slates, much stronger than reported in previous investigations using a pole-figure goniometer. For some samples it exceeds any reported preferred orientation in rocks and is even higher than preferred orientation in recrystallized metals that was the previous record (e.g. Hutchinson and Nes, 1992). This is extraordinary, since slates are fine-grained polyphase materials.

With all the interest in slates, slaty cleavage and slate properties it is surprising that there is not more quantitative information about crystal preferred orientation. Early studies with X-ray pole-figure goniometers were described in section 5. In investigations of slates from Wales (001) maxima of muscovite range from 2 to 16 mrd (Oertel, 1970; Oertel and Phakey; 1972, Tullis and Wood, 1975; Siddans, 1976; Wood and Oertel, 1980), in California slates 5-18 mrd (Oertel,
632 1974), in Michigamme slates from 3 to 10 mrd (Ho, 2001) and in Belgian slates from 3 to 20 mrd (Sintubin, 1994a; 1996). With synchrotron diffraction and Rietveld image analysis we find a range from 4 to 123 mrd, with 8 samples exceeding 30 mrd. The main reason probably is that previous studies relied on pole-figure goniometry which has limitations for fine-grained polyphase materials with very strong textures, particularly overlapping diffraction peaks, limitations in intensity corrections and difficulties in estimating the background.

633
634
635
636
637
The EBSD analysis of a sample from N Spain (Fig. 22) is remarkably similar to synchrotron X-ray analysis of Spanish slates (ENT1 and RZ1), with maxima ranging from 30-40 mrd (Fig. 19a).
638
639
Figure 23 is a plot of quartz volume fraction versus (001) muscovite texture strength. It demonstrates that there is no direct relationship. Samples with a very high phyllosilicate content can have strong textures (e.g. Sin2 with 65% phyllosilicates and an (001) maximum of 36 mrd for muscovite) but low phyllosilicate content can also produce strong textures (e.g. NS1 with only 21% phyllosilicates and an (001) maximum of 63 mrd). Other factors that have not been considered in this study, such as the tectonic history and metamorphic grade that may be related to CPO strength and pole distributions. Samples with strong lineations and crenulations display girdle-like textures.
640
641
642
643
644
645
646
647

648
649
650
651
652
The strength of muscovite preferred orientation correlates with the texture peak width (at half maximum) that ranges from 4 to 35° (Table 3, Fig. SM 3). Muscovite and chlorite texture maxima are correlated (Table 2, Fig. SM 4), though for weak preferred orientation chlorite may be stronger than muscovite but this should be explored further with well-defined samples from similar localities.
653
654

654 8.2. Absence of significant CPO of quartz.

655
656
657
658
659
660
661
Quite surprising has been the effective absence of significant crystal preferred orientation (CPO) for quartz (Fig. 20ab, 22), in spite of a highly flattened shape (SPO) (Fig. 18, 21, Table 3).
Concentrations of <1.5 mrd are most likely artifacts of the Rietveld analysis with extremely oriented phyllosilicates and diffraction peaks overlapping with quartz. There is no systematic orientation of positive rhombs in the foliation plane that could be attributed to mechanical Dauphiné twinning (e.g. Minor et al. 2018), except perhaps for Be-E2 and ENT1. In the literature of slates there are no data or discussion about quartz CPO in slates.
662
663 8.3. Suggestion of mechanisms leading to such extraordinary fabrics: March-Jeffery model
664 There are a few experimental studies to investigate fabric development in slates and slate analogs
665 (e.g. Means and Paterson, 1966; Paterson and Weiss, 1966; Tullis, 1976). But experiments are
666 difficult because they cannot reproduce the slow processes occurring at low temperature such as
667 growth and dissolution.

668 Some models have been developed to describe texture development by grain-shape-induced
669 rotation processes. Jeffery (1923) describes the dynamic orientation behavior of rigid, non-
670 equiaxed, particles in a deforming viscous medium. March (1932) uses this concept to infer
671 strain from the orientation pattern of planar, passive markers, deforming homogeneously with the
672 viscous medium. Tullis (1976) showed that the Jeffery model becomes identical to the March
673 model when the rigid particles attain an infinite oblate morphology.

674 The March model relates the phyllosilicate texture (pole density) with the finite strain ellipsoid,
675 following the equation:

$$
676 \quad \varepsilon_i = \rho_i^{-1/3} - 1
$$

677 with the strain ε (= Δl/l), the pole density ρ (in mrd), and the three principal directions i (1, 2, 3).

678 Applying this simple relationship is subject to a number of contraints: (1) the orientation
679 distribution has to have at least an orthorhombic symmetry; (2) the initial orientation distribution
680 is assumed to be random; (3) the marker grains are passive geometrical markers; mechanical
681 properties of the marker grains are ignored as well as their position in the material; grain
682 interaction is excluded; (4) deformation is homogeneous; (5) the orientation distribution reflects
683 the total finite strain; (6) there is no volume change during deformation.

684 It is clear that these constraints do not apply to real materials. Oertel (1983) extensively discusses
685 the significance of the different assumptions for the application of the March model to geological
686 materials. In spite of the above-mentioned constraints, the application of the March model for
687 strain estimates in geological materials has produced impressive results in cases where
688 independent strain markers were available, such as lapilli (Oertel, 1970), reduction spots (Tullis
689 & Wood, 1975) or crinoids (Weber, 1981).
690 Considering the individual phyllosilicate grains in the slate fabric, it is obvious that their
691 orientation behavior is determined by inter- and intracrystalline processes, and is thus far from
692 Marchian. But, on the average, every phyllosilicate grains seems to show a Marchian behavior,
693 behaving like passive markers rotating in a homogeneously deforming medium. In this respect
694 the Marchian behavior complies with the original definition of a cleavage, being a planar fabric
695 produced by the statistical preferred alignment of platy fabric elements (Ramsay and Huber,
696 1987).

697 The Jeffery-March approach relies on a viscous medium and rigid particles which may be
698 applicable to shales where platy phyllosilicates settle in water at surface conditions (e.g. Wenk et
699 al., 2017) as well as deep oceanic environments (e.g. Hornby, 1998; Vasin et al., 2013). This
700 alignment is increased during compaction as observed in natural environments (e.g. Lonardelli et
701 al., 2007; Sintubin, 1994) as well as experiments (e.g. Clark, 1970; Means and Paterson, 1966;
702 Voltolini et al., 2009).

703

704 8.4 Recrystallization and dissolution under stress

705 The microstructures in slates with relatively euhedral muscovite crystals, occasionally
706 interlayered with chlorite in a matrix of much stronger quartz, are more compatible with a
707 mechanism that emphasizes recrystallization and dissolution under stress. Growth under stress
708 can be controlled by thermodynamic considerations (e.g. Kamb, 1959; Paterson, 1973; Ortoleva
709 et al., 1982; Lee et al., 1986; Shimizu 1992) to predict alignment during recrystallization.
710 Muscovite is elastically highly anisotropic with a stiffness parallel to the (001) plates more than
711 three times higher than perpendicular to it (e.g. Vaughan and Guggenheim, 1986). Quartz is also
712 anisotropic, not quite by a factor two, but there are three stiff {01-11} directions and three soft
713 {10-11} directions, more or less at right angles (e.g. Heyliger et al., 2002). Thus if phyllosilicates
714 recrystallize in shales, platelets would preferentially align perpendicular to the principal
715 compressive component. Microstructures indicate that muscovite and chlorite grow by replacing
716 surrounding quartz, with no indications of significant rotations during progressive growth.

717 Dissolution under stress in sedimentary and low-grade metamorphic rocks has been described by
718 many authors (e.g. Weyl, 1959; Plessmann, 1966; Durney, 1972; De Boer, 1977; Sprunt and
719 Nur, 1977) and was applied to quartz (e.g. Powell, 1972; Morris, 1981; Wright and Platt, 1982).
This dissolution is most likely responsible for the flattened quartz grain shape (e.g. Knipe and White, 1977; White and Knipe, 1978; Knipe 1979, 1980) with dissolution and reprecipitation.

Strain has been explored based on the elliptical shape of quartz in slaty rocks from the Ardennes (Mukhopadhyay, 1973), estimating a shortening of 30 to 80%. Flat shapes of quartz caused by dissolution have also been described in siliceous sediments from Svalbard, Norway (Morris, 1981). However, if these grain shapes develop by dissolution grain shapes may not be an adequate strain estimate, like spherical radiolaria deforming into ellipsoids (e.g. Galvez and Orozco, 1979; Toriumi et al., 1988).

728

## 8.5. Significance of organic components

Sample DL1, a slate from the extensive Paleozoic Shoo Fly Formation in Northern California is a special case. The selected sample has a high quartz and mica content and displays only a moderate muscovite texture (5 mrd, Fig. 19c). This sample is adjacent to layers with high carbon content. Both have excellent cleavage but the latter display no significant preferred orientation and organic carbon may also reduce preferred orientation in DL1. "Organic" slates/shales, or argillites have been described in coal deposits, but very little is known about the transition of siliceous slates into carbon-rich rocks and corresponding microstructures and textures (e.g. Robinson, 1971; Large et al., 1994; Suarez et al., 2012; Taylor, 1991).

738

## 8.6. Multiple stages of deformation

The original slate orientation pattern may have been modified by different stages of tectonic events as secondary development of a pencil cleavage fabric at Ruhrsee. Phyllosilicates are interlayered in some larger muscovite-chlorite aggregations (Be-A, Cz-SLA1 and EPS39A, Fig. 19). Some slates do contain shear zones with macroscopic kinks (e.g. Ch1 and Ch2a) and do represent crenulation cleavage (e.g. Cosgrove, 1976; Gray and Durney, 1979; Weber, 1981). Bending of phyllosilicates has been documented by TEM (e.g. White and Knipe, 1978; Knipe, 1981; van der Pluijm et al., 1998) as well as SEM (Walsh, 2007).

The deviation of (001) pole figures of phyllosilicates from axial symmetry was quantified with the ratio large/narrow width, given in Table 3. The narrowest width of the 001 peak for
749 muscovite is plotted as function of the total texture strength in Figure SM 4. The majority of the
750 samples show a nearly axial symmetrical orientation distribution (ratio between 1 and 2). Four
751 samples show a very high ratio, ranging from 2.6 to 4.3 (Be-A, Ch2a, Ch2b, NS1).

752 The (010) pole figures show a maximum at 90° to the (001) girdle, representing a zone axis
753 (Housen et al., 1993). This zone axis is not generated by an alignment of (010) but by the girdle
754 distribution of (001) poles (Sintubin, 1998). Macroscopically, the zone axis is expressed as an
755 intersection lineation or as the long axis of a pencil cleavage (Fig. 5c). This pole figure pattern is
756 a typical of transformation textures, representing a polymodal orientation distribution of the
757 phyllosilicates (Sintubin, 1994, 1996).

758

759 *8.7. Comparison of CPO in slates with shales and gneiss.*

760 There is general agreement that most slates originate from clay-rich sedimentary rocks. Shale
761 preferred orientation has been of much interest because of the significance for hydrocarbon
762 exploration. Phyllosilicate textures in shales generally are axially symmetric and form largely
763 during sedimentation and subsequent compaction. An example is Kimmeridge shale from the
764 North Sea (Fig. 24a). Typical (001) maxima of illite-muscovite in shale range from 2-6 mrd (e.g.
765 Sintubin, 1994; Lonardelli et al. 2007; Kanitpanyacharoen et al. 2012; Vasin et al. 2013, Wenk
766 and Vasin 2018). But sedimentary microstructures observed in compacted shales are generally
767 not preserved in slates. Particularly there is no waviness of phyllosilicates that is common in
768 shales (e.g. Vasin et al., 2013, their Fig. 1). In some slates the original bedding plane is preserved
769 with markers such as fossils (Fig 5d) or compositional layering oblique to the cleavage (Fig. 2).
770 Clearly the CPO in shales is much weaker than in slates.

771 Rocks of similar composition, some of which may have transformed from slates at higher
772 metamorphic grade, are schists and gneiss. There is a wide variety of schists and gneiss, in
773 composition, microstructures and metamorphic grade. They are much coarser than slates and, to
774 sample large representative volumes, textures are often investigated with neutron diffraction.
775 Some gneiss is transformed from granite and has a high content of feldspars. Gneiss has a
776 relatively low mica content. Quartz generally does not show significant SPO but has a significant
777 texture attributed to plastic deformation.
Ullemeyer et al. (2006) describe gneiss samples from the Central Alps with a mica (001) CPO up to 4.3 mrd and quartz (001) 1.7 mrd. In highly foliated Tambo gneiss, also from the Alps (001) of muscovite is 3.6 mrd (Fig. 24c), of quartz 1.6 mrd and of plagioclase 2.5 mrd (Vasin et al., 2017). In gneiss from the Outokumpu drillhole in Finland biotite CPO reaches 23 mrd and quartz 2 mrd. For mylonitic gneiss from Montalto (S-Italy) muscovite ranges from 2-4 mrd and quartz 2-3 mrd (Fazio et al., 2017). In this study results obtained with different methods, optical, EBSD and time-of-flight neutron diffraction vary considerably. In recrystallized mylonitic gneiss from Palm Canyon (S California) biotite is up to 2.5 mrd, while quartz exceeds 8 mrd. (Pehl and Wenk, 2005).

In all these gneiss samples phyllosilicate textures are relatively weak compared with slates (Fig. 24b). They deform largely by dislocation glide and associated recrystallization.

## 8.8. Seismic anisotropy based on CPO.

The strong crystal alignment in many slates is expressed in anisotropy for seismic wave propagation. Already early experiments (e.g. Christensen, 1965) describe P-wave velocities in slate from Vermont, with velocities perpendicular to the cleavage over 20% slower than in the foliation plane. This was later conformed for other samples (e.g. Godfrey et al., 2000). Figure 25 shows velocity patterns obtained by calculating anisotropic elastic properties based on mineral composition and preferred orientation distributions for samples from the Ardennes (BE-E1) and from China (Be-2a). Averaging was done with a geometric mean over all orientations and phases but not taking grain shapes or porosity into account (Vasin et al., 2013). Both samples have over 50% quartz with a random orientation distribution and yet the seismic anisotropy is very strong. Sample Be-E1 has axially symmetric velocity surfaces, similar to the pole figures (Fig. 19a). There is strong P-wave anisotropy (>20%) and strong shear-wave splitting, close to 1 km/s for shear-waves passing parallel to the cleavage plane. In sample Ch2a preferred orientation is weaker and stretched out along a girdle (Fig. 19b). Like (001) poles of phyllosilicates, velocities show an orthorhombic pattern with highest P velocities parallel to the lineation. Shear-wave splitting is about 600 m/s (Fig. 25).
Since slates and related phyllosilicate-rich rocks such as phyllites and schists are common in continental crustal belts they contribute significantly to seismic anisotropy which should be taken into account in seismic surveys (e.g. Christensen and Mooney, 1995).

Interestingly the seismic anisotropy of slates, as displayed in Figure 25 is similar to shales, even though shales have much weaker preferred orientation (e.g. Hornby, 1998; Valcke et al. 2006; Vernik and Liu, 2007; Vasin et al., 2013; Kanitpanyacharoen, W., 2015). This can largely be attributed to a significant fraction of oriented porosity in shales which is not observed in slates. Also, the volume fraction of quartz in shales is generally lower.

814

815 *8.9. Different methods to measure textures.*

In this review we have compared determinations of crystal preferred orientation using different methods. The preferred orientation in slates demonstrates that high energy X-ray diffraction (λ~0.1 Å) is an efficient method for quantitative texture analysis of fine-grained multiphase rocks with strong preferred orientation. Some details will need to be improved in the future such as smaller integration increments for images and a finer ODF grid for the Rietveld refinement. Experiments are very fast, less than 20 minutes per sample, and this could even be improved by developing automatic sample changers. It would be interesting to compare synchrotron X-ray results with high resolution EBSD on the same samples. Notice that the texture strength of phyllosilicates based on high energy X-rays and EBSD is much higher than that based on conventional pole figure goniometry (10-20 mrd) and thus – if the March model would apply to slates – strains would have been greatly underestimated.

816

817

818

819

820

821

822

823

824

825

826

827

828

829

830

831

832

833

834

835

For synchrotron diffraction analysis there are also limitations. Especially for extreme preferred orientation. In this study azimuthal segments of diffraction images were integrated over 10°, diffraction images were recorded in 15° increments, ODF resolution in the MAUD Rietveld refinement was 7.5°. This limits the resolution of the texture refinement and the real texture strength may be considerably higher. The diffraction images in Figure 13 are all of samples with strong textures but you notice that even for Be-E1 (Fig. 13a) the diffraction ring Ch002, for example, displays also some intensities perpendicular to the foliation. In (001) pole figures (Fig. 19) this is expressed by minimum intensities, which are not zero (Table 2). Another artifact could be produced in the Rietveld refinement, since the quartz 101 diffraction is superposed on
the muscovite 006 diffraction (Fig. 14 and Fig. SM 1). To avoid this complication the region containing these peaks (2θ 1.6°-1.8°) has been excluded in some refinements (the overlap could cause both an exaggerated texture for quartz and a high background component for muscovite). In future studies such factors should be taken into account.

While X-ray diffraction averages over the whole sample volume, EBSD records orientations of individual crystals that were indexed and does not record very fine-grained components, thus the value of the lowest pole densities is lower for EBSD than for X-ray diffraction. Some local pole densities may be artificially high due to large grain size, e.g. for (0001) quartz in Figure 22.

844

## 9. Conclusions

Slates have an extraordinary cleavage and because of this have been used as tiles since Prehistoric times. The cleavage is caused by preferred orientation of phyllosilicates. Early CPO studies investigated preferred orientation with pole figure goniometry. Since then new methods such as high energy synchrotron X-ray diffraction and EBSD gave striking evidence of even stronger phyllosilicate preferred orientation. Crystallites in these fine-grained multiphase rocks are aligned even more perfectly than in highly deformed metals. Equally surprising is the virtual absence of preferred orientation in highly flattened quartz. This can best be attributed to growth and dissolution at relatively low-grade metamorphic conditions. The alignment of phyllosilicates produces strong anisotropy of elastic properties, relevant to seismic wave propagation. In the future the same techniques should be applied to systematic studies of exploring the transition regimes between shale and slate during diagenesis, as well as slate and schist during progressive metamorphism.
858 **Acknowledgments:** We are obliged to Timothy Horscroft for inviting us to prepare this review.

859 We acknowledged access to beamline 11-ID-C of the Advanced Photon Source and the Zeiss

860 EVO SEM at EPS, UC Berkeley. Yang Ren helped with APS experiments and Tim Teague

861 helped with sample preparation as well as SEM and XRD analyses. Luca Lutterotti, Mauro

862 Bortolotti and Roman Vasin helped with the Rietveld refinement. HRW is appreciative for

863 support from NSF (EAR 1343908) and DOE (DE-FG02-05ER15637) and from The Fulbright

864 Foundation for a fellowship in Salamanca. RY acknowledges support from IKSTP, Shaanxi,

865 China (No. 2015GY101). VC is grateful to his grant PA-18-ACB17-11, from the Program Marie

866 -Curie funded COFUND by the Government of Asturias (Spain) and the Spanish FICYT. MALS

867 was supported by the European Union and the Government of the Principality of Asturias

868 through the “Clarín-Cofund” program [grant number ACA17-32] within the Marie Skłodowska-

869 Curie COFUND (FP7) Actions. We appreciate receiving samples from Juan Gomez Barreiro

870 (Salamanca) and David Lawler (Grass Valley CA). We are greatly indebted to two reviewers,

871 Rob Knipe and Ben van der Plujim, as well as the editor Joao Hippertt for their detailed

872 comments that helped to improve the manuscript. HRW is grateful to all the contributions from

873 his co-authors, particularly to Rong Yu for her meticulous analytical work.
874 References

875 Asselberghs, E., 1924. Les ardoisières du Dévonien de l'Ardenne. Annales des Mines de Belgique, 1037-1098.

877 Alam, M.R., Swamidas, A.S.J., Gale, J. and Munaswamy, K., 2008. Mechanical and physical properties of slate from Britannia Cove, Newfoundland. Canadian Journal of Civil Engineering, 35(7), 751-755.

880 Attewell, P.B., Taylor, R.K., 1969. A microtextural interpretation of a Welsh slate. Inter. J. Rock Mech. Min. Sci. 6, 423-438.

882 Biedermann, A.R., Koch, C.B., Lorenz, W.E.A., Hirt, A.M. 2014. Low temperature magnetic anisotropy in micas and chlorite. Tectonophysics 629, 63-74.

884 Bogen, N.L., 1984. Stratigraphy and sedimentary petrology of the Upper Jurassic Mariposa Formation, Western Sierra Nevada, California. In: Croach, J.K., Bachman S.B. eds. Tectonics and Sedimentation along the California Margin. Pacific Section S.E.P.M. 38, 119-134.

888 Borradaile, G.J., Henry, B. 1997. Tectonic applications of magnetic susceptibility and its anisotropy. Earth-Science Reviews 42, 49-93.

890 Borradaile, G.J., Jackson, M. 2010. Structural geology, petrofabrics and magnetic fabrics (AMS, AARM, AIRM). J. Struct. Geol. 32, 1519-1551.

892 Brandolini, P., 1988. L'utilizzazione dell'ardesia in Liguria. Studi e ricerche di Geografia 11, 31-85.

894 Briggs, D. E. G., Raiswell, R., Bottrell, S. H., Hatfield, D., Bartels, C., 1996. Controls on the pyritization of exceptionally preserved fossils: an analysis of the Lower Devonian Hunsrück slate of Germany. Am. J. Sci. 296, 633-663.

897 Bultynck, P., Dejonghe, L., 2001. Devonian lithostratigraphic units (Belgium). Geologica Belgica 4, 39-69.

899 Cárdenes, V., Cnuddle, V., Cnuddle, J.P., 2014. Iberian roofing slate as a Global Heritage Stone Province Resource. Episodes 38 (2), 97-105.
901 Cárdenes, V., Ponce de León, M., Rodríguez, X.A. and Rubio-Ordoñez, A., 2019a. Roofing Slate
902 Industry in Spain: History, Geology, and Geoheritage. Geoheritage, 11(1): 19-34.

903 Cárdenes, V., Rubio-Ordoñez, A. and Ruiz de Argandoña, V.G., 2019b. Definition of roofing
904 slate lithotypes and their physical and petrographical characteristics. Key Materials in
905 Engineering, in press.

906 Cárdenes, V., Rubio-Ordoñez, A., Monterroso, C., Calleja, L., 2013. Geology and geochemistry
907 of Iberian roofing slates. Chem. Erde 73, 373-382. doi.org/10.1016/j.chemer.2012.11.004.

908 Cárdenes, V., Rubio-Ordonez, A., Wichert, J., Cnuddle, J.P., Cnuddle, V., 2014. Petrography of
909 roofing slates. Earth-Science Reviews 138, 435-453, doi.
910 org/10.1016/j.earscirev.2014.07.003.

911 Chernicoff, C.J., Zappettini, E.O., Santos, J.O.S., Beyer, E. and McNaughton, N.J., 2008.
912 Foreland basin deposits associated with Cuyania accretion in La Pampa Province, Argentina.
913 Gondwana Research, 13(2), 189-203.

914 Chiodi Filho, C., 2015. A vitoria das ardosias Bambui na Marca CE.

915 Cholach, P.Y., D.R. Schmitt, 2006. Intrinsic elasticity of a textured phyllosilicate aggregate:
916 relation to the seismic anisotropy of shales and schists, J. Geophys. Res., 111, B09410,
917 doi:10.1029/2005JB004158, 18 pp

918 Christensen, N.I., 1965. Compressional wave velocities in metamorphic rocks at pressures to 10
919 kbar. J. Geophys. Res. 70, 6147-6164.

920 Christensen, N.I., Mooney, W.D., 1965. Seismic velocity structure and composition of the
921 continental crust: A global view. J. Geophys. Res. 100, 9761-9788.

922 Clark, B.R., 1970. Mechanical formation of preferred orientation in clays. Am. J. Sci. 269(3),
923 250-266.

924 Clark, L.D., 1964. Stratigraphy and structure of part of the western Sierra Nevada metamorphic
925 belt, California. U.S. Geol. Survey Professional Paper 410, 70pp.

926 Cnudde, J.P., 1998. History of the Belgian slate industry. Homologation and sampling of natural
927 stone in Belgium. The part of the Applied Geology Laboratory of the University of Gent in
928 the homologation and the choice of natural stones. Natuurwetenschappelijk Tijdschrift 78(1-
929 4), 13-20.
930 Cornelius, H.P. 1950. Geologie der Err-Julier-Gruppe. II. Teil: Der Gebirgsbau. Beitr. Geol.
931 Karte Schweiz [N.F.] 70/2.

932 Cosgrove, J.W., 1976. The formation of crenulation cleavage. J. Geol. Soc. 132, 155-178

933 Davies, D.C., 1878. A Treatise on Slate and Slate Quarrying. Crosby, London, 186pp.

934 Day H.W., Bickford, M. E., 2004. Tectonic setting of the Jurassic Smartville and Slate Creek
935 complexes, northern Sierra Nevada, California. Geological Society of America Bulletin, 116,
936 11/12, 1515-1528, doi:10.1130/B25416.1

937 Day, H. W., Moores, E. M., Tuminas, A. C., 1985. Structure and tectonics of the northern Sierra
938 Nevada. Geological Society of America Bulletin, v. 96, p. 436-450.

939 De Boer, R.B., 1977. Pressure solution: theory and experiments. Tectonophysics 39, 287-301.

940 De Clercq, W., 2011. Roman rural settlements in Flanders. Perspectives on a 'non-villa'
941 landscape in extrema Galliarum. In: N. Roymans and T. Derks (Editors), Villa Landscapes in
942 the Roman North. Amsterdam University Press Amsterdam, pp. 332.

943 Del Soldato, M. and Coli, M., 2013. L'ardesia della Val Lavagna: pietra del paesaggio e della
944 cultura, 5° Congresso Nazionale Geologia e Turismo, Bologna.

945 Dieterich, J.H., 1969. Origin of slaty cleavage in folded rocks. Am. J. Sci. 267, 155-165.

946 Durney, D.W., 1972. Solution transfer, an important geological deformation mechanism. Nature
947 235, 315-317

948 Engelder T., Geiser P., 1979. The relationship between pencil cleavage and lateral shortening
949 within the Devonian section of the Appalachian Plateau, New York. Geology 7, 460-464

950 Eric, J.H., Stromquist, A.A., Swinney, C.M., 1955. Geology and mineral deposits of the Angels
951 Camp and Sonora quadrangles, Calaveras and Tuolomne Counties, California. Cal. Div.
952 Mines and Geology, Special Report 41, 55pp.

953 Fazio, E., Punturo, R., Cirrincione, R., Kern, H., Pezzino, A., Wenk, H.-R., Goswami, S.,
954 Mamtani, M.A., 2017. Quartz preferred orientation in naturally deformed mylonitic rocks
955 (Montalto Shear Zone – Italy): a comparison of results by different techniques, their
956 advantages and limitations. Int. J. Earth Sciences 106, 2259-2278. [doi:10.1007/s00531-016-
957 1424-y]
958 Fettes, D.J., Desmons, J. and Árkai, P., 2011. Metamorphic rocks: a classification and glossary
959 of terms. Recommendations of the International Union of Geological Sciences
960 Subcommission on the Systematics of Metamorphic Rocks. Cambridge University Press, 244
961 pp.

962 Fielitz, W., 1992. Variscan transpressive inversion in the northwestern central Rhenohercynian
963 belt of western Germany. J. Struct. Geology 14 (5), 547-563.

964 Fossen, H., 2018. Structural Geology, 2nd Edn, Cambridge University Press.

965 García-Guinea, J., Lombardero, M., Roberts, B., Taboada, J., 1997. Spanish roofing slate deposits.
966 Trans. Inst. Min. Metall., Sect. B 106, 205-214.

967 García-Guinea, J., Lombardero, M., Roberts, B., Peto, A., 1998. Mineralogy and microstructure
968 of roofing slate: thermo-optical behaviour and fissility. Mater. Constr. 48 (251), 37-48.

969 Godfrey, N.J., Christensen, N.I., Okaya, D.A., 2000. Anisotropy of schists: Contribution of
970 crustal anisotropy to active source seismic experiments and shear wave splitting
971 observations. J. Geoph. Res. 105 (B12), 27991-28007.

972 Gray, D.R., 1977. Some parameters which effect the morphology of crenulation cleavage. J.
973 Geol. 85, 763-780.

974 Gray D. R., Durney D. W., 1979. Crenulation cleavage differentiation: implications of solution-
975 deposition processes. J. Struct. Geol. 1, 73-80

976 Guggenheim, S., Chang, Y.-H., Koster van Groos, A.F., 1987. Muscovite dehydroxylation: high-
977 temperature studies. American Mineralogist 72, 537-550.

978 Gwyn, D., 2015. Welsh Slate: Archaeology and History of an Industry. Royal Commission n the
979 Ancient and Historical Monuments of Wales.

980 Haerinck, T., Adriaens, R., Debacker, T.N., Hirt, A.M., Sintubin, M. 2013. Paramagnetic
981 metamorphic mineral assemblages controlling AMS in low-grade deformed metasediments
982 and the implications with respect to the use of AMS as a strain marker. Journal of the
983 Geological Society, London 170, 283-280.

984 Haerinck, T., Wenk, H.-R., Debacker, T.N., Sintubin, M., 2015. Preferred mineral orientation of
985 a chloritoid-bearing slate in relation to its magnetic fabric. J. Struct. Geol. 71, 125-135,
986 doi.org/10.1016/j.jsg.2014.09.013.
987 Handy, M.R., Herwegh, M., Kamber, B.S., Tietz, R., Villa, I.M., 1996. Geochronologic, petrologic and kinematic constraints on the evolution of the Err-Platta boundary, part of a fossil continent-ocean suture in the Alps (eastern Switzerland). Schweiz. Mineral. Petrog. Mitt. 76, 453-474.

991 Heyliger, P., Ledbetter, H., Kim, S., 2002. Elastic constants of natural quartz. Journal of Acoustical Society of America 114, 644-650, doi.org/10.1121/1.1593063.

993 Hielscher, R., Schaeben, H., 2008. A novel pole figure inversion method: specification of the MTEX algorithm. J. Appl. Cryst. 41, 1024-1037.

995 Ho, N.-C., Peacor, D.R., van der Pluijm, B.A., 1996. Contrasting roles of detrital and authigenic phyllosilicates during slaty cleavage development. Journal of Structural Geology 18, 615-623.

998 Ho, N.-C., Peacor, D.R., van der Pluijm, B.A., 1995. Reorientation mechanisms of phyllosilicates in the mudstone-to-slate transition at Lehigh Gap, Pennsylvania. Journal of Structural Geology 17, 345-356.

1001 Ho, N.-C., van der Pluijm, B.A., Peacor, D.R., 2001. Static recrystallization and preferred orientation of phyllosilicates: Michigamme Formation, northern Michigan, USA. Journal of Structural Geology 23, 887-893.

1004 Hornby, B.E., 1998. Experimental laboratory determination of the dynamic elastic properties of wet, drained shales, Journal of Geophysical Research 103, 29945-29964.

1006 Housen, B.A., Richter, C., van der Pluijm, B.A., 1993. Composite magnetic anisotropy fabrics: experiments, numerical models, and implications for the quantification of rock fabrics. Tectonophysics 220, 1-12.

1009 Hughes, T., Horak, J., Lott, G. and Roberts, D., 2016. Cambrian age Welsh Slate: A global heritage stone resource from the United Kingdom. Episodes, 39(1), 45-52.

1011 Hunt, B., 1998. Chinese Slate—Telling the good from the bad. Roofing, Cladding and Insulation, October 1998. pp. 54–56.

1013 Hunt, B., 2006. Building stones explained—slate. Build. Stones 22 (1), 33–40.
1014 Hutchinson, W.B., Nes, E., 1992. Texture development during grain growth – a useful rule-of-thumb. Proc. Grain Growth in Polycrystalline Materials I. Materials Science Forum 94-96, 385-390.

1017 Iglesias Ponce de León, M., 1995. L'ardoise en Espagne: Histoire et économie, Université de Rennes 2, Rennes.

1019 Ishii, K., 1988. Grain growth and re-orientation of phyllosilicate minerals during the development of slaty cleavage in the South Kitakami Mountains, northeast Japan. J. Struct. Geol. 10, 145-154.

1022 Jeffery, G.B., 1923. The motion of ellipsoidal particles immersed in a viscous fluid. Proc. R. Soc. London 102, 161-179.

1024 Jinbiao, S., 2010. The characteristics of black roofing slate deposit in Western Hubei. Chin. Non-metallic Min. Ind. Her. 2, 52-53.

1026 Jope, E.M. and Dunning, G.C., 2011. The use of blue slate for roofing in medieval England. The Antiquaries Journal 34(3-4): 209-217. doi:10.1017/s0003581500059886.

1028 Kamb, W.B., 1959. Theory of preferred crystal orientation developed by crystallization under stress. Journal of Geology 67, 153-170.

1030 Kanitpanyacharoen, W., Kets, F.B., Wenk, H.-R., Wirth, R., 2012. Preferred orientation, microstructures and porosity analyses of Posidonia shales. Clays and Clay Minerals 60, 315-329, doi:10.1346/CCMN.2012.0600308.

1033 Kanitpanyacharoen, W., Parkinson, D.Y., de Carlo, F., Marone, F., Wenk, H.-R., MacDowell, A., Mokso, R., and Stampanoni, M., 2013. A comparative study of X-ray microtomography on shales at different synchrotron facilities: ALS, APS and SLS. J. Synchrotron Radiation 20, 1-9 [doi:10.1107/S0909049512044354].

1037 Klein, H., Hielscher, R., Schaeben, H., 2010. Texture analysis with MTEX – Free and open software toolbox. Solid State Phenomena 160, 63-68.

1039 Knipe, R.J., White, S.H., 1977. Microstructural variation of an axial plane cleavage around a fold. A H.V.E.M. study. Tectonophysics 39, 355-381.

1041 Knipe, R.J., 1979. Chemical changes during slaty cleavage development. Bull. Mineral. 102, 206-209.
1043 Knipe, R.J., 1981. The interaction of deformation and metamorphism in slates. Tectonophysics
1044 78, 249-272.

1045 Kocks, U.F., Tomé, C., Wenk, H.-R., 2000. Texture and Anisotropy. Preferred Orientations in
1046 Polycrystals and Their Effect on Materials Properties. 2nd paperback edition. Cambridge
1047 University Press, 676pp.

1048 Large, D.J., Christy, A.G., Fallick, A.E., 1994. Poorly crystalline carbonaceous matter in high
1049 grade metasediments: implications for graphitisation and metamorphic fluid composition.
1050 Contrib. Mineral. Petrog. 116, 108-116.

1051 Lee, J.H., Peacor, D.R., Lewis, D.D., Wintsch, R.P., 1986. Evidence for syntectonic
1052 crystallization for the mudstone to slate transition at Lehigh Gap, Pennsylvania, USA.
1053 Journal of Structural Geology 8, 767-780.

1054 Le Gall, B., 1992. The deep structure of the Ardennes Variscan thrust belt from structural and
1055 ECORS seismic data. Journal of Structural Geology 14, 531-546.

1056 Lin, W., Faure, M., Monie, P., Schaerer, U., Zhang, L., Sun, Y., 2000. Tectonics of SE China:
1057 New insights from the Lushan massif (Jiangxi Province). Tectonics 19, 852-871.

1058 Lindsay, J., 1974. A History of the North Wales Slate Industry. David and Charles, Newton
1059 Abbott. ISBN 0-7153 6264-X.

1060 Lipshie, S.R., Oertel, G., Christie, J.M., 1976. Measurement of preferred orientation of
1061 phyllosilicates in schists. Tectonophysics 34, 91-99, doi.org/10.1016/0040-1951(76)90178-5.

1062 Lonardelli I., Wenk, H.-R., Ren, Y., 2007. Preferred orientation and elastic anisotropy in shales.
1063 Geophysics 72, D33-D40, doi:10.1190/1.2435966.

1064 Lutterotti, L., Matthies, S., Wenk, H.-R., Schultz, A.J., Richardson, J.W., 1997. Combined
1065 texture and structure analysis of deformed limestone from time-of-flight neutron diffraction
1066 spectra. J. Appl. Phys. 81, 594-600.

1067 Lutterotti, L., Vasin, R., Wenk, H.-R., 2014. Rietveld texture analysis from synchrotron
1068 diffraction images. I. Calibration and basic analysis. Powder Diffraction 29, 76-84,
1069 doi.org/10.1017/S0885715613001346.

1070 March, A. 1932. Mathematische Theorie der Regelung nach der Korngestalt bei affiner
1071 Deformation. Zeitschrift für Kristallographie 81, 285-297.
1072 Martin-Hernandez, E., Hirt, A.M. 2003. The anisotropy of magnetic susceptibility in biotite,
1073 muscovite and chlorite single crystals. Tectonophysics 367, 13-28.

1074 Matthies, S., Vinel, G.W., 1982. On the reproduction of the orientation distribution function of
1075 texturized samples from reduced pole figures using the conception of a conditional ghost
1076 correction. Physica Status Solidi (b) 112, K111-114.

1077 Matthies, S., Wenk, H.-R., 2009. Transformations for monoclinic crystal symmetry in texture
1078 analysis. Journal of Applied Crystallography 42, 564-571,
1079 doi.org/10.1107/S0021889809018172.

1080 McWhirr, A., 1988. The Roman Swithland slate industry. Transactions of the Leicester
1081 Archaeological & Historical Society: 1-8.

1082 Means, W.D., 1981. The concept of steady-state foliation. Tectonophysics 78(1-4), 179-199.

1083 Means, W.D., Paterson, M.S., 1966. Experiments on preferred orientation of platy minerals.
1084 Contrib. Mineral. Petrol. 13(2), 108-133.

1085 Merriman, R.J., Roberts, B., Peacor, D.R., 1990. A transmission electron microscope study of
1086 white mica crystallite size distribution in a mudstone to slate transitional sequence. Contrib.
1087 Mineral. Petrol. 106, 27-40.

1088 Mookherjee, M., Mainprice, D., 2014. Unusually large shear wave anisotropy for chlorite in
1089 subduction zone settings. Geophys. Res. Lett. 41, 1506-1513. doi:10.1002/2014GL059334

1090 Morris, A.P., 1981. Competing deformation mechanisms and slaty cleavage in deformed,
1091 quartzose meta-sediments. J. Geol. Soc. London 138(4), 455-462.

1092 Mukhopadhyay, D., 1973. Strain measurements from deformed quartz grains in slaty rocks from
1093 the Ardennes and the northern Eifel. Tectonophysics 16(3-4), 279-296.

1094 Oertel, G., 1970. Deformation of a slaty, lapillar tuff in the Lake District, England. Geol. Soc.
1095 Am. Bull. 81, 1173-1188.

1096 Oertel, G., 1983. The relationship of strain and preferred orientation of phyllosilicate grains in
1097 rocks--a review. Tectonophysics 100, 413-447, doi.org/10.1016/0040-1951(83)90197-X.

1098 Oertel, G., Curtis, C. D., 1972. Clay ironstone concretion preserving fabrics due to progressive
1099 compaction: Geol. Soc. America Bull. 83, 2597-2606.
1100 Oertel, G., Curtis, C. D., Phakey, P. P., 1973. A transmission electron microscope and x-ray
diffraction study of muscovite and chlorite: Mineralog. Mag. 39, p.176-188.

1102 Oertel, G., Phakey, P.P., 1972. The texture of a slate from Nantlle, Caernarvon, North Wales.
1103 Texture 1, 1-8.

1104 Oncken, O., Plesch, A., Weber, J., Ricken, W. Schrader, S., 2000. Passive margin detachment
during arc-continent collision (Central European Variscides). In: Franke, W., Haak, V.,
1105 Oncken, O., Tanner, D. (eds). Orogenic Processes: Quantification and Modelling in the
1106 Variscan Belt. Geological Society, London, Special Publications 179, 199-216.
1107

1108 Ortoleva, P., Merino, E., Strickholm, P., 1982. Kinetics of metamorphic layering in
1109 anisotropically stressed rocks. Am. J. Sci. 282, 617-643.

1110 Passchier, C.W., Trouw, R.A.J., 2005. Microtectonics. 2nd Edn., Springer-Verlag.

1111 Paterson, M.S., 1973. Nonhydrostatic thermodynamics and its geologic applications. Reviews in
1112 Geophysics and Space Physics 11-2, 255-389.

1113 Paterson, M.S., Weiss, L.E., 1966. Experimental deformation and folding in phyllite. Geol. Soc.
1114 Amer. Bull. 77, 343-374.

1115 Pehl J., Wenk, H.-R., 2005. Evidence for regional Dauphiné twinning in quartz from the Santa
1116 Rosa mylonite zone in Southern California. A neutron diffraction study. J. Struct. Geol. 27,
1117 1741-1749 [doi:10.1107/S0021889805006059].

1118 Plessmann, W. von, 1964. Gesteinslösung, ein Hauptfaktor beim Schieferungsprozess. Geol.
1119 Mitt. Aachen 4. 69-82.

1120 Plessmann, W. von, 1966. Lösung, Verformung, Transport und Gefüge. Z. Dtsch. Geol. Ges.
1121 115, 650-663.

1122 Ponce de Leon, M., 2012. L'ardoise en Espagne. De la Prehistoire aux Hispano-Romains. In: G.
1123 Marchand and G. Querre (Editors), Roches et Sociétés de la Préhistoire. Presses
1124 Universitaires de Rennes, pp. 351-375.

1125 Popa, N.C., 1998. The (hkl) dependence of diffraction-line broadening caused by strain and size
1126 for all Laue groups in Rietveld refinement. Journal of Applied Crystallography 31, 176-180.
1127 Powell, C.McA., 1972. Tectonic dewatering and strain in the Michigamme slate, Michigan.

1128 Geol. Soc. Am. Bull. 83, 2149-2158.

1129 Price, W.R., Ronck, C.L., 2019. Quarrying for World Heritage Designation: Slate Tourism in

1130 North Wales. Geoheritage. doi:10.1007/s12371-019-00402-0

1131 Raiswell, R., Newton, R.J., Bottrell, S.H., Coburn, P.M., Briggs, D.E.G., Bond, D.P.G., Poulton,

1132 S.W., 2008. Turbidite depositional influences on the diagenesis of Beecher's Trilobite Bed

1133 and the Hunsrück Slate; sites of soft tissue pyritization, American Journal of Science 308,

1134 105-129.

1135 Ratcliffe, N.M., Stanley, R.S., Gale, M.H., Thompson, P.J. and Walsh, G.J., 2011. Bedrock

1136 geologic map of Vermont. U.S. Geological Survey Scientific Investigations Map 3184, 3

1137 sheets, scale 1:100,000. U.S. Geological Survey.

1138 Ramsay, J.G., Huber, M.I., 1983. The Techniques of Modern Structural Geology, Vol. 1: Strain

1139 Analysis. London, Academic Press.

1140 Reed, S.J.B., 2010. Electron Microprobe Analysis and Scanning Electron Microscopy in

1141 Geology. 2nd Edn. Cambridge University Press.

1142 Reks, I.J., Gray, D.R., 1982. Pencil structure and strain in weakly deformed mudstone and

1143 siltstone. J. Struct. Geol. 4, 161-176.

1144 Richards, A.J. (2007). Welsh Slate Craft. A brief account of Welsh slate working. Publisher:

1145 Gwasg Carreg Gwalch

1146 Rietveld, H.M., 1969. A profile refinement method for nuclear and magnetic structures, Journal

1147 of Applied Crystallography 2, 65-71.

1148 Robinson, D. 1971. The inhibiting effect of organic carbon on contact metamorphic

1149 recrystallization of limestones. Contr. Mineral. Petr. 32, 245-250.

1150 Rutter, E.H., 1976. The kinetics of rock deformation by pressure solution. Phil Trans. Roy. Soc.

1151 London A 283, 203-219.

1152 Sander, B., 1950. Einführung in die Gefügekunde der Geologischen Körper, Zweiter Teil: die

1153 Korngefüge. Springer Verlag, Wien.
1154 Sander, B., Sachs, G., 1930. Zur röntgenoptischen Gefügeanalyse von Gesteinen. Z. Krist. 75
1155 (5,6), 550-571.

1156 Schavemaker, Y.A., de Bresser J.H.P., Van Baelen, H., Sintubin, M., 2012. Geometry and
1157 kinematics of the low-grade metamorphic ‘Herbeumont shear zone’ in the High-Ardenne
1158 slate belt (Belgium). Geologica Belgica 15, 126-136.

1159 Schmidt, W., 1925. Gefügestatistik. Tschermaks Mineral. Petrog. Mitt. 38, 392-423.

1160 Sharma, A.K., Sharma, R., Dandi, H.R., 2019. Mineral resources of Himachal Pradesh.

1161 Shimizu, I., 1992. Nonhydrostatic and nonequilibrium thermodynamics of deformable materials.
1162 J. Geophys. Res. 97, B4, 4587-4597.

1163 Siddans, A.W.B. 1972. Slaty cleavage: a review of research since 1815. Earth Science Reviews
1164 8(2), 205-212

1165 Siddans, A.W.B., 1976. Deformed rocks and their textures. Philos. Trans. R. Soc. London 283,
1166 43-54.

1167 Sintubin, M, 1998. Mica (110) Pole figures in support of Marchian behaviour of phyllosilicates
1168 during the development of phyllosilicate preferred orientation. Materials Science Forum 273-
1169 275, 601-606.

1170 Sintubin, M., 1994b. Clay fabrics in relation to the burial history of shales. Sedimentology 41,
1171 1161-1169.

1172 Sintubin, M., 1996. Quantitative approach of the cleavage evolution in the Dinant Allochthon
1173 (Ardennes, France - Belgium); Annales de la Société géologique de Belgique 119, 145-158.

1174 Sintubin, M., 1993. The phyllosilicate texture as a strain marker in structural geology. Material
1175 Science Forum 133-136, 769-776.

1176 Sintubin, M., 1994a. Phyllosilicate preferred orientation in relation to strain path determination
1177 in the lower Paleozoic Stavelot-Venn Massif (Ardennes, Belgium). Tectonophysics 237, 215-
1178 231.

1179 Snow, C.A., Scherer, H., 2006. Terranes of the Western Sierra Nevada foothills metamorphic
1180 belt, California: A critical review. International Geology Review 48, 46-62,
1181 doi.org/10.2747/0020-6814.48.1.46.
1182 Sorby, H.C., 1853. On the origin of slaty cleavage. New Philos. J. (Edinburgh) 55, 137-148.

1183 Sorby, H.C., 1856. On the theory of the origin of slaty cleavage. Phil. Mag. 12, 127-129.

1184 Sprunt, E.S., Nur, A., 1977. Experimental studies of the effects of stress on solution rate. J. Geophys. Res. 82, 3013-3022.

1186 Stearns, B., Stearns, A., Meyer, J., 2000. Slate roofing and gradind in the new millenium. Interface, January 2000, 14-16.

1188 Suárez-Ruiz, I., Flores, D., Graciano Mendonça Filho, J., Hackley, P.C., 2012. Review and update of the applications of organic petrology: Part 1, geological applications. Int. J. Coal Geol. 99, 54-112, DOI: 10.1016/j.coal.2012.02.004.

1191 Taylor, G.H., Liu, S.Y., Teichmueller, M., 1991. Bituminite – a TEM view. Int. J. Coal Geology 18, 71-85.

1193 Tullis, T.E., Wood, D.S., 1975. Correlation of finite strain from both reduction bodies and preferred orientation of mica in slate from Wales. Geol. Soc. Am. Bull. 86, 632-638.

1195 Tullis, T.E., 1976. Experiments on the origin of slaty cleavage and schistosity. Geological Society of America Bulletin 87, 745-753.

1197 Twiss, J.J., Moores, E.M., 2006. Structural Geology, 2nd Edtn. W.H. Freeman, ISBN 0 7167 2252.

1199 Valcke, S.L.A., Casey, M., Lloyd, G.E., Kendall, J.-M., Fisher Q.J., 2006. Lattice preferred orientation and seismic anisotropy in sedimentary rocks, Geophysical Journal International 166, 652-666

1202 van der Pluijm, B.A., Ho, N.-C., Peacor, D.R., Merriman, R.J., 1998. Contradictions of slate formation resolved? Nature 392, 348-348.

1204 Van Noorden, M., Sintubin, M., Darboux, J.-R. 2007. Incipient strain partitioning in a slate belt: Evidence from the early Variscan Monts d'Arrée slate belt (Brittany, France). J. Struct. Geol. 29, 837-849.

1207 Van Noten, K., Hilgers, C., Urai, J., Sintubin, M., 2012. The complexity of 3D stress-state changes during compressional tectonic inversion at the onset of orogeny. In Faulting,
1209 Fracturing, and Igneous Intrusion in the Earth's Crust. Healy et al. Eds. Geol. Soc. London
1210 Spec Publ. 367, 51-69.

1211 Vasin, R., Kern, H., Lokajicek, T., Svitek, T., Lehmann, E., Mannes, D.C., Chaouche, M., Wenk, H.-R., 2017. Elastic anisotropy of Tambo gneiss from Promontogno, Switzerland: a comparison of crystal orientation and microstructure-based modeling and experimental measurements. Geophys. J. Int. 209, 1-20 [doi: 10.1093/gji/ggw487].

1215 Vasin, R., Wenk, H.-R., Kanitpanyacharoen, W., Matthies, S., Wirth, R., 2013. Anisotropy of Kimmeridge shale. J. Geophys. Res. 118, 1-26, doi:10.1002/jgrb.50259.

1217 Vaughan, M. T., Guggenheim, S., 1986. Elasticity of muscovite and its relationship to crystal structure. J. Geophys. Res. 91, 4657-4664.

1219 Vernik, L., Liu, X., 1997. Velocity anisotropy in shales: A petrophysical study, Geophysics, 62, 521-532.

1221 Voisin, L., 1987. Les Ardoisieres de l'Ardenne. Editions Terres Ardennaises, Charleville-Mezieres (La Fortelle)

1223 Voltolini M., Wenk, H.-R., Mondol, N.H., Bjørlykke, K., Jahren, J., 2009. Anisotropy of experimentally compressed kaolinite-illite-quartz mixtures. Geophysics 74, D13-D23, doi: 10.1190/1.3002557.

1226 Von Gehlen K., 1960. Die röntgenographische und optische Gefügeanalyse von Erzen, insbesondere mit dem Zählrohr-Texturgoniometer. Beitr. Mineral. Petrog. 7, 340-388.

1228 Wagner, H.W., 2014. Geologische Untersuchung und Materialprüfung an Dachschiefer-Alt deckungen. Mainzer Geowissenschaftliche Mitteilungen 42, 121-142.

1230 Wagner, H.W., Baumann, H., Negedank, J. and Roschig, F., 1997. Geological, petrographic, geochemical and petrophysical investigations on roofing slates. Mainzer Geowissenschaftliche Mitteilungen 25, 131-184.

1233 Wagner, W., Le Bail, R., Hacar, M., Stanek, S., 1994. European roofing slates part 1: remarks to the geology of mineral deposits. Z. Dtsch. Ges. Geowiss. 40, 68-74.

1235 Wagner, W., Le Bail, R., Hacar, M., Stanek, S., 1995. European roofing slates part 2: geology of selected examples of slates deposits. Z. Angew. Geol. 41, 21-26.
1237 Wallis, D., Lloyd, G.E., Phillips, R.J., Parsons, A.J., Walshaw, R.D. 2015. Low effective fault strength due to frictional-viscous flow in phyllonites, Karakoram Fault Zone, NW India. J. Struct. Geol. 77, 45-61, doi:10.1016/j.jsg.2015.05.010.

1240 Wallis, D., Hansen, L.N., Ben Britton, T., Wilkinson, A.J., 2016. Geometrically necessary dislocation densities in olivine obtained using high-angular resolution electron backscatter diffraction. Ultramicroscopy 168, 34-45.
1243 https://doi.org/10.1016/J.ULTRAMIC.2016.06.002

1244 Walsh, J.A., 2007. The use of the scanning electron microscope in the determination of the mineral composition of Ballachulish slate. Mater. Charact. 58, 1095-1103.

1246 Ward, C., Gómez-Fernandez, F., 2003. Quantitative mineralogical analysis of Spanish roofing slates using the Rietveld method and X-ray powder diffraction data. Eur. J. Mineral. 15, 1051-1062.

1249 Weber, K., 1981. Kinematic and metamorphic aspects of cleavage formation in very low-grade metamorphic slates. Tectonophysics 78, 291-306.

1251 Wenk, H.-R., 2012. Texture analysis by advanced diffraction methods. Chapter 7 in E.J. Mittemeijer and U. Welzel Eds. Modern Diffraction Methods. Wiley-VCH, p173-220.

1253 Wenk, H.-R., Chakmuradyan, A., Foltyn, S., 1996. Quantitative texture measurements of YBCO films on various substrates. J. Mater. Res. A205, 9-22.

1255 Wenk, H.-R., Kanitpanyacharoen, W., Ren, Y., 2019. Slate – a new record for crystal preferred orientation. J. Struct. Geol. 125 (50 year special issue), 319-324, doi.org/10.1016/j.jsg.2017.12.009.

1258 Wenk, H.R., Kanitpanyacharoen, W., Voltolini, M., 2010. Preferred orientation of phyllosilicates: comparison of fault gouge, shale and schist, Journal of Structural Geology 32, 478-489, doi:10.1016/j.jsg.2010.02.003.

1261 Wenk, H.-R., Lutterotti, L., Kaercher, P., Kanitpanyacharoen, W., Miyagi, L., Vasin, R., 2014. Rietveld texture analysis from synchrotron diffraction images: II. Complex multiphase materials and diamond anvil cell experiments. Powder Diffraction 29, 220-232, doi:10.1017/S0885715614000360.
1265 Wenk, H.-R., Matthies, S., Donovan, J., Chateigner, D., 1998. BEARTEX: a Windows-based program system for quantitative texture analysis. Journal of Applied Crystallography 31, 262-269.

1268 Wenk, H.-R., Vasin, R., 2017. Preferred orientation patterns of phyllosilicates in surface clays. Clays and Clay Minerals 65-5, 329-341, doi: 10.1346/CCMN.2017.064069.

1270 Weyl, P.K., 1959. Pressure solution and force of crystallisation-A phenomenological theory. J. Geophys. Res. 64(11), 2001-2025.

1272 White, S.H., Knipe. R.J., 1978. Microstructure and cleavage development in selected slates. Contrib. Mineral. Petrol. 66(2), 165-174.

1274 Williams, P.F., 1972. Relationship between axial plane foliation and strain. Tectonophysics 30, 181-196.

1276 Wood D. S., 1974. Current views of the development of slaty cleavage. Ann. Rev. Earth Sci. 2, 1-35

1278 Wood, D.S., Oertel, G., 1980. Deformation in the Cambrian slate belt of Wales. J. Geol. 88(3), 309-326.

1280 Wood, D.S., Oertel, G., Singh, J., Bennett, H.F., 1976. Strain and anisotropy in rocks. Philos. Trans.R. Soc. London 283, 27-42.

1282 Wright, T.O., Platt, L.B., 1982. Pressure dissolution and cleavage in the Martinsburg Shale. Am. J. Sci. 282(2), 122-135.

1284 Xue, J., 2002. A brief look at Chinese Slate. Traditional Roofing Magazine 2, 16-17.

1285 Zanazzi, P.F., Montagnoli, M, Nazzareni, S., Commodi, P., 2007. Structural effects of pressure on monoclinic chlorites: a single crystal study. Locality: Alpe Raguzzolo, Val Malenco, Italy. Amer. Mineral. 92, 655-661, DOI: 10.2138/am.2007.2341.

1288 Zimak, J., Losos, Z., Novotni, P., Dobe, P. Hladikova, J., 2002. Study of vein carbonates and notes to the genesis of the hydrothermal mineralization in the Moravo-Silesian Culm. J. Czech Geol. Soc. 47, 111-122.
1292 Table 1. Locations of slate samples used for synchrotron X-ray texture analysis

1293 (a) European slates

1294 1 Be-E1 La Fortelle, Herbeumont, Ardennes, Belgium, Devonian.

1295 2 Be-E2 La Fortelle, Herbeumont, Ardennes, Belgium, Devonian.

1296 3 Be-A Rursee, Eifel, Germany. Lower Devonian. Rod-shaped cleavages

1297 4 Sturtz85 Bundenbach, Hunsrück, Germany. Lower Devonian,

1298 5 Sin2 Monts d'Arrée, Brittany, France. Silurian Plougastel Formation.

1299 6 Cz-SLA1 Domasov, Bohemian Massiv, Czech Republic, Moravian-Silesian

1300

1301 (b) Alps, N Spain, China

1302 7 Brg1806e Base of Grevasalvas granite, Bergell Alps

1303 8 ENT1 Entrabasaguas (Luarca), Asturia, N Spain, with microfolds

1304 9 RZ1 Rozadais (Sobradelo), Galicia, N Spain, with lineation

1305 10 Ch1 Pudong, Shanghai, probably from JiangXi, China, Ordovician

1306 11 Ch2a Huangshan, probably from JiangXi, China, Ordovician, crenulation

1307 12 Ch2b Huangshan, probably from JiangXi, China, Ordovician

1308

1309 (c) Slates from Western US

1310 13 EPS39A Tuttletown, New Melones, California, USA, Jurassic

1311 14 EPS39B Tuttletown, New Melones, California, USA, Jurassic

1312 15 DL1 Iron Point, Placer Co. California, Shoo Fly Group, Paleozoic (Lawler)

1313 16 NS1 Spanish Creek Campground, N Quincy, California, N Sierra Nevada

1314 17 NS3 Spanish Creek Campground, N Quincy, California, N Sierra Nevada

1315 18 DL3 Black Butte, Pima Co. Arizona, Blade Group, Paleozoic (Lawler)

1316
<table><caption>**Table 2.** Composition (volume fraction %) and texture maxima in parentheses ((001) max-min, in mrd). Same order as Table 1. Abbreviations for accessories: Ab albite, Cc calcite, Ap apatite, Ilm ilmenite, Ru rutile, H hematite (volume fractions in parentheses).</caption><thead><tr><th></th><th>Muscovite</th><th>Chlorite</th><th>Quartz</th><th>Accessories</th></tr></thead><tbody><tr><td colspan="5">(a) European slates</td></tr><tr><td>Be-E1</td><td>36 (123.3-0.01)</td><td>15 (71.2-0.10)</td><td>41 (1.1)</td><td>Ab (8), Cc</td></tr><tr><td>Be-E2</td><td>38 (88.9-0.01)</td><td>20 (11.4-0.39)</td><td>33 (1.8)</td><td>Ab (9), Cc</td></tr><tr><td>Be-A</td><td>40 (4.0-0.05)</td><td>15 (4.9-0.03)</td><td>40 (1.6)</td><td>Ab (7), Ap</td></tr><tr><td>Sturtz85</td><td>28 (64.3-0.12)</td><td>36 (34.6-0.07)</td><td>36 (1.2)</td><td>Ab (7)</td></tr><tr><td>Sin2</td><td>47 (36-0.02)</td><td>18 (39-0.03)</td><td>35 (1.4)</td><td>Ru</td></tr><tr><td>Cz-SLA1</td><td>35 (7.6-0.3)</td><td>22 (7.2-0.2)</td><td>29 (1.3)</td><td>Ab (14), sphene</td></tr><tr><td colspan="5">(b) Alps, N Spain, China</td></tr><tr><td>Brg1806e</td><td>23 (20.0-0.15)</td><td>14 (18.0-0.26)</td><td>57 (1.7)</td><td>Ab (6)</td></tr><tr><td>ENT1</td><td>33 (31.6-0.07)</td><td>42 (19.3-0.07)</td><td>25 (3.8)</td><td>Monazite</td></tr><tr><td>RZ1</td><td>36 (33.6-0.05)</td><td>34 (26.2-0.06)</td><td>23 (1.5)</td><td>Ab (7), Ru</td></tr><tr><td>Ch1</td><td>26 (9.6-0.14)</td><td>18 (4.0-0.2)</td><td>36 (1.2)</td><td>Ab, Cc (7), Ap</td></tr><tr><td>Ch2a</td><td>25 (15-0.05)</td><td>17 (10-0.03)</td><td>40 (1.3)</td><td>Ab (18), Ilm</td></tr><tr><td>Ch2b</td><td>22 (8.2-0.23)</td><td>17 (7.4-0.17)</td><td>45 (1.4)</td><td>Ab (16), Ap, Ilm</td></tr><tr><td colspan="5">(c) Slates from Western US</td></tr><tr><td>EPS39A</td><td>28 (14-0.3)</td><td>18 (8.4-0.14)</td><td>37 (&lt;1.1)</td><td>Ab(5), Ru, illite (11)</td></tr><tr><td>EPS39B</td><td>30 (22.3-0.01)</td><td>26 (5.12-0.42)</td><td>33 (2.7)</td><td>illite (11)</td></tr><tr><td>DL1</td><td>29 (5.3-0.14)</td><td></td><td>71 (1.4)</td><td></td></tr><tr><td>NS1</td><td>11 (62.6-0.09)</td><td>10 (42.3-0.18)</td><td>79 (1.9)</td><td></td></tr><tr><td>NS3</td><td>39 (45.1-0.12)</td><td>27 (67.4-0.32)</td><td>28 (1.7)</td><td>Ab (6)</td></tr><tr><td>DL3</td><td>44 (19.0-0.05)</td><td>10 (8.7-0.02)</td><td>40 (3.3)</td><td>H (6), Ru</td></tr></tbody></table>
<table><caption>Table 3. Width of (001) texture peak at half maximum (°), ratio in parentheses, average grain aspect ratio (large/small) from microstructures (Fig. 6)</caption><thead><tr><th rowspan="2">Sample</th><th colspan="2">Muscovite</th><th colspan="2">Chlorite</th><th>Quartz</th></tr><tr><th>Width</th><th>Aspect</th><th>Width</th><th>Aspect</th><th>Aspect</th></tr></thead><tbody><tr><td colspan="6">(a) European slates</td></tr><tr><td>Be-E1</td><td>4/4 (1)</td><td>14</td><td>9/9 (1)</td><td>13</td><td>7</td></tr><tr><td>Be-E2</td><td>6/8 (1.3)</td><td>15</td><td>22/27 (1.2)</td><td>18</td><td>5</td></tr><tr><td>Be-A</td><td>30/130 (4.3)</td><td>10</td><td>35/89 (2.5)</td><td>10</td><td>1.5</td></tr><tr><td>Sturtz85</td><td>14/14 (1)</td><td>18</td><td>18/23 (1.3)</td><td>21</td><td>9</td></tr><tr><td>Sin2</td><td>16/21 (1.3)</td><td>25</td><td>18/21 (1.2)</td><td>20</td><td>9</td></tr><tr><td>Cz-SLA1</td><td>18/27 (1.5)</td><td>11</td><td>21/36 (1.7)</td><td>5</td><td>3</td></tr><tr><td colspan="6">(b) Alps, N Spain, China</td></tr><tr><td>Brg1806e</td><td>16/25 (1.6)</td><td>12</td><td>16/29 (1.8)</td><td>7</td><td>6</td></tr><tr><td>ENT1</td><td>18/21 (1.2)</td><td>25</td><td>22/25 (1.1)</td><td>16</td><td>8</td></tr><tr><td>RZ1</td><td>15/20 (1.3)</td><td>21</td><td>17/21 (1.2)</td><td>22</td><td>6</td></tr><tr><td>Ch1</td><td>26/49 (1.9)</td><td>15</td><td>39/77 (2.0)</td><td>4.5</td><td>5</td></tr><tr><td>Ch2a</td><td>15/39 (2.6)</td><td>25</td><td>14/50 (2.6)</td><td>12</td><td>4</td></tr><tr><td>Ch2b</td><td>20/60 (3.0)</td><td>10</td><td>37/61 (1.6)</td><td>20</td><td>5.5</td></tr><tr><td colspan="6">(c) Slates from Western US</td></tr><tr><td>EPS39A</td><td>15/26 (1.7)</td><td>6</td><td>18/39 (2.2)</td><td>6</td><td>2</td></tr><tr><td>EPS39B</td><td>19/26 (1.4)</td><td>7</td><td>22/36 (1.6)</td><td>25</td><td>5</td></tr><tr><td>DL1</td><td>35/42 (1.4)</td><td>14</td><td></td><td>No Chl</td><td>4</td></tr><tr><td>NS1</td><td>6/16 (2.7)</td><td>6</td><td>14/14 (1)</td><td>9</td><td>2.5</td></tr><tr><td>NS3</td><td>6/10 (1.7)</td><td>8</td><td>12/12 (1)</td><td>11</td><td>4</td></tr><tr><td>DL3</td><td>24/24 (1)</td><td>7.5</td><td>22/26 (1.2)</td><td>Little Chl</td><td>4</td></tr></tbody></table>

1323

1324

1325
1326 **Figure Captions**

1327

1328 **Figure 1.** Slate outcrops Bull Point, N Devon, UK (courtesy Martin Fowler)

1329

1330 **Figure 2.** Section perpendicular to the cleavage and bedding plane of slate from Valdeorras,

1331 Spain with slaty cleavage (horizontal) and bedding plane (inclined).

1332

1333 **Figure 3.** Slate quarry in Jiangxi (China) with perfectly planar cleavage surfaces.

1334

1335 **Figure 4.** Slate outcrop in one of the Valdeorras quarries (N Spain), illustrating the compositional

1336 and structural heterogeneities with high quality slates restricted to local areas.

1337

1338 **Figure 5.** Hand specimens displaying the morphology of slates. (a) High quality roofing slate

1339 from La Fortelle quarry, Belgium (Be-E1), (b) sample from La Fortelle with secondary

1340 deformation producing kinking, (c) rodshaped slate from Ruhrsee with two cleavages (Be-A2),

1341 (d) slate from Hunsrück, Germany with fossil Aspidosoma tischbeinianum; here the cleavage is

1342 parallel to the bedding plane (Sturtz85), (e) Slate with crenulation from China (Ch1).

1343

1344 **Figure 6.** View of the town of Matrelange in Wallonia (Belgium) with slates used on roofs and

1345 walls of historic buildings.

1346

1347 **Figure 7.** Historic roofing slate shingle, Bronze age from Castro da Cidade, Spain.

1348

1349 **Figure 8.** Cleaving slate with chisel and hammer from blocks cut with diamond saw in

1350 Valdeorras in 2019.

1351

1352 **Figure 9.** Map of roofing slate production in 2010 (Cardenes et al. 2014).
1353

1354

Figure 10. Production (blue) and import (red) of roofing slate around the World in 2010.

1355

1356

Figure 11. Laue X-ray diffraction image of slate (Sander, 1950, his Fig. 13).

1357

1358

Figure 12 (a) Synchrotron diffraction experiments with (b) sample morphology (aluminum rod is 2mm in diameter), and (c) Pole figure coverage with 11 images. The sample is rotated around the axis x (perpendicular to the cleavage plane) in 15° increments. Equal area projection on the cleavage plane.

1359

1360

1361

1362

1363

Figure 13. Typical diffraction images for 4 samples. (a) Be-E1 with very strong preferred orientation. (b) Be-A with a broader distribution of phyllosilicates. (c) ENT1 with large fraction of phyllosilicates and a relatively broad distribution. (d) NS1 with strong texture but dominant quartz. These images represent one of 11 from which the texture analysis was done.

1364

1365

1366

1367

1368

Figure 14. "Unrolled" diffraction images (bottom) and Rietveld fit (top) as function of 2θ integrated over 10° azimuthal sectors. The vertical coordinate is the azimuth on the diffraction image. Intensity differences as function of azimuth are indicative of preferred orientation. Muscovite and chlorite are highly oriented, quartz (with continuous lines) is random. Note the close similarity of observations and model, indicative of a reliable fit. (a) ENT1 with sharp diffraction peaks, (b) EPS39 with highly diffuse peak for chlorite.

1369

1370

1371

1372

1373

1374

1375

Figure 15. Chlorite EBSD diffraction image of slate BEI from Beiro do Rio (Spain). (a) Raw image, (b) superposed indexed diffractions.

1376

1377

1378

Figure 16 (a) (001) pole figure of muscovite for phyllite from Zillertal (Austria) measured with a Universal stage microscope (Sander, 1950, #109). (b) Pole figure of slate from Wales measured

1379
1380 with a pole figure goniometer (Oertel and Phakey, 1972). Dashed line is sedimentary bedding plane.

1381

1382

1383 Figure 17. Tectonic map of the Rhenohercian belt of Western Germany with the Stavelot-Venn anticlinorium. The Cambro-Ordovician unit is covered by the Lower Devonian unit, separated by the Monschau shear zone. Qualitative (001) pole figures of muscovite are indicated, projected on the slaty cleavage (Fielitz, 1992).

1387

1388 Figure 18. SEM-BE images of selected slate samples. (a) Be-E1 from La Fortelle, Ardennes. Strong alignment of phyllosilicates (grey muscovite, bright chlorite with a kinked crystal above the scale; quartz is dark but also flattened, (b) Slate from Hunsrück, Germany (Sturtz85). (c)
1389
1390
1391 Slate Sin2 from Brittany. (d) Slate Rz1 from N Spain with very strong alignment of muscovite and chlorite. (e) Slate Ch1b from China with an assemblage of interlayered muscovite and chlorite. (f) Quartz-rich slate NS1 from Northern California. In a quartz-rich band are euhedral chlorite crystallites. (g) Slate Ch2a from China with crenulations (compare Fig. 6e), (f) Be-A from Ruhrsee, Germany with two cleavages and also here interlayered muscovite and chlorite (compare Fig. 6c). Scale bars 20 microns.
1392
1393
1394
1395
1396

1397

1398 Figure 19a. Preferred orientation of white mica and chlorite. Samples in same order as in Table 1. (001) and (010) pole figures projected on cleavage plane in equal area projection. Contours in multiples of random distribution (mrd). Logarithmic intensity scale.

1401

1402 Figure 19b. Preferred orientation of white mica and chlorite. Samples in same order as in Table 1. (001) and (010) pole figures projected on cleavage plane in equal area projection. Contours in multiples of random distribution (mrd). Logarithmic intensity scale.

1405

1406 Figure 19c. Preferred orientation of white mica and chlorite. Samples in same order as in Table 1. (001) and (010) pole figures projected on cleavage plane in equal area projection. Contours in multiples of random distribution (mrd). Logarithmic intensity scale.
1407
1408
1409
1410 Figure 20a. Preferred orientation of quartz. Samples in same order as in Table 1. (0001) and (10-11) pole figures projected on cleavage plane in equal area projection. Contours in multiples of random distribution (m.r.d.). Linear intensity scale.

1413
1414 Figure 20b. Preferred orientation of quartz. Samples in same order as in Table 1. (0001) and (10-11) pole figures projected on cleavage plane in equal area projection. Contours in multiples of random distribution (m.r.d.). Linear intensity scale.

1417
1418 Figure 21. (a) SEM-BE image of slate BEI from Beiro do Rio (Spain). Dark grey is albite, lighter grey quartz, light is muscovite and very bright chlorite. Scale bar 10 µm. (b) EBSD map with identified phases indicated by color. Note flattened quartz grains (blue), muscovite (orange) in clusters and chlorite fine-grained (green). Cleavage is horizontal. Scale bar is 200 µm.

1422
1423 Figure 22. EBSD pole figures of sample BEI from Beiro de Rio (Spain). (001) and (010) pole figures of muscovite and chlorite, (0001) and (10-11) pole figures of quartz. Equal area projection on slate cleavage. Linear contour scales.

1426
1427 Figure 23. Plot of quartz content (volume %) as function of (001) maximum of muscovite in slates (from Table 2).

1429
1430 Figure 24. Comparison of muscovite (001) pole figures of (a) Kimmeridge shale (Vasin et al. 2013), (b) Ardennes slate (Wenk et al. 2019), (c) Tambo gneiss (Vasin et al. 2017). Equal area projection on foliation plane. Logarithmic contours.

1433
1434 Figure 25. Seismic velocities of slates from the Belgian Ardennes (Be-E1) and China (Ch2a), based on composition and preferred orientation. Equal area projection on the cleavage plane. For S-waves the polarization direction of the fast S-wave is indicated.
<figure><img src="image_2.png" /></figure>
<figure><img src="image_3.png" /></figure>

bedding plane

2 cm

slaty cleavage
<figure><img src="image_4.png" /></figure>
<figure><img src="image_5.png" /></figure>
<figure><img src="image_6.png" /><figcaption>(b)</figcaption></figure>

<figure><img src="image_7.png" /><figcaption>(c)</figcaption></figure>

<figure><img src="image_8.png" /><figcaption>(e)</figcaption></figure>

<figure><img src="image_9.png" /><figcaption>(d)</figcaption></figure>

<figure><img src="image_10.png" /></figure>
<figure><img src="image_11.png" /></figure>
<figure><img src="image_12.png" /></figure>
<figure><img src="image_13.png" /></figure>
<figure><img src="image_14.png" /></figure>
<figure><img src="image_15.png" /></figure>
<figure><img src="image_16.png" /></figure>

<figure><img src="image_17.png" /></figure>

> 20.000 T/year

<figure><img src="image_18.png" /></figure>

< 20.000 T/year
Import

<figure><img src="image_19.png" /></figure>

Export

Spain
China
Brazil
Ireland
Italy
Belgium
Germany
USA
United Kingdom
France

300

200

100

0

100

200

300

400

500

Million Tons
<figure><img src="image_20.png" /></figure>

c

a

b
<figure><img src="image_21.png" /><figcaption>(b)</figcaption></figure>

<figure><img src="image_22.png" /></figure>

<figure><img src="image_23.png" /><figcaption>(c)</figcaption></figure>
<figure><img src="image_24.png" /></figure>

<figure><img src="image_25.png" /></figure>

<figure><img src="image_26.png" /></figure>

<figure><img src="image_27.png" /></figure>
Azimuth

<figure><img src="image_28.png" /><figcaption>(a) ENT1</figcaption></figure>

Azimuth

<figure><img src="image_29.png" /><figcaption>(b) EPS39A</figcaption></figure>
<figure><img src="image_30.png" /><figcaption>(b)</figcaption></figure>

<figure><img src="image_31.png" /></figure>
<figure><img src="image_32.png" /><figcaption>(a)</figcaption></figure>

<figure><img src="image_33.png" /><figcaption>(b)</figcaption></figure>
<figure><img src="image_34.png" /></figure>

Mica (002) pole figures

-- approximate limit of shear zone
<figure><img src="image_35.png" /></figure>

<figure><img src="image_36.png" /></figure>

<figure><img src="image_37.png" /></figure>

<figure><img src="image_38.png" /></figure>

<figure><img src="image_39.png" /></figure>

<figure><img src="image_40.png" /></figure>

<figure><img src="image_41.png" /></figure>

<figure><img src="image_42.png" /></figure>
<figure><img src="image_43.png" /></figure>

(a)

<caption>Be-E2</caption>

<caption>Be-A</caption>

<caption>Sturtz85</caption>

<caption>Sin2</caption>

<caption>Cz-SLA1</caption>
<figure><img src="image_44.png" /></figure>
<figure><img src="image_45.png" /></figure>
<figure><img src="image_46.png" /></figure>
<figure><img src="image_47.png" /></figure>

<figure><img src="image_48.png" /></figure>

<figure><img src="image_49.png" /></figure>

<figure><img src="image_50.png" /></figure>

<figure><img src="image_51.png" /></figure>

<figure><img src="image_52.png" /></figure>

<figure><img src="image_53.png" /></figure>

<figure><img src="image_54.png" /></figure>

<figure><img src="image_55.png" /></figure>

<figure><img src="image_56.png" /></figure>

<figure><img src="image_57.png" /></figure>

<figure><img src="image_58.png" /></figure>

<figure><img src="image_59.png" /></figure>

<figure><img src="image_60.png" /></figure>

<figure><img src="image_61.png" /></figure>

<figure><img src="image_62.png" /></figure>

<figure><img src="image_63.png" /></figure>

<figure><img src="image_64.png" /></figure>

<figure><img src="image_65.png" /></figure>

<figure><img src="image_66.png" /></figure>

<figure><img src="image_67.png" /></figure>

<figure><img src="image_68.png" /></figure>

<figure><img src="image_69.png" /></figure>

<figure><img src="image_70.png" /></figure>

<figure><img src="image_71.png" /></figure>

mrd
(lin)
<figure><img src="image_72.png" /><figcaption>(a)</figcaption></figure>

<figure><img src="image_73.png" /><figcaption>(b)</figcaption></figure>
Muscovite

Chlorite

Quartz

(001)

(001)

(0001)

(010)

(010)

(10-11)

<figure><img src="image_74.png" /></figure>

<figure><img src="image_75.png" /></figure>

<figure><img src="image_76.png" /></figure>

<figure><img src="image_77.png" /></figure>

<figure><img src="image_78.png" /></figure>

<figure><img src="image_79.png" /></figure>
<figure><img src="image_80.png" /></figure>

Muscovite ((001) mrd)
<figure><img src="image_81.png" /><figcaption>(a)</figcaption></figure>

<figure><img src="image_82.png" /><figcaption>(b)</figcaption></figure>

<figure><img src="image_83.png" /><figcaption>(c)</figcaption></figure>

<figure><img src="image_84.png" /><figcaption>mrd</figcaption></figure>
<figure><img src="image_85.png" /><figcaption>Be E1</figcaption></figure>

<figure><img src="image_86.png" /><figcaption>P (km/s)</figcaption></figure>

<figure><img src="image_87.png" /><figcaption>fast S (km/s)</figcaption></figure>

<figure><img src="image_88.png" /></figure>

<figure><img src="image_89.png" /><figcaption>d S (10m/s)</figcaption></figure>

<figure><img src="image_90.png" /><figcaption>Be E1</figcaption></figure>