In everyday language, we commonly use the word “fabric.” When talking about fabrics that are used to make garments, we mean a patterned cloth made by weaving fibers in some geometric arrangement. But the word “fabric” is not used only to refer to material products. In a philosophical moment we might consider the “fabric of life,” by which we mean the underlying organization of life. As we found to be the case with many terms, the word fabric has a related yet somewhat different meaning in geology. To a structural geologist, the fabric of a rock is the geometric arrangement of component features in the rock, seen on a scale large enough to include many samples of each feature. The features themselves are called fabric elements. Examples of fabric elements include mineral grains, clasts, compositional layers, fold hinges, and foliations. Fabrics that form as a consequence of tectonic deformation of rock are called tectonic fabrics, and fabrics that form during the original formation of the rock are called primary fabrics (Chapter 1). It may sound picky, but structural geologists also make a distinction between microstructure and texture. Whereas texture is sometimes used as a synonym for microstructure, for example igneous texture, we restrict the term texture to crystallographic orientation patterns in an aggregate of grains (see supplementary material) and microstructure to their geometric arrangement.

Tectonic fabrics provide clues to the deformation state of the rock, the geometry of associated folding, the processes involved in deformation, the kinematics of deformation, the timing of deformation (if the fabric is defined by an arrangement of datable minerals), and ultimately about the tectonic evolution of a region. The purpose of this chapter is to explore common tectonic fabrics in rocks, foliations and lineations, and related fabric elements, and to introduce you to their characteristics and interpretation.

FABRIC TERMINOLOGY

We start by examining the inevitable vocabulary to discuss tectonic fabrics (see also Supporting Material). If there is no preferred orientation (i.e., alignment) of the fabric elements, then we say that the rock has a random fabric (Figure 9.1a). Undeformed sandstone, granite, or basalt are rocks with random fabrics. Deformed rocks typically have a non-random fabric or a preferred fabric, in which the fabric elements are aligned in some manner and/or are repeated at an approximately regular spacing (Figure 9.1b).

FIGURE 9.1. The basic categories of fabrics. (a) A random fabric. The fabric elements are dark, elongate crystals. The long axes of these crystals are not parallel to one another. (b) A (1-dimensional) preferred fabric, in which the long axes of elongate crystals are aligned with one another. (c) A foliation. The fabric elements are planar and essentially parallel to one another, creating a 2-dimensional fabric. (d) A lineation. The fabric elements are linear; in this example, we show the alignment of fabric elements in a single plane. [11.1]

There are two main classes of preferred fabrics in rock. A planar fabric, or foliation (Figure 9.1c), is one in which the fabric element is a planar or tabular feature (meaning it is shorter in one dimension than in the other two), and a linear fabric, or lineation (Figure 9.1d), is one in which the fabric element is effectively a linear feature (i.e., it is longer in one dimension relative to the other dimensions). Structural geologists may use the word “fabric” alone to imply the existence of a preferred fabric (as in, “that rock has a strong fabric”), but you should use appropriate modifiers if your meaning is not clear from context alone.

FIGURE 9.2. The distinction between continuous and spaced fabrics. (a) A continuous fabric. The lines represent a planar fabric element that continues to be visible no matter how small your field of view (at least down to the scale of individual grains). (b) A spaced fabric. The rock between the fabric elements does not contain the fabric. The circled areas represent enlarged views. [11.2]

We’re not quite done with terminology yet! Fabrics are complicated features, and there are lots of different adjectives used by structural geologists to describe them. For example, if you can keep splitting the rock into smaller and smaller pieces, right down to the size of the component grains, and can still identify a preferred fabric, then we say that the fabric is continuous (Figure 9.2a). In practice, if the fabric elements are closer than 1 mm (that is, below the resolution of the eye), the fabric is continuous. When there is an obvious spacing between fabric elements, we say that the fabric is spaced (Figure 9.2b).

FIGURE 9.3. Tectonites. (a) An S-tectonite. This fabric is dominantly a foliation, and the rock may tend to split into sheets parallel to the foliation. Within the planes of foliation, linear fabrics are not aligned, or are not present at all. (b) An L-tectonite. The alignment of linear fabric elements creates the dominant fabric, so the rock may split into rod-like shapes. In

L-tectonites, there is not a strong foliation. (c) An L/S-tectonite. The rock possesses a strong foliation and a strong lineation. [11.3]

Rocks with a penetrative tectonic fabric are also called tectonites. When linear fabric elements dominate, the rock is called an L-tectonite, whereas a rock with dominantly planar fabrics is called an S-tectonite, and, not surprisingly, rocks with both types of fabric elements are called LS-tectonites (Figure 9.3). Why create such jargon? Simply to highlight rocks whose internal structure has been substantially changed during deformation. Typically, the deformation that leads to the formation of a tectonite is accompanied by metamorphism, so the fabric is defined by grains that have been partially or totally recrystallized, and/or by new minerals that have grown during deformation (called neomineralization). Because most rocks are deformed, tectonites are among the most common rocks you will see; examples include, slates, schists and mylonites, all of which will be discussed below.

FOLIATIONS

A foliation is any type of planar fabric in a rock. We are admittedly a bit loose in our use of the term planar. Since, strictly speaking, a plane does not contain any curves or changes in orientation, the terms curviplanar or surface would be more appropriate. Although foliations are generally not perfectly planar, structural geologists nonetheless are in the habit of talking about planar fabrics. Thus, bedding, cleavage, schistosity, and gneissosity all qualify as foliations. Fractures, however, are not considered to be foliations, because fractures are breaks through a rock and are not a part of the rock itself. A rock may contain several foliations, especially if it has been deformed more than once. To keep track of different foliations, geologists add subscripts to the foliations: S0, S1, S2, and so on. S0 (or SS) is used to refer to bedding, S1 is the first foliation formed after bedding, and S2 forms after S1. The temporal sequence of foliation development is defined by cross-cutting relationships, but in complexly deformed areas, it may be quite a challenge to determine which foliation is which unless independent constraints on (relative) time are available.

There are many types of tectonic foliations that are distinguished from one another simply on the basis of what they look like. The physical appearance of a foliation reflects the process by which it formed, and the process, in turn, is controlled partly by the composition and grain size of the original lithology, and partly by the metamorphic conditions under which the foliation formed. Different names are used for the different types of foliations, summarized below:

Spaced cleavage (a) Disjunctive cleavage (e.g., stylolitic cleavage);

(b) Crenulation cleavage.

Continuous cleavage (a) Coarse cleavage (e.g., pencil cleavage);

(b) Fine cleavage (e.g., slaty cleavage).

Phyllitic cleavage Continuous cleavage with a distinctive silky luster in low-grade metamorphic rock (lower greenschist facies).

Schistosity Mica-rich foliation with a distinctive high sheen in low- to medium-grade metamorphic rock (greenschist facies).

Gneissic layering Coarse compositional banding or gneissosity in high-grade metamorphic rock (amphibolite and granulite facies).

In the following discussion, we examine different types of foliations roughly in order of increasingly higher metamorphic conditions—cleavage first, then schistosity, then gneissic layering. This progression using natural samples of each is shown in Figure 9.4.

Figure 9.4. Foliations as a function of metamorphic grade, from low (sub-greenschist) metamorphic grade to high (granulite) metamorphic grade: slate, phyllite, schist, gneiss and migmatite.

Cleavage

The term cleavage in rocks has been defined in different ways by different people, so the use of this term in the literature is confusing. We advocate a non-genetic definition, in which cleavage is defined as a secondary fabric element, formed under low-temperature conditions, that imparts on the rock a tendency to split along planes. The point of this definition is to emphasize:

That cleavage is a feature of the rock that forms subsequent to the original formation of the rock.
That the term cleavage is used, in practice, for tectonic planar fabrics formed at or below lower greenschist facies conditions (i.e., ≤ 300°C). The term is not used when referring to the fabrics in schists or in gneiss.
That in a rock with cleavage, there are planes of weakness across which the rock may later break when uplifted and exposed at the surface of the earth, even though cleavage itself forms without loss of cohesion. By this definition, an array of closely spaced fractures is not a cleavage.

We recognize four main categories of cleavage that are differentiated from one another by their morphological characteristics (or, by how they look in outcrop). These are disjunctive cleavage, pencil cleavage, slaty cleavage, and crenulation cleavage.

Disjunctive Cleavage

We find disjunctive cleavage, a type of spaced cleavage, mostly in sedimentary rocks that have been subjected to tectonic stress under sub–greenschist facies metamorphic conditions. It is defined by an array of subparallel fabric elements, called cleavage domains, in which the original rock fabric and composition have been markedly changed by the process of pressure solution. Domains are separated from one another by intervals, called microlithons, in which the original rock fabric and composition are more or less preserved (Figure 9.5).


	FIGURE 9.5. Spaced, disjunctive cleavage (or solution cleavage) in limestone (Harz Mountains, Germany). Cleavage is marked by the narrow dark bands that cut across the original, lighter-colored argillaceous limestone. (a) Outcrop view; (b) close-up view of central portion. In (c) cleavage domain and microlithon of spaced cleavage are illustrated in gently beds. [11.5]

The adjective “disjunctive” implies that the cleavage domains cut across a preexisting foliation in the rock (usually bedding), without affecting the orientation of the preexisting foliation in the microlithons. Because pressure solution is always involved in the formation of a disjunctive cleavage, other terms such as pressure-solution cleavage and stylolitic cleavage are also used for this structure. If the context is clear, some geologists may refer to the structure simply as spaced cleavage, though, as we see later, crenulation cleavage is also a type of spaced cleavage. Earlier in this century, many geologists incorrectly considered cleavage domains to be brittle fractures formed by loss of cohesion. The old term “fracture cleavage” should therefore be avoided when referring to disjunctive cleavage, because a cleavage cannot be composed of fractures. Such arrays of closely spaced fractures should be called a fracture or joint array.

FIGURE 9.6. Evolution of spaced disjunctive cleavage. (a) Pre-cleavage fabric of the rock. In the area indicated by the arrow in the mesoscopic image, there happens to be a greater initial concentration of clay. The microscopic image indicates that the clay flakes are randomly oriented. (b) As shortening and pressure solution occur, the zone in which there had initially been a greater clay concentration evolves into an incipient cleavage domain. At this stage, grains are being preferentially dissolved on the faces perpendicular to S1, and the clay flakes are collapsing together. (c) Ultimately, a clearly defined cleavage domain is visible. In the domain, the clay flakes are packed tightly together and only small relicts of the soluble mineral grains are visible. [11.6]

Now that we’ve gotten through the cleavage terminology, let’s see how disjunctive cleavage forms. Consider a horizontal bed of argillaceous (clay-rich) limestone or sandstone that is subjected to a compressive stress (σ1 in Figure 9.6). Dissolved ions created by pressure solution are transported away from the site of dissolution through a water film that adheres to the grain surfaces. The ions may then precipitate at crystal faces where compressive stress is less, precipitate nearby in the pressure shadows adjacent to rigid grains, or enter the pore fluid system to be carried out of the local rock environment entirely. Note that in order for pressure solution to occur, a thin layer of water molecules must be chemically bonded to grain surfaces. If the water is not bonded, a grain in the water can only sustain isotropic stress, because fluids cannot support shear stresses, and pressure solution won’t occur.

The distribution of clay in rocks is not uniform. Pressure solution occurs more rapidly where the initial concentration of clay is high. How clay affects pressure solution rates remains enigmatic. Perhaps swelling clay, which contains interlayers of bonded water, increases the number of diffusion pathways available for ions and thus multiplies the diffusion rate. Alternatively, the highly charged surfaces of clay grains may act as a chemical catalyst for the dissolution reaction. Field observations suggest that a rock must contain >10% clay in order for solution or stylolitic cleavage to develop. Where it occurs, the process preferentially removes more soluble grains. Thus, in an argillaceous limestone, calcite is removed, and clay and quartz are progressively concentrated. In argillaceous sandstone, the process is effectively the same as that in argillaceous limestone, except that quartz is the mineral that preferentially dissolves, and clay alone is concentrated. As the framework grains of carbonate and quartz are removed, the platy clay grains collapse together like a house of cards. Concentration of clay in the domain further enhances the solubility of the rock, so there is positive feedback. Eventually, a discrete domain develops in which there is a selvage, the material filling the domain, composed of mostly clay (and quartz) with some relict corroded calcite grains. In the selvage, the clay flakes are packed together with a dimensionally preferred orientation. If deformation continues, the domain continues to thicken as pressure solution continues along its edges. As a result, compositional contrast between cleavage domains and microlithons becomes so pronounced that it defines a new stratification in the rock that nearly obscures the original bedding. From this description you see that spaced cleavage formation is identical to the processes by which bedding-parallel stylolites form as a consequence of compaction loading; hence the use of the term stylolitic cleavage.

FIGURE 9.7. Cross-sectional sketches of morphological characteristics used for cleavage description and classification.

(a) Sutured domains; (b) planar domains; (c) wavy domains; (d) an anastomosing array of wavy domains. In (e) the description of spaced cleavage based on domain spacing is shown. [11.7]

The spacing of cleavage domains in a rock mostly depends on the initial clay content. If the clay content is high, the domains are closely spaced. Spacing also changes with progressive deformation. As strain increases, more cleavage domains nucleate and thus the domain spacing decreases. Thirdly, domain spacing may also be related to the magnitude of differential stress, for experiments suggest that the rate of pressure solution is proportional to the magnitude of differential stress. During the process of domain formation, the fabric of the microlithons remains relatively unaffected, though high-resolution microscopic analysis (such as transmission electron microscopy) may indicate the presence of incipient pressure solution features or newly crystallized grains of the soluble mineral in microlithons.

Before we leave the topic of disjunctive cleavage, we need to present the vocabulary that geologists use to describe disjunctive cleavage in the field. The classification and description of disjunctive cleavage is based on the characteristics of surface morphology of domains, and on domain spacing (Figure 9.7). If the clay content in the host rock is low, domain surfaces are severely pitted. In cross section, such domains resemble the toothlike or jagged sutures on a skull; this type of cleavage domain is called a sutured domain. If the clay content is high, cleavage domains tend to have thick selvages that have smooth borders; these domains are called non-sutured. When viewed under high magnification with a microscope, you will see that thick domains, in some cases, are composed of a dense braid of threadlike sutured domains. Cleavage domains can be either wavy, if the domain undulates, or planar, if the domain does not. If wavy domains are closely spaced, such that the domains merge and bifurcate and give the fabric the appearance of braided hair, the cleavage is called anastomosing.

The average spacing between domains is also a useful criterion for classification of cleavage. Figure 9.7e shows a simple field classification based on spacing. If the spacing between domains is greater than about 1 m, then the rock really doesn’t have an obvious fabric. Such isolated pressure-solution domains, when formed in response to tectonic stress, are called tectonic stylolites. The adjective “tectonic” distinguishes these structures from bedding-parallel stylolites formed by compaction. In general practice, if domains are spaced between about 10 cm and 1 m apart, we call the feature a weak cleavage; a spacing of 1–10 cm defines a moderate cleavage, and a spacing of less than 1 cm denotes a strong cleavage. When cleavage domains are less than 1 mm apart, the cleavage is continuous (see slaty cleavage, below). We caution you again that different authors use these adjectives differently, so our definitions are only generalizations.

Pencil Cleavage

If a fine-grained sedimentary rock (shale or mudstone) breaks into elongate pencil-like shards because of its internal fabric, we say that it has a pencil cleavage (Figure 9.8). Typically, pencils are 5–10 cm long and 0.5–1 cm in width. In outcrop, pencil cleavage looks as if it results from the interaction of two fracture sets (and in some locations, it is indeed merely a consequence of the intersection between a fracture set and bedding), but actually the parting reflects an internal alignment of clay grains in the rock.

FIGURE 9.8. Pencil cleavage in shale, with horizontal bedding trace (Virginia, USA). The roughly equal bedding and cleavage intensity creates the characteristic pencil shape. [11.8]

Pencil cleavage forms because of the special characteristics of clay. The strong shape anisotropy of clay flakes creates a preferred orientation parallel to bedding when they settle out of water and are compacted. This preferred orientation imparts the tendency for clay-rich rocks to break on bedding planes that is displayed by shale. Shale, mudstone, and slate are all names that are used to describe clay-rich rocks; mudstone is the general term for these rocks, whereas shale and slate are foliated clay-rich rocks. After compaction, a strain ellipsoid representing the state of strain in the shale would look like a pancake parallel to bedding (Figure 9.9a). Now, imagine that the shale is subjected to layer-parallel shortening. Cleavage formation processes begin to take place: large detrital phyllosilicates fold and rotate, while fine grains undergo pressure solution along domains perpendicular to the shortening direction, and new clay crystallizes. Microfolding may occur during this stage of the process, but because of the fine grain size these microfolds are only visible under the microscope. In addition, quartz may begin to dissolve, and as these framework grains are removed, clay flakes collapse so that their basal planes are perpendicular to the plane of shortening.

FIGURE 9.9. Sketches illustrating the progressive development of slaty cleavage via the formation of pencil structures. (a) Compaction during burial of a sedimentary rock produces a weak preferred orientation of clay parallel to bedding. A representative strain ellipsoid is like a pancake in the plane of bedding. (b) Shortening parallel to layering creates an incipient tectonic fabric. Superposition of this fabric on the primary compaction fabric leads to the formation of pencils. The representative strain ellipsoid is elongate, with the long axis parallel to the pencils. (c) Continued tectonic shortening leads to formation of slaty cleavage at a high angle to bedding. The phyllosilicates are now dominantly aligned with the direction of cleavage, and a representative strain ellipsoid is oblate and parallel to cleavage. [11.9]

The plane defined by new or rotated clay grains is roughly perpendicular to the shortening direction, so it forms a tectonic foliation at high angles to the original bedding (Figure 9.9c). At an early stage during this process, the new tectonic fabric is comparable in degree of development to the initial bedding-parallel fabric. At this stage, the strain ellipsoid representing this state would look something like a big cigar, and the rock displays pencil cleavage (Figure 9.9b). In sum, pencil cleavage is a fabric found in weakly deformed shale in which the tendency of the shale to part on bedding planes is about the same as the tendency for it to part on an incipient tectonic cleavage that is at a high angle to bedding. Deformation in most areas produces a fabric that is much stronger than the original bedding parting, so pencils are not as common as the more evolved stage, represented by slaty cleavage.

Slaty Cleavage

With pencil cleavage a snapshot of an early stage in the process, slaty cleavage represents the fully evolved step of foliation formation at low temperatures. As shortening perpendicular to cleavage planes accumulates, clay throughout the rock develops a preferred orientation at an angle to the original sedimentary fabric and this orientation dominates over the primary fabric (Figure 9.9c). The finite strain ellipsoid for cleavage development at this stage has the shape of a pancake that parallels the tectonic fabric. Formation of slaty cleavage occurs by much the same process as did the formation of disjunctive cleavage in argillaceous sandstone or limestone, but the resulting domains are so closely spaced that effectively there are no uncleaved microlithons, and the entire rock mass displays the tectonically induced preferred orientation. When a rock has this type of continuous fabric, we say that it contains slaty cleavage (Figure 9.10).

FIGURE 9.10. An inclined syncline with a well-developed slaty cleavage that parallels the axial plane of the fold (southern Appalachians, USA). [11.10]

In other words, slaty cleavage is defined by strong dimensionally preferred orientation of phyllosilicates in a very clay-rich rock, and the resulting rock, which is considered to be a low-grade metamorphic rock, is called a slate. Slaty cleavage tends to be smooth and planar. This characteristic, coupled with the penetrative nature of slaty cleavage, means that slates can be split into thin sheets, which made them popular roofing materials in the nineteenth century and early parts of the twentieth century.

FIGURE 9.11. Photomicrograph of a continuous slaty) cleavage from Newfoundland (Canada); width of view is 2 mm. [11.11]

At microscopic magnifications the detailed character of slaty cleavage becomes apparent. In many cases only a network of anastomosing surfaces around partially dissolved quartz or feldspar grains is preserved (Figure 9.11). Because slaty cleavage forms under temperature conditions that mark the onset of metamorphism (250°C–350°C), the mineralogy of slate tends to resemble that of shale. However, there is a notable decrease in the amount of interlayered water in clays; that is, smectite, the water-bearing clay, transforms to illite. As metamorphic temperatures increase, the illite becomes more muscovite-like, though it is still fine grained. Thus, rocks with higher-grade slaty cleavage display a distinct sheen on cleavage planes. They are also significantly harder than shale, so they ring when hit with a hammer. Such slates provide excellent material for the construction of pool tables and, for classic roofing tiles.

Phyllitic Cleavage

When metamorphic conditions reach the lower greenschist facies, the clay and illite in a pelitic rock react to form muscovite and chlorite. If reaction occurs in an anisotropic stress field, these phyllosilicates grow with a strong preferred orientation. Rock that is composed of strongly aligned fine-grained white mica and/or chlorite is called phyllite (Figure 9.4) and the foliation that it contains is sometimes called phyllitic cleavage. The mineralogy and fabric of phyllites give the rock a distinctive silky luster. Phyllitic cleavage is intermediate between slaty cleavage and schistosity.

Cleavage Refraction

Commonly we find that cleavage orientation parallels or fans systematically around the orientation of a fold’s axial plane (Figure 9.12). Cleavage fans change from convergent in mechanically strong (competent) beds (e.g., sandstone, limestone) to divergent around the axial plane in relatively weak (incompetent) units (e.g. shale, marl). As a result, cleavage changes orientation from bed to bed, a pattern that is called cleavage refraction. Cleavage refraction is the change in cleavage attitude that occurs where cleavage domains cross from one lithology into another of different competency, and reflects variation in the local strain field between beds. In graded beds, this change in cleavage orientation occurs gradually across the bed, producing curved cleavage surfaces (Figure 9.12). Changes in domain spacing typically accompany cleavage refraction, as this is also controlled by lithology.

Figure 9.12. Cleavage in folds, with continuous (slaty) cleavage in the shale (dark) and discontinuous cleavage in sandstone (light) fanning systematically around the fold. Note the sharp refraction from shale to sandstone base and gradual cleavage orientation change in upward-fining sandstone layer (oval); hand for scale.

If the deformed unit is dominated by shale, then the regional slaty cleavage forms approximately parallel to the axial plane of regional folds. In cases where the cleavage is not parallel to the axial plane, but cuts obliquely across the folds, we say that the cleavage is cross-cutting. The occurrence of cross-cutting cleavage may indicate that the cleavage was superimposed on preexisting folds, or that there were local complexities in the strain field. For example, rotation of a thrust sheet during folding may cause fold hinges to become oblique to the regional shortening direction.

FIGURE 9.13. The relationship between cleavage, axial plane, and enveloping surface in folds with transecting cleavage. The counterclockwise cleavage transection that is illustrated may be indicative of dextral transpression. Note the obliquity of the bedding-cleavage intersection lineation to the hinge line. [11.23]

Transecting cleavage is a term for cross-cutting cleavage that forms in transpressional environments (meaning there are components of both pure and simple shear). Counterclockwise transecting cleavage, in which the cleavage cuts counterclockwise relative to roughly synchronous fold hinges (Figure 9.13), indicates a component of dextral shear (dextral transpression). Similarly, a component of sinistral shear may produce a clockwise transecting cleavage. However, the meaning of transecting cleavage and its use as a displacement indicator remains controversial.

Crenulation Cleavage

A lithology containing a closely and evenly spaced foliation that is shortened in a direction at a low angle to this foliation will crinkle like the baffles in an accordion. In fine-grained lithologies like slate or phyllite, these microfolds are closely spaced and the spacing tends to be uniform. The axial planes of the crenulations define a new foliation called crenulation cleavage (Figure 9.14).

FIGURE 9.14. The two basic categories of crenulation cleavage. (b) Symmetric crenulation cleavage; (c) asymmetric (sigmoidal) crenulation cleavage. The arrows indicate a possible component of shear associated with this crenulation geometry. [11.13]

Like mesoscopic folds, crenulations can be symmetric (Figure 9.14b) or asymmetric (Figure 9.14c). In a given outcrop, both symmetric and asymmetric crenulation cleavages may occur. For example, on the limbs of a fold, asymmetric crenulation occurs, whereas in the hinge zone of the fold, the crenulation is symmetric, analogous to the geometry of parasitic folds described earlier (Chapter 8).

FIGURE 9.15. Photomicrograph of crenulation cleavage with incipient differentiation (Pyrenees, Spain). Folded and annealed muscovite (light) traces microfolds around areas with quartz (gray) concentration; width of view is ~0.5 mm. [11.14]

A prerequisite for the formation of crenulation cleavage is the existence of a preexisting strong lamination or foliation. Crenulation cleavage won’t form in a sandstone, but may form in a shale with a strong bedding foliation, or in a rock that already contains a slaty or phyllitic cleavage (Figure 9.15). By definition, crenulation cleavage in outcrop is a later foliation that was superimposed on a preexisting (S0 or S1) foliation.

FIGURE 9.16. Differentiation during the formation of crenulation cleavage. (a) Fairly homogeneous composition, before migration of the quartz. (b) Quartz accumulates in the hinges of the crenulations, and the phyllosilicates are concentrated in the limbs; the result is the formation of compositionally distinct bands in the rock. (c) Complete transposition of the S1 foliation into a new S2 cleavage or schistosity. [11.15]

Crenulation cleavage forms under conditions that are also amenable to the occurrence of fluid-based pressure solution. When the starting rock contains a mixture of quartz and clay or fine-grained mica, the quartz is preferentially removed from the limbs of the microfolds and precipitates in the hinges as the crenulations form (Figure 9.16). Gradually, phyllosilicates concentrate on the limbs and quartz is concentrated in the hinges. This mineralogical differentiation can be so complete that the old foliation disappears entirely and is replaced entirely by a new foliation, which is defined not only by preferred orientation of the phyllosilicates, but also by micro-compositional layering (Figure 9.16b). If quartz is largely removed by progressive solution, a new foliation eventually develops and the crenulated appearance of the rock fades (Figure 9.16c). If this happens, all traces of the original fabric, the one predating the crenulation, may be destroyed, leaving the mistaken impression of a deformed rock with only one foliation.

Synthesis: Bedding-Cleavage Relationships

The relationship of bedding and cleavage and their geometries can provide powerful clues to the regional interpretation of an area, such as the relative position of an outcrop with respect to a large regional fold. Cleavage refraction can further help to determine the facing (“younging”) of folds. Take the example of an initially recumbent fold (F1) with an S1 axial planar slaty cleavage that is folded by a second folding generation (F2) creating an upright, open fold with crenulation cleavage S2 (Figure 9.17).

FIGURE 9.17. Cleavage-bedding relationships and cleavage refraction in upright and overturned limbs of upward-facing and downward-facing folds. A-D are described in the text. [11.24]

As a consequence, a portion of the F2 fold faces up, while another portion of the fold faces down. If the strata in this fold still contain the S1 cleavage, and if this cleavage is axial planar to the F1 fold in clay-rich horizons, then, on the overturned limb of the upward-facing part of the F2 fold (A), bedding dips more steeply than cleavage, and on the upright limb (B) cleavage dips more steeply than bedding. On the upright limb of the downward-facing part of the large fold (C), bedding dips more steeply than cleavage, whereas on the overturned limb cleavage is steeper (D). You might also find a crenulation cleavage (S2) that is axial planar to the F2 fold, especially in its hinge region. The geometry of cleavage refraction further indicates the younging direction (i.e., upright and overturned limbs) in the large fold. Figure 9.17 is another one of those information-rich diagrams exploring a range of geometries and relationships that will prove to be quite helpful in the field; that is, once you have figured it out. After taking a suitable amount of time to study this illustration, we’ll continue with our description of foliations by a look at higher grade metamorphic rocks.

Schistosity

When metamorphic conditions get into the middle greenschist facies, the minerals in a pelitic lithology react to form coarser-grained mica and other minerals. When these reactions again take place in an anisotropic stress field, the mica has a strong preferred orientation. The resulting rock is a schist, and the foliation it displays is called schistosity. The specific assemblage of minerals that forms depends not only on the pressure and temperature conditions at which metamorphism takes place, but also on the chemical composition of the protolith and on the degree to which chemicals are added or removed from the rock by migrating fluids. Conveniently, a schist is named by the assemblage of metamorphic index minerals that it contains (e.g., a garnet-biotite schist). In schist that contains porphyroclasts (relict large crystals) or porphyroblasts (newly grown large crystals), the schistosity tends to be wavy, as the micas curve around the large crystals.

Figure 9.18. Fold hinges in transposed gneiss near Parry Sound, Grenville Orogen (Ontario, Canada); hand for scale. [11.16]

Gneissosity

Foliated gneiss is a metamorphic rock in which the foliation is defined by compositional banding (Figure 9.18). Commonly, light and dark bands of felsic and mafic mineralogy alternate. The light-colored layers are rich in feldspar and quartz, whereas darker layers contain more of the minerals amphibole/pyroxene (and/or biotite). This color banding in gneiss is called gneissic layering or gneissosity. Under the metamorphic condition for gneiss formation (amphibolite to granulite facies), muscovite reacts to form feldspar, so the rock contains no schistosity. Gneiss can be derived from a sedimentary protolith, in which case it is called a paragneiss, or an igneous protolith, in which case it is called an orthogneiss. It is often difficult to decide whether a particular rock is an orthogneiss or a paragneiss, however; this requires careful field and petrologic analysis. A special type of gneiss, called augen gneiss, contains relatively large feldspar clasts floating in a finer-grained matrix [lonely mappers felt that altered feldspar grains look like eyes, or “Augen” in German].

FIGURE 9.20. Formation of a gneiss. (a) Inheritance from an original lithology; (b) creation of new compositional banding via transposition; (c) metamorphic differentiation; (d) lit-par-lit intrusion. [11.17]

How does the compositional banding in gneiss form? There may be several processes involved in the formation of gneissic layering (Figure 9.20). First, it may occur by inheritance from original compositional contrasts. If the protolith (the rock from which it formed) was a stratified sequence with layers of different composition (such as alternating sandstone and shale), then metamorphism will transform this sequence into a compositionally banded metamorphic rock. Secondly, it may result from transposition via folding of an earlier layering. Transposition is a common process during deformation under high-temperature conditions, as we also discussed elsewhere (see Chapters X and X). Rock containing compositional layering that is subjected to intense, high-grade deformation may develop isoclinal folds. If the hinges of the fold are detached and a new sequence of compositional layers has formed, then we say that the new layering is a type of transposed foliation. The layering in a rock with a transposed foliation does not represent the original stratigraphy of the rock, though it may have been derived from it. Thirdly, gneissic layering may be formed by metamorphic differentiation when the thermodynamics governing diffusion during metamorphism causes certain ions to be excluded from the formation of new metamorphic minerals in a layer. The excluded ions accumulate to form different minerals in an adjacent layer. Thus, minor differences in the original composition of successive layers may be amplified into major compositional changes after metamorphism. The resulting rock with alternating layers of different composition is a gneiss. Finally, gneissic banding may originate from an igneous process called lit-par-lit intrusion (French for “layer-by-layer”), when melts inject as thin sills along many weak planes in the protolith, with interlayering of sills and host rock defining the gneiss. Usually the process of injection is accompanied by passive folding, so the igneous nature of the contacts can be obscured.

When metamorphic temperatures are sufficiently high, a rock begins to melt, but not all minerals melt at the same temperature. Quartz, some feldspar, and muscovite melt at lower temperatures than mafic minerals like amphibole, pyroxene, and olivine. Therefore, when a rock of intermediate composition begins to melt, certain minerals become liquid while others remain solid. The minerals that stay solid until higher temperatures are achieved are called refractory minerals. When only part of a rock melts, we say that it has undergone partial melting. A rock that is undergoing partial melting is a mixture of pockets of melt and lenses of solid, both of which are quite soft. Shortening will cause the mass to flow, and the contrasting zones of melt and solid fold and refold much like chocolate and vanilla batter in marble cake. When this happens in rock, the resulting semi-chaotic mixture of light and dark layers is called a migmatite. Because of their origin and often chaotic nature, the analysis of structures in migmatite may provide little or no information about the regional deformation.

Mylonites

Mylonitic Foliations

Earlier we introduced the fault rock mylonite that forms at elevated pressure and temperature condition, and that involves crystal plastic processes. Most mylonites show at least one well-developed foliation that is often at a low angle to the boundary of the shear zone. This penetrative foliation is known as the S-foliation, derived from the French word for foliation, schistosité (Figure 9.21). Its angle with the shear zone boundary may be as little as a few degrees, at which point it is hard to distinguish the S-foliation from a foliation or banding that parallels the shear zone boundary, called the C-foliation after the French word for shear, cisaillement (Figure 9.21). Late shear surfaces that offset and crenulate the C and S foliations are called C’ foliations or shear bands (Figure 9.21b).

FIGURE 9.21. (a) Photomicrograph of a quartzite mylonite with S and C foliations, and (b) C-C′ foliations in a micaceous mylonite (b); width of view is ~1 mm and ~15 cm, respectively. [12.10]

Multiple mylonitic foliations may seem confusing, at first, but when correctly identified these fabrics provide powerful information on the processes and kinematics of shear zones, including displacement sense (see below). Mylonitic foliations reflect grain-shape fabrics or represent discrete shear surfaces. Thin-section mylonites in quartzites, granites and marbles often reveal the presence of an S-foliation that is defined by elongate grains, which reflects the activity of grain-scale crystal plastic processes that tend to elongate grains toward the extension axis of the finite strain ellipsoid. The C-foliation can also be expressed as compositional bands or by grain size changes (Figure 9.21a), tracking the orientation of the shear zone boundary.

Shear-Sense Indicators

Ductile shear zones concentrate displacement at deeper levels in the crust, where recognizable markers that determine offset, such as bedding, are often absent. Consider a greenschist-facies shear zone in a large granitic body. The granitic rocks on either side of the mylonite are indistinguishable, so there is nothing at first glance to predict the sense of displacement, let alone the magnitude of displacement. Sense of displacement describes the relative motion of opposite sides of the zone (left-lateral or right-lateral, up or down, and so on), whereas magnitude of displacement is the distance over which one side moves relative to the other. The solution to this challenge is to look for shear-sense indicators in ductile shear zones.

Geometry and Orientation

The recognition and interpretation of shear-sense indicators require that we view a shear zone in a consistent orientation. Most mylonites contain at least one foliation and a lineation, which we use as an internal reference frame. In the field we look for outcrop surfaces (or cut an oriented sample in the lab) that are perpendicular to the mylonitic foliation and parallel to the lineation (Figure 9.22a).

FIGURE 9.22. (a) Schematic, right-lateral ductile shear zone showing mylonitic foliation (S) and lineation. The optimal surface for study is the XZ-plane of the finite strain ellipsoid. (b) Apparent difference in shear sense, which results from observing the shear-sense indicators in different surfaces. Note that a surface perpendicular to both lineation and foliation gives no shear-sense information. [12.4]

We make the reasonable assumption in this case that the lineation coincides with the movement direction of the shear zone. When two foliations are present, this surface is also generally perpendicular to their intersection. This plane, which parallels the XZ-plane of the finite strain ellipsoid, maximizes the expression of the rotational component of the deformation; in all other surfaces this component is less.

Next, we must place the orientation of our surface in the context of the region. Say, we find that a right-lateral displacement is the surface of observation. Were we to look at this same surface from the opposite side, the displacement would appear to be left-lateral (Figure 9.22b). This is not a paradox, but simply a matter of reference frame; we encountered the same situation with fold vergence (Chapter 8). Because the displacement sense is the same in geographic coordinates, it is a good habit to analyze surfaces in the same orientation across the field area to avoid confusion. If this is not possible, make careful field notes of the orientation of the surface in which you determined shear sense. Back in the laboratory the rock saw offers complete control, provided you oriented the sample prior to removing it from the field. Having cautioned you sufficiently about orientation, let us now look at types of shear-sense indicators, which we organize into three groupings:

grain-tail complexes;
disrupted grains;
composite foliations.

Other shear-sense indicators are folds and crystallographic fabrics, which will not be discussed in this chapter.

Grain-Tail Complexes

Mylonites commonly contain large grains or aggregates of grains that are surrounded by a finer matrix; for convenience, we use the term grain in a general sense to describe both large single grains and coherent grain aggregates. These grains may have tails of material with a composition and/or grain shape and size that differ from the matrix, such that they are distinguishable. For example, large feldspar grains connected by thin layers of finer-grained feldspathic material are common in gneiss (Figure 9.23).


FIGURE 9.23. Grain-tail complexes. (a) A K-feldspar clast with a tail of fine-grained plagioclase of the σ-type (California, USA). (b) A plagioclase clasts with tail of K-spar of the δ-type (Ontario, Canada). Width of view is ∼10 cm. GE: 45°20'40.83"N, 80° 0'21.46"W (approximate). [12.5]

The tail may represent highly attenuated, preexisting mineral grains, it may be a consequence of dynamic recrystallization of material at the rim of the grain, or it may be material formed by syn-kinematic metamorphic reactions. During deformation, the grains act as rigid bodies and we may be able to determine the sense of displacement from the tails. Based on their relationship with the shear-zone foliation, we recognize two types of grain-tail complexes: σ-type and δ- type. A third type, θ-complexes, has been proposed, but their interpretation is equivocal and we do not further discuss them.

FIGURE 9.24. Grain-tail complexes as shear-sense indicators. (a) σ-type complex, (b) δ-type complex, and (c) the evolution of σ-type complex into δ-type grain-tail complex. [12.6]

[Animation]

Grain-tail complexes of the σ-type are characterized by wedge-shaped tails that do not cross the reference plane when tracing the tail away from the grain (Figure 9.24a). Sometimes the tail is flat at the top and the other side is curved toward the reference plane. Overall it has a stair-stepping geometry in the direction of displacement. This grain-tail geometry looks like the Greek letter σ (at least in the case of right-lateral displacement), hence the name σ-type. Figure 9.23a is a field example of a σ-type complex in a quartzo-feldspathic shear zone. Obviously, in the case of left-lateral displacement the geometry is a mirror image.

In δ-type grain-tail complexes the tail wraps around the grain such that it cross cuts the reference plane when tracing the tail away from the grain (Figure 9.24b). If you rotate the Greek letter δ over 90° you will see why we use this symbol. Figure 9.23b is a field example of a δ-type complex in another feldspathic shear zone. The rotation on the δ-type complex is counterclockwise for left-handed and clockwise for right-handed displacement. The stair-stepping geometry that we find in σ-type grain-tail complexes is no longer a characteristic of displacement in δ-type complexes.

It is common to find both σ- and δ-types in one surface, as well as their mixtures, because they are related (Figure 9.24c). One explanation for a mixed type is a varying relationship between the rate of recrystallization/neocrystallization and rotation of the grain. If tail formation is fast relative to rotation, the tails are of the σ-type. If, on the other hand, the rotation of the grain is faster, the tail will simply be dragged along and wrap around the grain (δ-type). The case of preexisting tails, which often occur with pegmatites that are incorporated into a shear zone, falls in the latter category.

The presence of both σ- and δ-type grain-tail complexes may indicate different rates of tail growth, different initial grain shape, different times of tail formation, or different coupling (see later section). From these variables it is clear that we should use grain-tail complexes with considerable caution for strain quantification purposes, but there is no doubt about their power as shear-sense indicators. In practice, geometries of the σ-type may be more difficult to recognize, but δ-type complexes offer intuitively obvious and unequivocal information on shear sense.

Fractured Grains and Mica Fish

Minerals in some mylonites, such as feldspar, may have deformed by fracturing while others, like quartz, display crystal plastic deformation. As long as the approximate orientation of these fractures before shear is known, we can determine the shear sense from their displacement (Figure 9.25).

FIGURE 9.25. Placing dominos between a pair of hands to demonstrate sense of shear from fractured grains, if the fractures are at a low angle (a) or high angle (b) to the shear plane. Note that rotation according to the “domino model” of individual segments of fractured grains in (a) is the same as that for rotated grains.

[12.8]

Fractures oriented at low angles to the mylonitic foliation have a displacement sense that is consistent with the overall shear sense of the zone; these fractures are called synthetic fractures (Figure 9.25a). Fractures at angles greater than ∼45° to the foliation show an opposite sense of movement; these are called antithetic fractures (Figure 9.25b). The opposite motion is not contradictory, as we can see from a simple experiment. Place a series of dominos upright between your hands, and move your hands in opposite directions. You will notice that as long as the angle of the dominos with your hands remains greater than ∼45°, the displacement between individual dominos is opposite (antithetic) to the relative movement direction of your hands. At lower angles you will find that displacement of the dominos has a motion that is the same (synthetic) to the motion of your hands. An example is shown in Figure 9.26a. Fractured minerals and clasts behave similarly in shear zones, and some therefore call this the domino model for shear sense.


FIGURE 9.26. (a) Photomicrograph of fractured feldspar grain showing bookshelf- or domino-type, antithetic displacement. (b) Mica fish in a quartz mylonite. Right-lateral displacement in both images; width of view ∼0.5 mm. [12.7]

Feldspar and quartz are not the only minerals useful for determining shear-sense in mylonites. It is quite common to find large phyllosilicate grains, such as mica and biotite in quartzo-feldspathic rocks, and phlogopite in marbles, that display a characteristic geometry. The micas are connected by a mylonitic foliation, and their basal (0001) planes are typically oriented at an oblique angle to the mylonitic foliation, such that they point in the direction of the instantaneous elongation axis. In this orientation they show a stair-stepping geometry in the direction of shear (Figures 9.26b), which is geometrically similar to σ-type grain-tail complexes that also step up in the shear direction. The evolution of mica fish is shown in Figure 9.27.

FIGURE 9.27. The formation of mica fish; (a) to (c) show successive stages in mica-fish development, with C-foliation marking the shear plane. [12.9]

When phyllosilicates are large enough to be seen in hand specimens they look like scales on a fish (hence they are called mica fish) and you can use a simple field test to determine their approximate orientation. The basal planes of phyllosilicates are excellent reflectors of light, so when you turn the foliated shear-zone sample in the sun you encounter an orientation that is particularly reflective. When this happens and the sun is behind you, you are looking in the direction of shear. The method is affectionately known as the “fish flash,” and was obviously developed by those fortunate geologists who work in sunnier parts of the world.

Composite Foliations: C-S and C-C′ Structures

Most mylonites show at least one well-developed foliation that is generally at a low angle to the boundary of the shear zone. This penetrative, mylonitic foliation is known as the S-foliation. Its angle with the shear zone boundary may be as little as a few degrees, at which point it is hard to distinguish from an occasional foliation that parallels the shear zone boundary, called the C-foliation. A third foliation, showing discrete synthetic shear displacements oblique to the shear zone boundary, is called the C′-foliation.

FIGURE 9.28. Characteristic geometry of (a) C-S and (b) C-C′ composite foliations in a dextral shear zone. The C-surface parallels the shear zone boundary and is a surface of shear accumulation (i.e., not parallel to a plane of principal finite strain). The S-foliation is oblique to the shear-zone boundary and approximates the XY-plane of the finite strain ellipsoid. The C′-foliation in (b) synthetically offsets an earlier foliation (C or composite C/S). [12.11]

Thin-section study of well-developed mylonites in quartzites, granites and marbles often shows the presence of a foliation that is defined by elongate grains. This foliation reflects the activity of crystal-plastic processes that tend to elongate grains toward the extension axis of the finite strain ellipsoid, and is called the S-foliation. It is uncertain whether the S-foliation exactly tracks the XY-plane of the finite strain ellipsoid; nor is it certain whether S- and C-foliations form simultaneously or sequentially. These distinctions are only important when C-S structures are used to determine the degree of non-coaxiality or the kinematic viscosity number (Chapter 6). For their purpose as a shear-sense indicator, these questions are somewhat academic, because in all cases the geometry of a C-S structure gives the same shear sense. The long axis of elongate grains in the S-surface points up in the direction of shear, and the shear direction is perpendicular to the intersection line of the S- and C-foliations (Figure 9.28a).

Lastly, another common foliation as shear-sense indicator is a series of oblique, discrete shears that are found in strongly foliated mylonites. These small shears, called C′-surfaces because they accumulate shear strain, are particularly common in phyllosilicate-rich mylonites and crenulate or offset the mylonitic foliation (Figure 9.28b). C′- foliations are, therefore, also called shear bands or extensional crenulations. The offset on C′-surfaces is in the same direction as the overall displacement in the shear zone (i.e., displacement along C). The C′-surfaces contrast with S-surfaces that do not appear to displace the C-surface, suggesting that C′-surfaces form late in a mylonite’s evolution. As with C-S structures, the strain significance of C-C' structures is incompletely understood, but their formation reflects a component of extension along the main anisotropy (the C-surface) of the mylonite. Thus, the shear sense on C-C′ structures is synthetic to the sense of shear of the zone as a whole (Figure 9.28b).

Summary of Shear-Sense Indicators

Shear sense in ductile shear zones is only reliably determined when two or more different indicators give a consistent sense of displacement. So we close this section with a summary diagram (Figure 9.29) showing common shear-sense indicators that may be encountered in a ductile shear zone.


FIGURE 9.29. Summary diagram of shear-sense indicators in (a) dextral shear and (b) sinistral shear zone. A copy of this figure on a transparency (for right- and left-lateral shear) makes a handy addition to your field notebook. [12.12]

Of these indicators, C-S and δ-clasts are most readily interpretable. Planar objects may have synthetic and antithetic shear, and C′ surfaces may reflect opposing motions; for example, boudins may show antithetic bookshelf faulting or synthetic displacement shear. The occurrence of a particular indicator will vary within a zone and even within the same outcrop, as a function of the dominance and mineralogy of grains and presence of foliations. Except for fractured grains, the shear sense can be determined from any of these indicators without knowing their (original) relationship to the shear-zone boundary. Figure 5.32 is another handy reference tool for your field notebook.

FOLIATION KINEMATICS AND METAMORPHISM

Mineral growth at elevated temperatures provides an opportunity to determine the relative timing of metamorphism and deformation. We define pre-kinematic growth when the mineral growth occurs before deformation, syn-kinematic growth when growth occurs during deformation (Figure 9.30), and post-kinematic growth, you guessed it, when minerals grow after deformation. The shape and internal geometry of minerals and their relationship to external fabric elements, in particular foliations, help us to determine this relative temporal relationship. Such newly grown minerals are called porphyroblasts.

FIGURE 9.30. Photomicrograph of syn-kinematic garnet. Long dimension of view is ∼1.5 mm. [13.6]

The beautiful crystals that decorate geodes may come to mind when thinking of mineral growth, but rocks at depths beyond a few kilometers rarely have any open spaces. Mineral growth under these conditions mostly occurs by the replacement of preexisting phases. For example, garnet grows by the consumption of another mineral, and phases that are not involved in the reaction (accessory phases such as zircon, monazite), or phases that are left over because the rock does not contain the right mix of ingredients, may become inclusions. These inclusions may form ordered trails that define an internal foliation (Si) in the porphyroblast; for contrast, the foliation outside the blast is called the external foliation (Se). The relationship between Si and Se allows us to determine the relative timing of mineral growth and deformation (Figure 9.31).

FIGURE 9.31. Schematic diagram showing diagnostic forms of porphyroblasts that grow before (pre-kinematic), during (syn-kinematic), and after (post-kinematic) layer-parallel shear, layer-perpendicular shortening and layer-parallel shortening. The relative timing is based on the relationship between internal foliation (Si; dashed lines) and external foliation (Se; solid lines). [13.5]

Pre-kinematic growth is characterized by an internal foliation (Si) whose shape is unrelated to the external foliation (Se), and typically the internal and external foliations do not connect. At the other end of the spectrum we find post-kinematic growth, which shows an external foliation that continues into the grain seemingly without any disruption. In the intermediate case (syn-kinematic growth), the timing of growth and deformation coincide. The evidence for syn-kinematic growth is an external foliation that can be traced into a blast containing an internal foliation with a pattern different but connected to Se. A classic example is a snowball garnet (early stage is shown in Figure 9.30), where Si spirals around the core of the garnet until it connects with Se, which generally displays a simpler pattern.

In many instances we may not find evidence for syn-kinematic growth as compelling as that of snowball garnets. Nevertheless, it seems generally true that mineral reactions are triggered by deformation, because the stress gradients that exist during deformation promote material transport and grain growth (e.g., by fluid-assisted or “wet” diffusion; Chapter 7). Thus, mineral assemblages and mineral compositions may be metastable until deformation triggers the reactions that produce equilibrium assemblages for the ambient metamorphic conditions. Lastly, mineral growth during deformation is common, but is not a rule, with environments of contact metamorphism as one example exception.

FOLIATIONS AND STRAIN

Do foliations provide constraints on the nature of strain in a rock or region? Unfortunately, we can only offer a wishy-washy answer: it depends. We look at cleavage, porphyroblasts and mylonitic fabrics for an answer.

Cleavage and Strain

In order to determine the relationship of cleavage to strain, we seek strain markers in a cleaved rock. Red slate is a good lithology for such studies, because it often contains reduction spots, which are small regions where iron in rock is reduced and turns greenish in color (Figure 9.32).

FIGURE 9.32. Reduction spots in slate that are elliptical on the cleavage surface and ellipsoidal in 3D (Vermont, USA). They can be used as strain markers if they were formed as spherical regions around a reducing phase prior to deformation; coin for scale. [11.18]

If these spots start out as early diagenetic spheres around an inclusion, they make ideal strain markers because they behave in a totally passive manner, meaning that they have no mechanical contrast with the host rock. In studies of reduction spots in slate, geologists find that the deformed spots are flattened ellipsoids, and that the plane of flattening (i.e., the XY principal plane of the finite strain ellipsoid) is essentially parallel to the slaty cleavage. In these cases, therefore, cleavage appears to approximate the orientation of a principal plane of strain, with the shortening direction being perpendicular to cleavage.

But cleavage is probably not strictly parallel to the XY principal plane of strain in all situations. For example, when cleavage occurs in flexural slip folds, or adjacent to fault zones, there may be a component of shear on the cleavage planes themselves, and when this happens the cleavage by definition is not a principal plane. Moreover, a spaced cleavage may initiate as a principal plane of strain, but subsequent folding of the bed in which cleavage occurs may rotate it away from the bulk principal strain directions. Reduction spot studies give estimates of a total shortening strain (e1) for the formation of slaty cleavage in the range of 50–60%.

FIGURE 9.33. To lose or not to lose volume; that’s the question. (a) A block of height l and width w is shortened and forms a cleavage. If there is volume loss strain, then w′ < w, but l′= l (we assume no change in the third dimension; that is, the intermediate strain axis (Y) equals 1). If there is volume constant strain, then l′′ > l, as in (c). [11.19]

Key to understanding strain during low-grade cleavage formation is the question of whether cleavage formation is a volume-constant strain or a volume-loss strain process. Imagine a cube of rock that is 10 cm long on each edge. If it is shortened in one direction, but does not stretch in the other directions (X = Y = 1 > Z), it must lose volume during deformation (Figure 9.33). Alternatively, if shortening in one direction causes it to expand by an equal proportion in one other direction (X > Y = 1 > Z or plane strain) or in two other directions (X > Y > Z), then the strain may be volume-constant. Remember that formation of low-grade cleavage involves significant activity in the way of dissolution and new growth of grains. If the ions of soluble minerals enter the pore water system, they can be carried out of the local rock system by the movement of groundwater. Thus, volume loss during cleavage formation is a likely possibility; in fact, studies of deformed markers (such as the fossil graptolites, whose regular protrusion spacing can be used to measure finite strain) and geochemical studies (Figure 9.34) demonstrate that rock volume may have decreased by as much as 50% during cleavage formation.

FIGURE 9.34. Volume loss may occur by preferential removal of certain elements. If the composition of microlithons (Q-domains) is considered constant, then the amount of volume loss can be calculated (see equation). Based on relatively immobile elements TiO2, Y, and Zr, volume loss is calculated as ∼45%. [11.20]

The analysis of finite strain and volume loss indicate that plane strain conditions are representative for cleavage formation, and that calculated flattening strains are apparent rather than real (field of apparent flattening; Chapter 6). But not all strain analyses of cleaved rocks yield indications of large volume loss. Probably the degree of volume loss is affected by the degree to which the rock is an open or closed geochemical system during deformation, and whether the fluids passing through the rock during deformation are saturated or undersaturated with respect to the soluble mineral phase. Volume-loss scenarios also have an important implication concerning the nature of rock–water interactions during deformation. If we assume that most of the volume loss reflects dissolution and removal of certain minerals, then we can calculate the minimum volume of fluid required for a given amount of volume loss if we know the solubility of the mineral and the amount that can be held in solution (saturation). Such calculations suggest that the amounts of fluid needed for the proposed amounts of volume loss are very large. The results of these calculations are usually expressed in terms of the ratio of fluid volume to rock volume (called the fluid/rock ratio). In the example in Figure 9.34, with quartz solubility of 0.019 kg/l and 20% undersaturated fluid, this ratio is >>100. But whether such large fluid/rock ratios are geologically reasonable remains a matter of debate.

Mylonitic Foliations and Strain

In Chapter 6 we presented several methods to quantify strain in deformed rocks, including the use of grain shapes. But do any of these methods apply to mylonitic shear zones? Two approaches provide a first-order estimate of the amount of finite strain in shear zones, namely grain rotation and foliation orientation, though these methods are not without their limitations either as you will discover from further literature study.

Rotated Grains

The formation of a ductile shear zone involves an internal rotation (vorticity; Chapter 6) that can be recorded in rotated grains. From grains that preserve this rotation we can determine components of finite strain. We saw this earlier with δ-type grain-tail complexes, but is also found in minerals that grow and incorporate matrix grains during rotation. In particular the mineral garnet shows this behavior, in which “trapped” matrix grains eventually produce a spiraling trail. Such garnets are often called snowball garnets, for apparent reasons (Figure 9.35).

FIGURE 9.35. Snowball garnet (Sweden). Garnet is isotropic and appear black under crossed polarizers; width of view is ∼5 mm. [12.13]

Let’s starts with a simple analog experiment. Place a ball bearing or a marble between oppositely sliding hands and you will see that its rotation is directly related to the motion of your hands (Figure 9.36). In fact, the amount of rotation of the marble is proportional to the relative displacement of your hands (i.e., the amount of simple shear); mathematically, this relationship is straightforward:

β = tan ψ = γ Eq. 9.1

where β is the rotation angle in radians (1 radian is 180°/π), ψ is the angular shear, and γ is the shear strain. However, if the ball bearing is greasy, the rotation angle may be less, because there is some slip between your hands and the ball bearing. This adds to Equation 9.1 a factor that describes the coupling between matrix and grain:

β = Ω tan ψ = Ω γ Eq. 9.2

where the parameter Ω describes the mechanical coupling between the ball bearing and your hands.

FIGURE 9.36. (a) A simple ball-bearing experiment (b) that illustrates the relationship between rotation [β] and shear (ψ, γ). [12.14]

The value of Ω is equal to 1 for full coupling (clean ball bearing), less than 1 for partial coupling (greasy ball bearing), and 0 for no coupling. In kinematic terms, coupling describes the degree by which internal vorticity is converted to spin. There is no single value for Ω that is unique for natural rocks, but if we assume that grains rotate in a viscous (Newtonian) fluid, considerable slippage will occur at the contact between matrix and grain, and we obtain a theoretical value for Ω of 0.5. Thus, by measuring the rotation angle of the spiraling snowball garnet we can determine the shear strain given some assumption about coupling. Ignoring decreased coupling, so Ω = 1, gives us an estimate of the minimum shear strain.

Foliation Orientation

Next we look at the strain distribution in a foliated shear zone where the host foliation is progressively rotated (Figure 9.37).

FIGURE 9.37. Small-scale, right-lateral shear zone in anorthosite (feldspathic gneiss), showing deflection of the mylonitic (S) foliation (Grenville Orogen, Ontario, Canada); width of view is ∼20cm (mirror image); C-foliation E-W, S-foliation SW-NE. (b) Heterogeneous strain in shear zones resulting in progressive foliation deflection, quantified by variably oriented and elongated ellipses [12.16]

We find that the S-foliation is increasingly parallel to the shear zone boundary as we approach the center of the zone. If we assume that the trace of the S-foliation tracks the X-axis of the finite strain ellipsoid (Figure 9.38), the shear strain is determined by the equation:

γ = 2/tan 2φ′ Eq. 12.3

where φ′ is the angle between the S-foliation and the shear-zone boundary.

FIGURE 9.38. Angular relationship, φ′, between S-foliation and shear-zone boundary, and shear strain, γ, in a perfect shear zone (kinematic vorticity number, Wk, is 1). Dashed lines are traces of circular sections of the strain ellipse. [12.17]

Note that small differences in the value of φ′ (which occur in the center of the zone) significantly change the calculated magnitude of the shear strain. For example, compare the results for φ′ values of, say, 10° and 5°. As before, in practice, we use this method to obtain a minimum strain from shear zones.

LINEATIONS

A lineation is any fabric element that can be represented by a line, meaning that one of its dimensions is much longer than the other two. There are multiple types of lineations. Some are associated with other geologic structures (such as folds or boudins), some are visible only on specific surfaces in a rock body, while others reflect the arrangement and shape of mineral grains or clasts within the rock. Some lineations reflect strain axes or kinematic trajectories, while others appear to have no kinematic significance. For their description, we broadly group lineations into three categories: form lineations, surface lineations, and mineral lineations.

Form Lineations

The hinge of any fold is by definition a linear feature. If folds are closely spaced, the fold hinges effectively define a rock fabric that we can measure as a fold hinge lineation (Figure 9.39a).

FIGURE 9.39. Examples of form lineations (l). (a) Fold and crenulation hinges, and (b) mullions. [11.25]

Similarly, a crenulation lineation is defined by the hinge lines of the microfolds in a crenulated rock. Why bother measuring crenulation lineations if a crenulation cleavage is present in the rock? The reason is that sometimes the interval in which crenulations occur is not thick enough for cleavage domains to be measurable. For example, deformation and metamorphism of thick sandstone beds separated by thin interbeds of shale will create quartzite beds separated by thin layers of phyllite. Even if the phyllite layer is so thin that a crenulation cleavage plane cannot be measured, it may be possible to see and measure the crenulation lineation. When intense deformation detaches the limbs of folds, as occurs during fold transposition, isolated fold hinges may be left in the rock. Such hinges are called rods and typically occur in a multilayer composed of phyllite (or schist) and quartzite; the quartz layers are relatively rigid and define visible folds, the limbs of which may be thinned so severely that they pinch out, and the quartz flows into the hinge zone, where it is preserved as a rod. Rodding may also occur in mylonites, because the progressive folding in mylonites may generate rootless isoclinal folds whose limbs are detached and whose hinges (the rods) have rotated into parallelism with the shear direction. Mullions are cusp-like corrugations that form at the contact between units of different competencies in a deformed multilayered sequence (Figure 9.39b); the axes of mullions are a lineation. Typically, the more rigid lithology occurs in convex bulges that protrude into the ductile lithology, and the bulges connect in pointed troughs. Because of their mechanical origin, mullions cannot be used as a facing indicator.


Figure 9.40. Boudins in sandstone with quartz-filled necks (Spain); (b) geometry of boudins and boudin necks in a folded layer, defining a form lineation (l).

Boudins are tablet-shaped lenses of a relatively rigid lithology, embedded in a weaker matrix, that have collectively undergone layer-parallel stretching (Figure 9.40). In the third dimension, these long tabular bodies are separated by narrow boudin necks that are linear objects. In rare cases we find boudins that record extension in two directions, lovingly called chocolate-tablet boudinage. Other elongate objects in rock that are useful lineations include elongate pebbles and elongate pumice fragments. Again, when the long axes of such elongate objects in a rock are aligned, then they define a measurable lineation. The elongation of embedded objects is generally a manifestation of deformation, which we used in Chapter 6 for strain quantification. When using objects for structural analysis, however, it is important to make sure that their alignment is a tectonic and not a primary feature, such as clasts alignment during deposition.

Surface Lineations

An intersection lineation is a linear fabric element formed by, as the name suggests, the intersection of two planar fabric elements (Figure 9.41a). One intersection lineation that structural geologists often use in the field is the bedding-cleavage intersection, which is manifested by the traces of cleavage domains on a bedding plane (or vice versa). When cleavage is parallel to the axial plane of a fold, the bedding-cleavage intersection must parallel the hinge line in mostly cylindrical folds. The field application is powerful, as it allows one to predict regional fold geometry in areas with otherwise sparse outcrop.

FIGURE 9.41. Sketches of surface lineations. (a) Intersection lineation of bedding (S0) and (axial plane) cleavage (S1) in a fold, and (b) slip lineation on a (normal) fault surface. (c) Slip lineation (marked by pencil) on veined fault surface. [11.27]

Slip lineations form on surfaces that move in opposite directions (Figure 9.41b). This occurs, for example, on fault surfaces, but also at the interface between beds in flexural slip folds. There are two basic types of slip lineations: groove lineations, formed by plowing of surface irregularities, and fiber lineations that are formed when vein mineral fibers precipitate along a sliding surface (Figure 9.41c). Slip lineations, of course, are parallel to the slip direction and their roughness may indicate slip sense.

Mineral Lineations

A rock with a mineral lineation means that the fabric element defining this lineation is the size of a mineral grain or a cluster of mineral grains. Mineral lineations are commonly present on the foliation plane of metamorphic rocks, and particularly in foliated mylonites (Figure 9.42).

Figure 9.42. Mineral lineation on foliation surfaces (Swiss Alps). [11.28]

There are several types of mineral lineations. Not all of these have the same tectonic significance, so it is important to determine what type of mineral lineation is present in the rock. Some minerals, such as kyanite and amphibole, grow such that they are very long in one direction relative to the other two directions. If the long axes of the crystals are aligned in a rock, they create a mineral lineation. The alignment may be due either to growth of the crystal in a preferred direction (controlled by differential stress or by flow-controlled diffusion) or because elongate grains are rotated toward a principal strain direction during deformation. This type of linear fabric is taken to indicate the direction of stretching in the rock and is therefore called a stretching lineation. It is quite common in deformed metamorphic terranes to find a consistently oriented mineral lineation that reflects the regional transport direction of the rocks; for example, parallel to displacement in areas of regional thrusting.

We earlier mentioned elongate objects, such as stretched pebbles in a conglomerate. Elongate grains or grain clusters produce the same phenomenon, only on a smaller scale. Quartz, which deforms quite readily under metamorphic conditions, may deform into long ribbons that define a distinct lineation in outcrop, called rods.

Interpretation of Lineations

In structural analysis, the bedding-cleavage intersection lineation is widely used, because this lineation offers a clue to the orientation of folds in a region where the hinges themselves may not be exposed. Other types of lineations, however, may be more difficult to interpret, because there are at least two alternative interpretations for their origin. First, a lineation can parallel a principal strain; specifically, the direction of stretching or elongation. When we talk about stretching lineations, as defined by elongate mineral grains or pebbles, we are implying that the lineation is roughly parallel to the direction of maximum elongation (the X-axis of the finite strain ellipsoid). Other lineations, like boudin necks, are roughly perpendicular to the stretching direction.

Second, a lineation can parallel the shear direction, meaning that it is parallel to a vector defining the motion of one part of a rock with respect to another. Slip lineations (fibers or grooves) are good examples of displacement-direction lineations. However, the shear direction is not parallel to the stretching direction, except in special cases. Only in zones of high shear strain do the shear direction and finite elongation direction approach parallelism; in other words, grains are stretched and mineral clusters are smeared out in the direction of shear. Clearly, you must understand the process of their formation when interpreting the meaning of lineations.

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v1.3; Last update: 15-Sep-2023