Carbonate Petrography

Carbonate petrography is the study of limestones, dolomites and associated deposits under optical or electron microscopes greatly enhances field studies or core observations and can provide a frame of reference for geochemical studies.

25 strangest Geologic Formations on Earth

The strangest formations on Earth.

What causes Earthquake?

Of these various reasons, faulting related to plate movements is by far the most significant. In other words, most earthquakes are due to slip on faults.

The Geologic Column

As stated earlier, no one locality on Earth provides a complete record of our planet’s history, because stratigraphic columns can contain unconformities. But by correlating rocks from locality to locality at millions of places around the world, geologists have pieced together a composite stratigraphic column, called the geologic column, that represents the entirety of Earth history.

Folds and Foliations

Geometry of Folds Imagine a carpet lying flat on the floor. Push on one end of the carpet, and it will wrinkle or contort into a series of wavelike curves. Stresses developed during mountain building can similarly warp or bend bedding and foliation (or other planar features) in rock. The result a curve in the shape of a rock layer is called a fold.

Point counting technique used in petrography

What is point counting?




Point investigation may suggests that of describing rock in an unbiased and quantitative means.

A rock which could be represented within the field as “Fine-grained Sandstone” has additional way more data to relinquish up once examined more closely. a skinny section viewed underneath the magnifier could show grains of variable sizes, overgrowths on those grains, overgrowths of various clays, every of variable designs and habits, pore areas between the grains, pores part crammed, grains part dissolved, and so on, in nearly infinite detail. most components of the detail area unit of interest to 1 or alternative of the specialist geoscientists answerable for creating an initial exploration appraisal of an opportunity, or for developing a promising prospect, or for exploiting the known field. it's the task of the petrographer to amass and document this data and purpose investigation is that the principal technique used.

Point investigation may be a applied mathematics technique. It involves watching an outsized range of points on the slide, recording precisely what's seen at every purpose and so collecting an outline from all the knowledge recorded. so as to be a statistically valid illustration, the amount of points represented is usually three hundred – five hundred.

Moving the slide that range of times, and guaranteeing constant purpose isn't represented quite once, may be a important task. concerning forty years agone, Swift Engineering developed a tool that enraptured the slide in equal increments on a line; at the tip of a line, the petrographer would re-set it to the beginning and move it up a line, very much like employee performed a “carriage come back and line feed” operation.

Added to the current, Swift, and later previous Scientific Instruments, developed a 20-channel device for recording up to twenty completely different classes of things seen. however twenty is way too few for the big range of minerals, clays, pore varieties and alternative materials of interest found in an exceedingly skinny section, even before we tend to try and record the habit and morphology, the combos, and therefore the diagenetic relationships.
Number of points to count and “1.5 times grain size”

By contrast, if we were to take the same thin-section dimensions as previously but now with a rock which has an average grain size of “Very Coarse Sand Upper”, we would find that the “step size (at least) 1.5 times grain size” adage means we run out of rock to look at when we have stepped about 130 times. We simply cannot get enough material to look at. What has gone wrong this time?

We have to look again at why we are point-counting. The reason is to collect data from a representative sample of the rock, and hence to be able to make predictions or deductions about the rock as a whole. Where did “400 points” come from? And why “1.5 times grain size”? What do they contribute to the determination of a strategy for finding a representative sample of rock?

If you are an oil company sedimentologist feeling hard-done-by because you have to describe 400 points on your slides, think of coal petrologists working to the ICCP recommendation of 900 points. But even they are getting off lightly: ASTMS recommend 2000 – 4000 for clinker. Where do these numbers come from? Felix Chayes is frequently referred to as the father of mathematical geology and in 1953 he published a discussion of how variability in the rock should determine the target point count. This should be obvious: if we have counted 300 points already and they were all quartz, the 301st point is odds-on not to add any new information, the relative percentages will almost certainly remain unchanged; if we have counted 300 points already and 100 were quartz, 100 feldspar and 100 microcline, then the 301st pint is bound to change the relative percentages, but only slightly; but if, at the other extreme, every one of those 300 points was a different mineral, then the 301st is going to effect a large percentage change, no matter what we see there. Logically, if we want a statistically representative (or, to put it another way, accurate) description of the rock, we should keep counting until the changes effected by further counting are negligible.

The “magic number” approach to setting point count targets is a hangover from a bygone age. If one has no simple way to place a grid over a slide, then step size is a simple, if crude and inefficient, means of determining a sampling strategy. The dictum “1.5 times average grain size” comes from a desire not to count a grain twice. This itself is dubious thinking (if we were sampling the electorate in Britain, would we only ask one person in London, because we don’t want to ask Londoners twice? I appreciate that this is different in kind and in principle, but should at least serve to provoke thought) but is worse if it leads us to count the wrong number of points. It is surely obvious that, if there are fewer but larger grains, we shouldn’t be counting less points. (It may be less obvious that there is an implication that we should, on the contrary, be counting more points; statistical theory is occasionally counter-intuitive and we will come back to this point).

Number of steps is most frequently determined by accountants rather than geologists. With the predominance of out-sourcing, number of points to count will usually be determined by available budget. More correctly it should be determined by our degree of tolerance for errors taken with our expectation (based on background knowledge and experience) for spatial and statistical variation in the rock.

If we start with a knowledge of what the data will be used for, and hence how much we can tolerate errors, we can devise a point-counting strategy to suit. For example, if the petrophysicists wish to know whether certain trace minerals are present, we should count until we have a reasonable confidence that, were they present, we would have seen them. Statistics is very good at measuring and determining this kind of confidence: tests exist which, if applied correctly, will give us a numerical measure of confidence. We can ask the petrophysicist: “How sure do you want us to be – 90% confident? 95% confident? 99% confident?” and, based on her/his answer, determine how many points we should count.

The additional input to this calculation of confidence is the geologists expectation of variation in the rock. We can, however, refine this as we go along, by determining the actual degree of variation in this particular thin section. This would add inordinately to the effort if tallying manually, even using one of the ready-reckoned graphs as published by, for instance, xyz; but, if using a computer to do the tallying of counts, it is trivial to ask the computer to tell us when to stop counting, based on our target for percentage confidence.

Different answers might come from the engineers, who are only going to average our data into blocks hundreds or thousands of times the size of our thin sections, or from specialist geologists, who may be using the data as input to models for determining diagenetic history, basin development, etc. In each case, we can determine a strategy from the requirements – are we trying to catch odd occurrences of extreme events (trace minerals) or quantify relative abundances of significant occurrences (e.g. clay habit)? – and, if necessary, take the worst case from each to allow the data to be of maximum use to all parties. The down side is that someone has to think: this becomes a genuine test of skill, using a large amount of background information on the specific rocks or on rocks from similar environments, instead of a simplistic application of a universal formula handed down through the generations and having little or no relevance. It is therefore a task for the geologist, not for the company accountant.

Is There Oil Beneath My Property?



Is There Oil Beneath My Property? First Check the Geological Structure

Accumulations of crude oil and gas require a source rock, which is an organic-rich sedimentary rock that produces petroleum as the organisms decompose following burial. Petroleum migrates from the source rock until it encounters impermeable rock or the earth's surface, where it is lost in seeps. Porous and permeable reservoir rocks allow petroleum to accumulate below impermeable layers. Geologic structures provide traps to further concentrate the accumulating petroleum. Anticlines are the best traps and the first recognized by geologists searching for oil. These structures can form "oil pools" with stacked levels of water (lowest), oil and natural gas (highest) against the impermeable rock above the reservoir.



Other types of traps have been recognized including faults, unconformities, and facies changes. Modern petroleum exploration is sophisticated, but there is no way short of drilling to determine the presence of oil. "Wildcat" wells only have a 1 in 10 chance of discovery, which explains why future oil supplies will be more difficult and costly to find.........
.........       

Oil Trap

A set of conditions that hold petroleum in a reservoir rock and prevents its escape by migration.

Major types of Petroleum Traps.  In all cases, impermeable rock encloses or caps the petroleum.


Types of fossils


What it a fossil? 


The word fossil is derived from the Latin fossilis meaning an object that has been dug up from the ground. Fossils are the evidence for the existence of once-living animals and plants and may be either the preserved remains of an organism or evidence of its activity. 

Types of fossils

Trace fossils


Trace fossils are the preserved impressions of biological activity. They provide indirect evidence for the existence of past life. They are the only direct indicators of fossil behaviour. As trace fossils are usually preserved in situ they are very good indicators of the past sedimentary environment. Trace fossils made by trilobites have provided an insight into trilobite life habits, in particular walking, feeding, burrowing, and mating behaviour.

Chemical fossils 


When some organisms decompose they leave a characteristic chemical signature. Such chemical traces provide indirect evidence for the existence of past life. For example, when plants decompose their chlorophyll breaks down into distinctive, stable organic molecules. Such molecules are known from rocks more than 2 billion years old and indicate the presence of very early plants. 

Coprolites


Coprolites are fossilized animal feces. They may be considered as a form of trace fossil recording the activity of an organism. In some coprolites recognizable parts of plants and animals are preserved, providing information about feeding habits and the presence of coexisting organisms. 

Body fossils 


Body fossils are the remains of living organisms and are direct evidence of past life. Usually only hard tissues are preserved, for example shells, bones, or carapaces. In particular environmental conditions the soft tissues may fossilize but this is generally a rare occurrence. Most body fossils are the remains of animals that have died, but death is not a prerequisite, since some body fossils represent parts of an animal that are shed during its lifetime. For example, trilobites shed their exoskeleton as they grow and these molts may be preserved in the fossil record.

Fault Terminology

Faults are much more complex and compound features that can accommodate large amounts of strain in the upper crust. The term fault is used in different ways, depending on geologist and context. A simple and traditional definition states:

A fault is any surface or narrow zone with visible shear displacement along the zone.

This definition is almost identical to that of a shear fracture, and some geologist use the two terms synonymously. Sometimes geologists even refer to shear fractures with millimeter- to centimeter-scale offsets as microfaults. However, most geologists would restrict the term shear fracture to small-scale structures and reserve the term fault for more composite structures with offsets in the order of a meter or more. 
The thickness of a fault is another issue. Faults are often expressed as planes and surfaces in both oral and written communication and sketches, but close examination of faults reveals that they consist of fault rock material and subsidiary brittle structures and therefore have a definable thickness. However, the thickness is usually much smaller than the offset and several orders of magnitude less than the fault length. Whether a fault should be considered as a surface or a zone largely depends on the scale of observation, objectives and need for precision.
Faults tend to be complex zones of deformation, consisting of multiple slip surfaces, subsidiary fractures and perhaps also deformation bands. This is particularly apparent when considering large faults with kilometerscale offsets. Such faults can be considered as single faults on a map or a seismic line, but can be seen to consist of several small faults when examined in the field. In other words, the scale dependency, which haunts the descriptive structural geologist, is important. This has led most geologists to consider a fault as a volume of brittlely deformed rock that is relatively thin in one dimension:

A fault is a tabular volume of rock consisting of a central slip surface or core, formed by intense shearing, and a surrounding volume of rock that has been affected by more gentle brittle deformation spatially and genetically related to the fault.

The term fault may also be connected to deformation mechanisms (brittle or plastic). In a very informal sense, the term fault covers both brittle discontinuities and ductile shear zones dominated by plastic deformation. This is sometimes implied when discussing large faults on seismic or geologic sections that penetrate much or all of the crust. The term brittle fault (as opposed to ductile shear zone) can be used if it is important to be specific with regard to deformation mechanism. In most cases geologists implicitly restrict the term fault to slip or shear discontinuities dominated by brittle deformation mechanisms, rendering the term brittle fault redundant:

A fault is a discontinuity with wall-parallel displacement dominated by brittle deformation mechanisms.

By discontinuity we are here primarily referring to layers, i.e. faults cut off rock layers and make them discontinuous. However, faults also represent mechanical and displacement discontinuities.  A kinematic definition, particularly useful for experimental work and GPS-monitoring of active faults can therefore be added:

 Faults appear as discontinuities on velocity or displacement field maps and profiles. The left blocks in the undeformed map  a) and profile (b) are fixed during the deformation. The result is abrupt changes in the displacement field (arrows) across faults.
A fault is a discontinuity in the velocity or displacement field associated with deformation.

Faults differ from shear fractures because a simple shear fracture cannot expand in its own plane into a larger structure. In contrast, faults can grow by the creation of a complex process zone with numerous small fractures, some of which link to form the fault slip surface while the rest are abandoned.

Geometry of faults


Normal (a), strike-slip (sinistral) (b) and reverse (c) faults. These are end-members of a continuous spectrum of oblique faults. The stereonets show the fault plane (great circle) and the displacement vector (red point).
Non-vertical faults separate the hanging wall from the underlying footwall. Where the hanging wall is lowered or down thrown relative to the footwall, the fault is a normal fault. The opposite case, where the hanging wall is up thrown relative to the footwall, is a reverse fault. If the movement is lateral, i.e. in the horizontal plane, then the fault is a strike-slip fault. Strike-slip faults can be sinistral (left-lateral) or dextral (right-lateral) (from the Latin words sinister and dexter, meaning left and right, respectively).
Although some fault dip ranges are more common than others, with strike-slip faults typically occurring as steep faults and reverse faults commonly having lower dips than normal faults, the full range from vertical to horizontal faults is found in naturally deformed rocks. If the dip angle is less than 30 the fault is called a low-angle fault, while steep faults dip steeper than 60 . Low-angle reverse faults are called thrust faults, particularly if the movement on the fault is tens or hundreds of kilometres. 

Listric normal fault showing very irregular curvature in the sections perpendicular to the slip direction. These irregularities can be thought of as large grooves or corrugations along which the hanging wall can slide.



A fault that flattens downward is called a listric fault, while downward-steepening faults are sometimes called antilistric. The terms ramps and flats, originally from thrust fault terminology, are used for alternating steep and sub-horizontal portions of any fault surface. For example, a fault that varies from steep to flat and back to steep again has a ramp-flat-ramp geometry
Irregularities are particularly common in the section perpendicular to the fault slip direction. For normal and reverse faults this means curved fault traces in map view, as can be seen from the faults of the extensional oil field. Irregularities in this section cause no conflict during fault slippage as long as the axes of the irregularities coincide with the slip vector. Where irregularities also occur in the slip direction, the hanging wall and/or footwall must deform. For example, a listric normal fault typically creates a hanging-wall rollover.

The main faults in the North Sea Gullfaks oil field show high degree of curvature in map view and straight traces in the vertical sections (main slip direction). Red lines represent some of the well paths in this field. 
A fault can have any shape perpendicular to the slip direction, but non-linearity in the slip direction generates space problems leading to hanging or footwall strain.

The term fault zone traditionally means a series of sub-parallel faults or slip surfaces close enough to each other to define a zone. The width of the zone depends on the scale of observation – it ranges from centimetres or meters in the field to the order of a kilometre or more when studying large-scale faults such as the San Andreas Fault. The term fault zone is now also used inconsistently about the central part of the fault where most or all of the original structures of the rock are obliterated, or about the core and the surrounding deformation zone associated with the fault. This somewhat confusing use is widespread in the current petroleum related literature, so any use of the term fault zone requires clarification.

A horst (a), symmetric graben (b) and asymmetric graben (c), also known as a half-graben. Antithetic and synthetic faults are shown.
Two separate normal faults dipping toward each other create a down thrown block known as a graben. Normal faults dipping away from each other create an up thrown block called a horst. The largest faults in a faulted area, called master faults, are associated with minor faults that may be antithetic or synthetic. An antithetic fault dips toward the master fault, while a synthetic fault dips in the same direction as the master fault. These expressions are relative and only make sense when minor faults are related to specific larger-scale faults.

Displacement, slip and separation 

 Illustration of a normal fault affecting a tilted layer. The fault is a normal fault with a dextral strike-slip component (a), but appears as a sinistral fault in map view (b, which is the horizontal section at level A). (c) and (d) show profiles perpendicular to fault strike (c) and in the (true) displacement direction (d).

Displacement, slip and separation

The vector connecting two points that were connected prior to faulting indicates the local displacement vector or net slip direction. Ideally, a strike-slip fault has a horizontal slip direction while normal and reverse faults have displacement vectors in the dip direction. In general, the total slip that we observe on most faults is the sum of several increments (earthquakes), each with its own individual displacement or slip vector. The individual slip events may have had different slip directions. We are now back to the difference between deformation sensu stricto, which only relates the undeformed and deformed states, and deformation history. In the field we could look for traces of the slip history by searching for such things as multiple striations.

Classification of faults based on the dip of the fault plane and the pitch, which is the angle between the slip direction (displacement vector) and the strike.
A series of displacement vectors over the slip surface gives us the displacement field or slip field on the surface. Striations, kinematic indicators and offset of layers provide the field geologist with information about direction, sense and amount of slip. Many faults show some deviation from true dip-slip and strike-slip displacement in the sense that the net slip vector is oblique. Such faults are called oblique-slip faults. The degree of obliquity is given by the pitch (also called rake), which is the angle between the strike of the slip surface and the slip vector (striation).
Unless we know the true displacement vector we may be fooled by the offset portrayed on an arbitrary section through the faulted volume, be it a seismic section or an outcrop. The apparent displacement that is observed on a section or plane is called the (apparent) separation. Horizontal separation is the separation of layers observed on a horizontal exposure or map, while the dip separation is that observed in a vertical section. In a vertical section the dip separation can be decomposed into the horizontal and vertical separation. Note that this horizontal separation is different from. These two separations recorded in a vertical section are more commonly referred to as heave (horizontal component) and throw (vertical component). Only a section that contains the true displacement vector shows the true displacement or total slip on the fault.

The relationship between a single fault, a mapped surface and its two fault cutoff lines. Such structure contour maps are used extensively in the oil industry where they are mainly based on seismic reflection data.
A fault that affects a layered sequence will, in three dimensions, separate each surface (stratigraphic interface) so that two fault cut off lines appear. If the fault is non-vertical and the displacement vector is not parallel to the layering, then a map of the faulted surface will show an open space between the two cut-off lines. The width of the open space, which will not have any contours, is related to both the fault dip and the dip separation on the fault. Further, the opening reflects the heave (horizontal separation) seen on vertical sections across the fault.

Stratigraphic separation

 (a) Missing section in vertical wells (well C) always indicates normal faults (assuming constant stratigraphy). (b) Repeated section (normally associated with reverse faults) occurs where the normal fault is steeper than the intersecting well bore (well G).
Drilling through a fault results in either a repeated section or a missing section at the fault cut (the point where the wellbore intersects the fault). For vertical wells it is simple: normal faults omit stratigraphy (figure a), while reverse faults cause repeated stratigraphy in the well. For deviated wells where the plunge of the well bore is less than the dip of the fault, such as the well G (figure b), stratigraphic repetition is seen across normal faults. The general term for the stratigraphic section missing or repeated in wells drilled through a fault is stratigraphic separation. Stratigraphic separation, which is a measure of fault displacement obtainable from wells in subsurface oil fields, is equal to the fault throw if the strata are horizontal. Most faulted strata are not horizontal, and the throw must be calculated or constructed.
Credits: Haakon Fossen (Structural Geology)

Types of fractures

What is a fracture? 

Strictly speaking, a fracture is any planar or sub-planar discontinuity that is very narrow in one dimension compared to the other two and forms as a result of external (e.g. tectonic) or internal (thermal or residual) stress. Fractures are discontinuities in displacement and mechanical properties where rocks or minerals are broken, and reduction or loss of cohesion characterizes most fractures. They are often described as surfaces, but at some scale there is always a thickness involved. Fractures can be separated into shear fractures (slip surfaces) and opening or extension fractures (joints, fissures and veins). In addition, closing or contraction fractures can be defined.

Three types of fracture.

The orientation of various fracture types with respect to the principal stresses.
Fractures are very narrow zones, often thought of as surfaces, associated with discontinuities in displacement and mechanical properties (strength or stiffness).

A shear fracture or slip surface is a fracture along which the relative movement is parallel to the fracture. The term shear fracture is used for fractures with small (mm- to dm-scale) displacements, while the term fault is more commonly restricted to discontinuities with larger offset. The term slip surface is used for fractures with fracture-parallel movements regardless of the amount of displacement and is consistent with the traditional use of the term fault. Fractures are commonly referred to as cracks in material science and rock mechanics oriented literature.
Extension fractures are fractures that show extension perpendicular to the walls. Joints have little or no macroscopically detectable displacement, but close examination reveals that most joints have a minute extensional displacement across the joint surfaces, and therefore they are classified as true extension fractures. Extension fractures are filled with gas, fluids, magma or minerals. When filled with air or fluid we use the term fissure. Mineral-filled extension fractures are called veins, while magma-filled fractures are classified as dikes. Joints, veins and fissures are all referred to as extension fractures. 
Contractional planar features (anticracks) have contractional displacements across them and are filled with immobile residue from the host rock. Stylolites are compactional structures characterized by very irregular, rather than planar, surfaces. Some geologists now regard stylolites as contraction fractures or closing fractures, as they nicely define one of three end-members in a complete kinematic fracture framework together with shear and extension fractures. Such structures are known as anticracks in the engineering-oriented literature. Rock mechanics experiments carried out at various differential stresses and confining pressures set a convenient stage for studying aspects of fracture formation. 

Experimental deformation structures that develop under extension and contraction. Initial elastic deformation is seen for all cases, while ductility increases with temperature (T) and confining pressure (Pc). YP, yield point.
Similarly, numerical modeling has added greatly to our understanding of fracture growth, particularly the field called linear elastic fracture mechanics. In the field of fracture mechanics it is common to classify the displacement field of fractures or cracks into three different modes. Mode I is the opening (extension) mode where displacement is perpendicular to the walls of the crack. Mode II (sliding mode) represents slip (shear) perpendicular to the edge and Mode III (tearing mode) involves slip parallel to the edge of the crack. Modes II and III occur along different parts of the same shear fracture and it may therefore be confusing to talk about Mode II and Mode III cracks as individual fractures. Combinations of shear (Mode II or III) fractures and tension (Mode I) fractures are called hybrid cracks or fractures. Furthermore, the term Mode IV (closing mode) is sometimes used for contractional features such as stylolites. The mode of displacement on fractures is an important parameter, for instance when fluid flow through rocks is an issue.

Mode I, II, III and IV fractures.

Extension fractures and tensile fractures 

Extension fractures ideally develop perpendicular to s3 and thus contain the intermediate and maximum principal stresses (2y ¼ 0 ). In terms of strain, they develop perpendicular to the stretching direction under tensile conditions, and parallel to the compression axis during compression tests. Because of the small strains associated with most extension fractures, stress and strain axes more or less coincide.
Joints are the most common type of extension fracture at or near the surface of the Earth and involve very small strains. Fissures are extension fractures that are more open than joints, and are characteristic of the uppermost few hundred meters of the solid crust, where they may be up to several kilometers long. 

Fissures formed in Thingvellir, Iceland, along the rift axis between the Eurasian and Laurentian plates. The fissures are open extension fractures in basalt, but the vertical displacement (right-hand side down) indicates a connection with underlying faults.
Extension fractures are typical for deformation under low or no confining pressure, and form at low differential stress. If extension fractures form under conditions where at least one of the stress axes is tensile, then such fractures are true tensile fractures. Such conditions are generally found near the surface where negative values of s3 are more likely. They can also occur deeper in the lithosphere, where high fluid pressure reduces the effective stress. Many other joints are probably related to unloading and cooling of rocks.

Shear fractures 

Shear fractures show fracture-parallel slip and typically develop at 20–30 to s1, as seen from numerous experiments under confined compression. Such experiments also show that they commonly develop in conjugate pairs, bisected by s1. Shear fractures develop under temperatures and confining pressures corresponding to the upper part of the crust. They can also form near the brittle–plastic transition, where they tend to grow into wider bands or zones of cataclastic flow. Such shear factures result in strain patterns otherwise typical for plastic deformation.

While extension fractures open perpendicular to s3, shear fractures are oblique to s3 by an angle that depends mostly on rock properties and state of stress.

Brittle and plastic deformation show different stress-strain curves (blue versus red curves): the more ductile the deformation, the greater the amount of plastic deformation prior to fracturing. It is also interesting to note the relationship between confining pressure (depth) and strain regime (contractional or extensional). The experimental data indicate that the brittle-plastic transition occurs at higher confining pressure under extension than under contraction. 

Lineations in the brittle regime

Lineations in the Brittle Regime

Some lineations occur only on fracture surfaces. They are not fabric-forming elements and are more characteristic of the brittle regime in the upper crust. These lineations form by mineral growth in extension fractures, as striations carved on the walls of shear fractures and faults, by intersections between fractures and by fracture curvatures that form early during fracturing. 
Mineral lineations in the brittle regime tend to be restricted to fiber lineations, where minerals have grown in a preferred direction on fractures. The growth of minerals on fractures usually requires that the fractures open to some extent, either as true extension fractures or as shear fractures with a component of extension. Furthermore, the minerals must grow in a preferred direction for a lineation to be defined. Minerals such as quartz, antigorite, actinolite, gypsum and anhydrite may appear fibrous on fractures. 

Two perpendicular sets of mineral lineations on a fault surface in serpentinite. The two directions indicate movements under two different stress fields. Leka Ophiolite, Scandinavian Caledonides.
Mineral fibers are found in many extensional or Mode I fractures. The orientation of the fibers is commonly taken to represent the extension direction. Curved fibers are sometimes seen, implying that the extension direction has changed during the course of deformation, or that shear has occurred during or after the formation of the fibers. 

Fiber lineation (talk) in extension fracture where the fibers have grown perpendicular to the fracture walls as the fracture opened.
Even though extension is involved in the formation of fibrous mineral lineations, this does not imply that such lineations are restricted to extension fractures. Because of the irregular shapes of most shear fractures, a component of extension may occur in extensional stepovers, and minerals may grow as walls separate. The mineral fibers then grow on the lee side of steps or other irregularities, precipitated from fluids circulating on the fracture network. Thus,the senseof slipis detectablefrom the relation between fiber growth and fracture geometry.

Formation of fiber lineation in irregular shear fractures. (a) Early stage. (b) Final stage. In (c) the upper wall is removed for inspection. Groove lineations (striations) are found on surfaces that have not opened during faulting.
Non-planar shear fractures can contain extensional (pull-apart) segments where fibers can grow and form a local lineation.

Striations or slickenlines are lineations found on shear fractures and form by physical abrasion of hanging-wall objects into the footwall or vice versa. The smooth and striated slip surface itself is called a slickenside

Slickensides with slickenlines, formed by cataclastic shearing of epidote on a post-Caledonian normal fault in the Precambrian of West Norway.
Slickensides tend to be shiny, polished surfaces coated by a 1 mm thick layer of crushed, cohesive fault rock. Hard objects or asperities can carve out linear tracks or grooves known as fault grooves. The term groove lineation can be used for this type of slickenlines. 

Fault grooves on striated fault surface separating limestone from sandstone. Moab Fault, Utah.
Such mechanically formed slickensides may show similarities with glacially striated surfaces. Close examination of many slickensided slip surfaces reveals that they are formed on mineral fill or that they actually are fiber lineations.

There are two principal types of slickenlines: those that form by mechanical abrasion (striations) and those formed by fibrous growth (slip fiber lineations).

Minerals may grow during the movement history of a fault, and it is common to find a combination of fiber lineations and striations. Such lineations may be of the same or different ages, and sometimes two or more different sets of lineations of different minerals exist that may or may not be striated due to mechanical abrasion.
Geometric striae relate to the irregular or corrugated shape of a slip surface. Such irregularities may have a preferred orientation or axis in the slip direction and appear as lineations on an exposed wall. A special type of geometric striae is the cigar-shape seen on the walls of deformation band cluster zones. Geometric striae and physical striae or slickenlines commonly coexist. 

Lineations in a zone of deformation bands. The lineation points down dip and is an expression of the cigarshaped geometry of undeformed
volumes of rock between the deformation bands. Compass for scale. Entrada Sandstone, San Rafael Desert, Utah.
Intersection lineations are found on fractures where the main slip plane is intersected by secondary fractures such as Riedel fractures or tensile fractures. The lines of intersection typically (but not necessarily) form a high angle to the slip direction, in marked contrast to striae and mineral lineations that tend to parallel the slip direction. Finally, we mention a type of lineation that is typical for some deformed limestones. It forms perpendicular to well-developed pressure solution seams where shortening occurs across the fracture surface and is composed of tubular structures known as slickolites. These structures tend to point in the direction of contraction and slickolites are thus a kind of lineation that is kinematically different from the other lineations.
Credits: Haakon Fossen (Structural Geology)

Lineations related to plastic deformation

Plastic deformation lineations

Strain data can be represented in (a) the Flinn diagram (linear or logarithmic axes) or (b) the Hsu ¨ diagram. The same data are plotted in the two diagrams for comparison. 
Penetrative lineations are found almost exclusively in rocks deformed in the plastic regime. Where the lineation forms the dominating fabric element so that the S-fabric is weak or absent, the rock is classified as an L-tectonite. It can be seen from rocks with strain markers that most L-tectonites plot in the constrictional field of the Flinn diagram, i.e. X >> Y Z (Figure above). A balanced combination of a foliation (S-fabric) and a penetrative lineation (L-fabric) is more common, and such a rock is referred to as an LS-tectonite. LS-tectonites tend to plot close to the diagonal in the Flinn diagram. S-tectonites, which have no or just a hint of linear fabric, typically plot in the flattening field of the Flinn diagram.

Mineral lineations 

A penetrative linear fabric is typically made up of aligned prismatic minerals such as amphibole needles in an amphibolite, or elongated minerals and mineral aggregates such as quartz–feldspar aggregates in gneiss. Mineral lineations can form by several processes:

Minerals and mineral aggregates can form a linear fabric by means of recrystallization, dissolution/ precipitation or rigid rotation.

Physical rotation of rigid prismatic minerals in a soft matrix can in some cases occur during deformation. An example is amphibole or epidote crystals in micaschist, where statically grown amphiboles become aligned in zones of localized deformation. In most cases the competence contrast between elongate-shaped minerals and their matrix is not high enough for rotation to be important. Instead, synkinematic recrystallization by means of plastic deformation mechanisms or a dissolution/precipitation process reshapes minerals and mineral aggregates. In addition, precipitation of quartz in pressure shadows or strain shadows is a common way to facilitate growth of minerals or mineral aggregates in a preferred direction. Even crushing or cataclasis of brittle minerals and mineral aggregates enclosed in a ductile matrix can reshape mineral aggregates to linear fabric elements.

Cataclasis, pressure solution and recrystallization all contribute to change the shape of minerals and mineral aggregates during deformation.

In a homogeneously strained rock, if a mineral aggregate had a spherical shape at the onset of deformation, its shape after deformation would represent the strain ellipsoid. In most cases the original shape is unknown so that the final shape only gives us a qualitative impression of the shape of the strain ellipsoid. Nevertheless, deformed mineral aggregates in gneisses have been used for strain analysis, although the difference in viscosity between the aggregates and their surroundings may add uncertainty to the results. In cases where the initial shape is known and the competence contrast is small, such analyses are particularly useful. Deformed conglomerates or oolites are examples of rocks where linear shapes can quantitatively be related to strain. These and other lineations defined by the shape of deformed objects are named stretching lineations and the related fabric is called a shape fabric. Stretching lineations and shape fabrics are extremely common in plastically deformed rocks such as gneisses.

Stretching lineation in quartzite conglomerate. The long axes of the pebbles are plunging to the right. The Bergsdalen Nappes, West Norway Caledonides.
Stretching of minerals and mineral aggregates into a penetrative stretching lineation forms the most common type of lineation in deformed metamorphic rocks.

Quartz, calcite and some other common minerals can grow well-aligned fibrous crystals that define linear elements. Such mineral lineations are referred to as mineral fiber lineations. Fibers may grow in the instantaneous stretching direction (ISA1), but may also grow perpendicular to the face of the opening walls. Besides, once formed they may rotate away from this direction as a result of progressive deformation. In addition to their occurrence in veins in retrograde metamorphic rocks and un-metamorphosed sedimentary rocks such as over-pressured mudstones, they are commonly found in strain shadows of porphyroclasts in metamorphic rocks during low-grade metamorphism. Fibers do not form if pressure and temperature get too high, and seldom above middle greenschist-facies conditions. This rather restricted occurrence makes fiber lineations less common and also less penetrative than many other lineation types. 
Rodding describes elongated mineral aggregates that are easily distinguished from the rest of the rock. Quartz rods are common in micaschists and gneisses where striped quartz objects occur as rods or cigars in the host rock. Rods are often considered as stretching lineations, but are commonly influenced by other structure-forming processes. They may represent isolated fold hinges, or be related to boudinage or mullion structures, or to deformed veins with an originally elongated geometry.

Intersection lineations 

Intersection lineations appearing on bedding or foliation surfaces that are intersected by a later foliation.
Many deformed rocks host more than one set of planar structures. A combination of bedding and cleavage is a common example. In most cases such planar structures intersect, and the line of intersection is regarded as an intersection lineation. Where the first tectonic cleavage (S1) cuts the primary layering or bedding (S0), the resulting intersection lineation (L1) appears on the bedding planes. Intersection lineations formed by the intersection of two tectonic foliations are also common. In most cases intersection lineations are related to folding, with the lineation running parallel to the axial trace and the hinge line. Note that for transected folds  there will be an angle between the intersection lineation and the axial trace. 
In some deformed rocks an intersection lineation appears only locally. In most cases, however, their frequency and distribution are large enough that the lineation can be considered penetrative. Like other lineations, intersection lineations may be folded about later folds, showing their use in tracing the deformation history of a rock.

Fold axes and crenulation lineations 

Fold axes are generally regarded as linear structures, despite being theoretical lines related to the geometrical shape of the folded surface. Some rocks have a high enough density of parallel fold axes that they constitute a fabric. This is often the case with phyllosilicate-rich metamorphic rocks, where small-scale folds or crenulations constitute a crenulation lineation. Crenulation lineations are thus composed of numerous millimetre to centimetre-scale fold hinges of low-amplitude folds. They are commonly seen in multiply deformed phyllites, schists and in micaceous layers in quartz-schists, mylonites and gneisses. Crenulation lineations are closely associated with intersection lineations but are different in that they are comprised of fold hinges identifiable to the naked eye. During folding of layered rocks crenulation cleavages and crenulation lineations form at an early stage, while larger folds form later on during the same process. It is therefore of interest to compare the orientation of early crenulation lineations with related but slightly younger fold axes. If there is a difference in orientation this could be related to how layers rotated during deformation. We are already familiar with the concept of transected folds, where the lineation makes an angle to the hinge line.

Cleavage transecting the axial surface of a transected fold.
Boudinage 

Cylindrical pinch-and-swell structures (above) and boudins (below) represent linear elements in many deformed rocks.
Boudins are competent rock layers that have been stretched into segments. Individual boudins are commonly much longer in one dimension than the other two and thus define a lineation. Such linear boudins form where the X-axis of the strain ellipsoid is significantly larger than Y. Chocolate-tablet boudins can form when X Y. When occurring in folded layers, boudins typically appear on the limbs of the fold with their long axes oriented in the direction of the fold axis. 

Common connection between folding and boudinage. The fold hinges are thickened while the limbs are extended and boudinaged. The strain ellipse is indicated.
In general, boudinage structures are most easily recognized in sections perpendicular to the long axes of the boudins. Because of this fact they may be difficult to recognize as linear features in deformed rocks. It is also true that boudins are restricted to competent layers and therefore more restricted in occurrence than most other lineations.

Mullions 

Mullion is the name that structural geologists use for linear deformation structures that are restricted to the interface between a competent and an incompetent rock. The term mullion has been used in several different ways in the literature, ranging from striations on fault surfaces (fault mullions) to layer-interface structures formed during layer-parallel extension as well as contraction. We will relate the term to layer-interface structures where the viscosity contrast is significant. In such cases the cusp shapes of mullions always point into the more competent rock, i.e. the one with the higher viscosity at the time of deformation. 

Mullion structures form lineations at the interface between rocks of significantly different competence (viscosity).
Such mullions are closely related to buckle folds in the sense that their formation is predicted by a contrast in viscosity, they form by layer parallel shortening, and their characteristic wavelength is related to the viscosity contrast. But they differ from buckle folds in having shorter wavelengths and they are restricted to a single layer interface. A common place to find mullion structures in metamorphic rocks is at the boundary between quartzite and phyllite or micaschist. Mullions also occur on the surface of quartz pods in micaschists.

Pencil structures 

The formation of pencil structures occurs as a result of discrete interference between compaction cleavage and a subsequent tectonic cleavage, or between two equally developed tectonic cleavages. Pencil structures have a preferred orientation and form a lineation in un-metamorphosed and very low-grade metamorphic rocks.
Credits: Haakon Fossen (Structural Geology)

Cleavage development

Cleavage Development

There are many types of cleavages and a rich terminology is available. To efficiently deal with cleavages and foliations it is useful to keep an eye on crustal depth and lithology. Crustal depth is related to temperature (and pressure), and with increasing temperature we first obtain increasing mobility of minerals and at yet higher temperatures the possibility that minerals will recrystallise. Around 350–375 C we leave the realm of cleavage and enter that of schistosity and mylonitic foliations. Lithology and mineralogy are important because different minerals react differently to stress and temperature. Phyllosilicates are particularly important in cleavage development. In general, if there are no phyllosilicates in the rock there will not be a very strong cleavage or schistosity. Cleavage formation in calcareous rocks is controlled by the mobility of carbonate and the easy formation of stylolites.

Cleavage is the low-temperature version of foliation and is best developed in rocks with abundant platy minerals.

We will here consider the most common types of foliation that develop due to deformation during prograde metamorphism, i.e. as a rock is being buried to progressively greater depths.

Compaction cleavage 

The first secondary foliation forming in sedimentary rocks is related to their compaction history. Reorientation of mineral grains and collapse of pore space result in accentuation and reworking of the primary foliation (bedding). For a clay or claystone, the result is a shale with a marked compaction cleavage

Theoretical cleavage development in a mudstone.
In this process there is also dissolution going on, and in some quartzites we can find pressure solution seams. Such structures are much more common in limestone, in which dissolution of carbonate produces sub-horizontal and irregular seams where quartz or carbonate has been dissolved and where clay and other residual minerals are concentrated. The seams are stylolites or pressure solution seams, and the foliation can be called a stylolitic cleavage, which is a type of (pressure) solution cleavage. The spacing of the seams in calcareous rocks is usually several centimetres, and the cleavage is therefore a widely spaced cleavage. In fact, the stylolitic surfaces may be too far apart to define a cleavage. In contrast, the compaction cleavage in shale is recognizable under the microscope and therefore a continuous cleavage. These non-tectonic cleavages are usually regarded as S0 foliations.

Early tectonic development and disjunctive cleavage 

A tectonic foliation commonly results when a sedimentary rock is exposed to tectonic stress that leads to progressive horizontal shortening of sedimentary beds - a condition typical of the foreland regions of orogenic belts. In limestone and some sandstones, the first tectonic foliation to form is a pressure solution cleavage that typically is stylolitic (toothlike, showing a zigzag suture in cross section). If s1 is horizontal, a vertical pressure solution cleavage forms that will make a high angle to S0 and previously formed compaction-related stylolites. 
Pressure solution is also important when shales are exposed to tectonic stress. In this case extensive dissolution of quartz causes concentration and reorientation of clay minerals. At some point the secondary cleavage will be as pronounced as the primary one, and clay minerals will be equally well oriented along S1 and S0. The shale will now fracture along both S1 and S0 into pencil-shaped fragments, which explains why the cleavage is known as pencil cleavage. Pencil cleavage also occurs where two tectonic cleavages develop in the same rock due to local or regional changes in the stress field. Such purely tectonic pencil cleavage is associated with some thrust ramps, formed close in time during the same phase of deformation. 

Pencil cleavage in shale in the Caledonian foreland fold-and-thrust belt near Oslo.
If the tectonic shortening persists, it will eventually dominate over the compactional cleavage. More and more clay grains become reoriented into a vertical orientation as quartz grains are being dissolved and removed, somewhat similar to a collapsing house of cards. Micro-folding of the clay grains may also occur. The result is a continuous cleavage that totally dominates the structure and texture of the rock. The rock is now a slate and its foliation is known as slaty cleavage
The formation of slaty cleavage occurs while the metamorphic grade is very low, so that recrystallization of clay minerals into new mica grains appears to have just started. A close look at a well-developed slaty cleavage reveals that a change has taken place in terms of mineral distribution. There are now domains dominated by quartz and feldspar, known as QF-domains, that separate M-domains rich in phyllosilicate minerals. The letters Q, F and M relate to quartz, feldspar and mica, and we need a microscope to discern the individual domains, which are considerably thinner than 1 mm. The QF-domains are typically lozenge- or lens-shaped while the M-domains form narrower, enveloping zones. As shown below, many types of cleavages and foliations show such domainal structures, and they can all be referred to as domainal cleavages

Disjunctive cleavage types. Stylolitic (limestones) and anastomosing (sandstones) cleavages are usually spaced, while continuous cleavages in more fine-grained rocks are separated into rough and smooth variants, where the rough cleavage can develop into the smooth version. All disjunctive cleavages are domainal, and the cleavage domains (C) are separated by undeformed rock called microlithons (M).
The term disjunctive cleavage is commonly used about early tectonic domainal cleavage in previously unfoliated rocks such as mudstones, sandstones and limestones. This term implies that the cleavage cuts across, rather than crenulating (folding), pre-existing foliations.
It was once thought that slaty cleavage formed by physical grain rotation. We now know that so called wet diffusion or pressure solution is what chiefly produces the domainal structure that characterizes slaty cleavage. Grains of quartz and feldspar are dissolved perpendicular to the orientation of the cleavage and achieve lensoid shapes (disk shapes in three dimensions). Where this happens, phyllosilicates are concentrated and M-domains form. The importance of dissolution or pressure solution allows us to use the term (pressure) solution cleavage also for slaty cleavage.

Cleavage forms through grain rotation, growth of minerals with a preferred direction and, most importantly, wet diffusion (pressure solution) of the most solvent minerals in the rock.

Wet diffusion implies that dissolved minerals diffuse away through a very thin film of fluid located along grain boundaries. The material is precipitated in so-called pressure shadows of larger and more rigid grains in the QF-domains or becomes transported out of the rock. In fact, very significant amounts of matter appear to have left most slate belts, which represent interesting aspects in terms of thermodynamics and fluid flow through the upper crust.


Greenschist facies: from cleavage to schistosity 

New phyllosilicate minerals grow at the expense of clay minerals in shales and slates when they enter the field of green schist facies metamorphism. A phyllite forms and the cleavage changes into a phyllitic cleavage. The new mica minerals grow with their basal plane more or less perpendicular to the Z-axis of the strain ellipsoid, and more or less perpendicular to s1. The newly formed mica grains are thus parallel and a phyllitic cleavage is established. The cleavage is still a continuous one, and the development of QF- and M-domains is more pronounced than for slaty cleavage. The domainal cleavage becomes better developed because dissolution (wet diffusion) becomes more efficient as green schist facies temperatures are reached. 

Phyllitic cleavage bears similarities to crenulation cleavage when viewed under the microscope. The difference is that phyllitic crenulations are microscopic and thus invisible to the naked eye. In this case a phyllitic cleavage dies out towards a folded competent lamina.
When original claystone reaches upper green schist facies and perhaps lower amphibolite facies, the mica grains grow larger and become easily visible in a hand sample. At the same time, the foliation becomes less planar, wrapping around quartz–feldspar aggregates and strong metamorphic minerals such as garnet, kyanite and amphibole. The foliation is no longer called a cleavage but a schistosity, and the rock is a schist

(a) Phyllitic cleavage in lower greenschist facies phyllite. (b) This cleavage formed higher into the greenschist facies, showing very well-developed QF- (middle) and M-domains and coarser grain size.
Schistosity is also found in quartz-rich rocks such as quartz schists and sheared granites. Here the M- and QF-domains are on the millimetre or even centimetre scale and they appear more regular and planar than for mica schists. This is why quartz schists and sheared granites split so easily into slabs that can be used for various building purposes. In summary, while wet diffusion (solution) and grain reorientation dominate the formation of slaty cleavage, recrystallization is more important during the formation of schistosity.

Secondary tectonic cleavage (crenulation cleavage) 

An already established tectonic foliation can be affected by a later cleavage (S2 or higher) if the orientation of the ISA changes locally or regionally at some point during the deformation, or if a later cleavage-forming deformation phase occurs. Because cleavages tend to form perpendicular to the maximum shortening direction (X), a new cleavage will form that overprints the preexisting one. In many cases this occurs by folding the previous foliation into a series of microfolds, in which case the cleavage is called a crenulation cleavage. Hence, a crenulation cleavage is a series of micro-folds at the centimetre scale or less with parallel axial surfaces. Depending on the angle between the existing foliation and the secondary stress field, the crenulation cleavage will be symmetric or asymmetric. 

Asymmetric crenulation cleavage affecting a mylonitic foliation. The cleavage is discrete in the middle, micaceous layer, while it is zonal and less well developed in the more quartz-rich adjacent layers.
A symmetric crenulation cleavage has limbs of equal length, while an asymmetric crenulation cleavage is composed of small, asymmetric folds with S- or Z-geometry. 
Crenulation cleavage through which the earlier foliation can be traced continuously is known as zonal crenulation cleavage. In the opposite case, where there is a sharp discontinuity between QF- and M-domains, the cleavage is called a discrete crenulation cleavage. The M-domains here are thinner than the QF-domains and mimic micro-faults. Discrete and zonal crenulation cleavages can grade into each other within a single outcrop. 
Crenulation cleavage is restricted to lithologies with a pre-existing well-developed foliation that at least partly is defined by phyllosilicate minerals. It is commonly seen in micaceous layers while absent in neighbouring mica-poor layers. The domainal thickness of the affected foliation is connected with the wavelength of the new crenulation cleavage: thicker domains produce longer crenulation wavelengths. This is the same relationship between layer thickness and wavelength, where the viscosity contrast was shown to be important. A close connection is also seen between crenulation cleavage and folding. 

Crenulation cleavage affecting the phyllitic cleavage. The crenulation cleavage is seen to be axial planar to decimetre-scale folds.
We can find any stage of crenulation cleavage development, from faint crenulation of foliations to intense cleavage development where recrystallization and pressure solution have resulted in a pronounced domainal QF-M-structure.In the latter case the original foliation can be almost obliterated in a hand sample although usually observable under the microscope. Progressive evolution of crenulation cleavage is accompanied by progressive shortening across the cleavage, and eventually a crenulation cleavage can transform into a phyllitic foliation.

Courtesy © Text and Photos from Structural Geology (book) by Haakon Fossen 


Folding: mechanisms and processes


Every geologist mapping or describing folds in the field probably has the same question in mind: how did these structures actually form? As geologists we tend to look for a simple history or mechanism that can explain our observations reasonably well. Folding is no exception, and there are different approaches and process-related terms. One approach is to consider the way force or stress acts on a layered rock, which leads to the three-fold classification and terminology. Other terms are related to how the layer(s) react to force and stress, for instance whether layers fold by layer-parallel shearing, orthogonal flexure or some other mechanism that is controlled by rock rheology. Still other classes of folding, such as kinking and chevron folding, are related to fold geometry. For this reason, several different fold mechanisms are defined, and many of them overlap in definition. This is why terms such as buckling, kink folding and bending can be confusing when discussed in terms of mechanisms such as flexural slip and simple shear. In summary, we are dealing with differences in orientation of stress axes relative to the layering, kinematics, and mechanical and rheological properties, and thus mechanisms that emphasize different aspects of folding. The most important distinction between the ways folds form probably lies in whether the layering responds actively or passively to the imposed strain field. Now considering active folding (buckling), where the competence or viscosity contrast between the folding layer and its host rock is important. We will then look at passive folding, where layers are simply passive markers with no rheological influence, and then consider bending, where forces are applied across the layering. The following sections will then discuss models known as flexural folding mechanisms (flexural slip, flexural shear and orthogonal flexure), which can contribute to both active folding and bending. Finally, we will discuss kinking and the formation of chevron folds.

Active folding or buckling (Class 1B folds)


Active folding or buckling is a fold process that can initiate when a layer is shortened parallel to the layering, as shown schematically. Folds appear to have formed in response to layer-parallel shortening. A contrast in viscosity is required for buckling to occur, with the folding layer being more competent than the host rock (matrix). The result of buckling is rounded folds, typically parallel and with more or less sinusoidal shape.

Two folded layers of different thickness. The upper and thinner one shows a smaller dominant wavelength than the lower one.
Buckling occurs when a competent layer in a less competent matrix is shortened parallel to the length of the layer.

If an isotropic rock layer has perfectly planar and parallel boundaries and is perfectly parallel with a constantly oriented s1 or ISA1, then it will shorten without folding even though there is a significant viscosity contrast between the layer and the host rock. However, if there are small irregularities on the layer interfaces, then these irregularities can grow to form buckle folds with a size and shape that depend on the thickness of the folded layer and its viscosity contrast with its surroundings.

Buckling or active folding implies that there is layer parallel shortening and a viscosity contrast involved, and also irregularities on which folds can nucleate.

Buckling of single, competent layers in a less competent matrix is relatively easy to study in the laboratory and has also been explored numerically. Single-layer folds formed by buckling have the following characteristics:

Alternating Class 1B and 3 folds are commonly seen in folded layers. CompetentlayersexhibitClass1Bgeometry.
  • The fold wavelength–thickness ratio (L/h) is constant for each folded layer if the material is mechanically homogeneous and if they were deformed under the same physical conditions. Such folds are often called periodic folds. If the layer thickness varies, then the wavelength is changed accordingly.  
  • The effect of the folding disappears rapidly (about the distance corresponding to one wavelength) away from the folded layer.   
  • The folds in the competent layer approximate Class 1B folds (constant layer thickness). If there are two or more folded competent layers then the incompetent layers in between are folded into Class 1A and Class 3 folds. Cusp (pointed) hinges point to the more competent layers.  
Strain distribution in the hinge zone of a folded limestone layer in shale. Outer-arc stretching is separated from inner-arc shortening by a neutral surface.
  • The outer part of the competent layer is stretched while the inner part is shortened. The two parts are typically separated by a neutral surface. Note that layer-parallel shortening, which always takes place prior to folding, can reduce or eliminate the outer extensional zone.  
  • The normal to the axial surface or axial cleavage indicates the direction of maximum shortening (Z).
Experiments and theory show that homogeneous shortening (T) occurs initially, together with the growth of irregularities into very gentle and long-amplitude fold structures. When the most accentuated folds achieve opening angles around 160–150 , the role of layer-parallel shortening decays. From that point on the folds grow without any significant increase in layer thickness. 
Buckling has been modeled under the assumption of linear or Newtonian viscosity. It is likely that most rocks show non-linear rheological behavior during plastic deformation, which has consequences for the buckling process. A power-law rheology is then assumed, where the exponent n > 1. The higher the n-exponent, the quicker the fold growth and the less the layer parallel shortening T. Many natural folds show low T-values, and, together with low L/h ratios (L/h < 10), this indicates a non-linear rheology. However, the differences between the results from viscous and power-law rheology models are not great. Buckle folds are most easily recognized as single competent layers, but can also occur where several competent layers occur in parallel. Ld/h is significantly less for multilayer than for single layer buckling. Where two thin layers are close they will behave more like a single layer whose thickness is the sum of the two thin layers, as seen from the experimental results.
Folding of multi-layered rocks. Far-apart layers act as individual layers (left). The closer they get, the more they behave as a single layer with thickness larger than that of the thickest of the individual layers.
Where we have alternating thick and thin layers, the thin layers will start to develop folds first. At some point the thick layers will start to fold (with longer wavelength) and take control over the further development. The result is relatively large folds controlled by thick layers together with small, second-order folds formed earlier in the process.
Illustration of how folding initiates in thin layers. Once the thicker layer starts to fold, the smaller folds in the thin layer become parasitic and asymmetric due to flexural flow.
Several mechanisms can be involved during buckling. The simple stones can collectively be termed flexural folding and are separated into orthogonal flexure, flexural slip and flexural flow. In addition there is always the possibility of having volume change, particularly in the hinge zone.
Buckled multi-layers. Note how the largest folds affect the entire layer package.
Passive folding (Class 2 folds) 

Passive folding is typical for rocks where passive flow occurs, i.e. where the layering exerts no mechanical influence on the folding. In these cases the layering only serves as a visual expression of strain with no mechanical or competence contrast to neighbouring layers. Such layers are called passive layers. Perfectly passive folds produced by simple shear are Class 2 (similar) folds, and passive folds that are associated with simple shear, or atleast a significant component of simple shear, are called shear folds.

Formation of Class 2 folds by (a) simple shearing and (b) pure shearing of a gently curved layer. No viscosity contrast is involved, meaning that the folds can be regarded as passive.
Passive folds generated by simple shearing are perfectly similar folds.

Passive folds of perfect Class 2 geometry can easily be generated by differentially shearing a card deck. Drawing lines perpendicular to the cards prior to shearing helps visualize the fold. However, the formation of passive folds is not restricted to simple shear. Passive folds can form in response to any kind of ductile strain, for instance sub-simple shear, transpression and even coaxial strain. Hence simple shear is only one of an infinite spectrum of kinematic models that can produce passive folds.

Passive folding produces harmonic folds where the layering plays no mechanical role and therefore no influence on the fold shape. 

Examples of passive folding are found where passive layers enter shear zones or otherwise are affected by heterogeneous strain. Drag folds along faults are examples typical for the brittle regime, although many layered sequences contain beds of quite different mechanical properties so that slip occurs between layers. Passive folds are frequently found in mylonite zones, particularly in mono-mineralic rocks such as quartzite, marble and salt.
Passive harmonic folding of quartzite in a Caledonian mylonite zone. The similar geometry of this Z-fold and its setting in a Caledonian shear zone indicate that it is a shear fold.
Bending 

Bending occurs when forces act across layers at a high angle, unlike buckle folds where the main force acts parallel to a layer. This is also the case for passive folding, and the two are closely related. However, bending is generally thought of as something that is more directly forced upon the layers by geometries and kinematics of the bounding rock units. Several aspects of bending have been studied in great detail by engineers because of its importance in the field of construction engineering, such as in horizontal beams supported by vertical pillars.

Bending occurs when forces act across layers, and may involve more than one mechanism.

Classic geologic results of bending are the forced folds created in sedimentary layers blanketing faulted rigid basement blocks. The displacement is forced upon the sediments by fault movement on a pre-existing basement fault, and the sediments are soft enough to respond by monoclinal folding until at some critical point they rupture and the fault starts propagating up-section. Such structures are particularly well exposed in the Colorado Plateau–Rocky Mountains area, where numerous Laramide-related uplifts have created such structures. Bending as such is a boundary condition- or external load-related model, not a strain model, particularly not when a free surface is involved such as during forced folding mentioned above. In other words, there are many ways that folding and strain can accumulate internally in a fold during bending.

Examples of bending in various settings and scales: (a) between boudins; (b) above thrust ramps; (c) above reactivated faults; and (d) above shallow intrusions or salt diapirs.
An obvious response to bending is deformation by simple shear, in which case we are back to passive folding. The simple shear passive folding model may work if we have a wide fault zone underneath the fold or if the fold is very narrow. In most cases the fold widens upward, telling us that we have to modify the simple shear model. In this case trishear comes in handy. Trishear distributes shear in a triangular zone ahead of a propagating fold, and seems to work very well for several mapped examples. Still, trishear cannot explain all features seen in many forced folds. Field studies show evidence of bedding parallel slip or shear. This is manifested by striations on weak bedding-parallel surfaces or by bedding-parallel deformation bands. Also the related flexural mechanism described below as orthogonal flexure can result from bending loads. 
There are many other examples of bending. One is fault-bend folds, for instance where thrust sheets are passively bent as they move over a ramp structure. Such folds are commonly modelled as kink folds, again related to flexural slip. They may also be modelled by means of simple shear, which is commonly done for fault-bend folds formed above non-planar (e.g. listric) faults. 
Differential compaction, where a sedimentary sequence compacts more in one area than in another due to different degrees of compaction of the underlying layers, is also a type of bending. This is common across the crests of major fault blocks in post rift-sequences in sedimentary basins, but can also occur along salt diapirs and shallow intrusions. Folds formed by differential compaction are gentle. 
Forceful intrusion of magma or salt can also bend roof layers. Again the strain accumulation mechanism may vary, with flexural slip being a common constituent.
In the plastic regime, bending is less common because of the high ductility of all or most parts of the deforming rocks. However, bending is frequently associated with rigid boudins.

Passive folding of layers between boudins.
Flexural slip and flexural flow (Class 1B) 

(a) Flexural slip, showing opposite sense of slip on each limb, decreasing towards the hinge zone. (b) Flexural flow, where fold limbs are being sheared. Ideally, layer thickness is preserved in both models.
Flexural slip implies slip along layer interfaces or very thin layers during folding. It is one of three kinematic models of folding (the others being flexural flow and orthogonal flexure) that maintains bed thickness and thus produces Class 1B or parallel folds. Simple flexural slip experiments can be performed simply by folding double sandwiches with jelly. The sandwich maintains its thickness even though slip occurs between the pieces of bread, until the fold becomes too tight. It is a prerequisite for flexural slip that the deforming medium is layered or has a strong mechanical anisotropy. 
In nature, the anisotropy could be mica-rich thin layers in a quartzite or mylonite, or thin shale layers between thicker sandstone or limestone beds in sedimentary rocks. Flexural slip can occur in the middle crust where plastic deformation mechanisms would be involved, but is perhaps more common where sedimentary strata are folded in the upper crustal brittle regime. In the latter case, bedding surfaces act like faults, and slicken lines will sometimes develop on slipping surfaces. 
Maximum slip occurs at the inflection points and dies out toward the hinge line, where it is zero. The sense of slip is opposite on each limb, and the slip is consistent relative to the hinges, where sense of slip changes. Relative slip on the convex side of a flexural slip fold is always toward the fold hinge, whereas on the concave side slip is opposite.

Slicken lines on folded weak layers and constant bed thickness reveal flexural slip.

In cases where strain is more evenly distributed in the limbs in the form of shear strain, as is more commonly the case in the plastic regime, flexural slip turns into the closely related mechanism called flexural shear or flexural flow. Flexural flow experiments are conveniently done by bending a soft paperback book or a deck of cards (remember to draw circles for strain markers). During this process slip occurs between individual paper sheets. If we put strain markers on our paperback, we would see that strain is zero in the hinge zone and increasing down the limbs. This is so because the shear strain is directly related to the orientation (rotation) of the layers: the higher the rotation, the higher the shear strain.
For originally horizontal layers folded into an upright fold, shear strain is directly related to dip (g ¼ tan (layer dip)), and the sense of shear is opposite on each side of the axial trace. This results in a characteristic strain distribution in the fold. For example, the neutral surface separating extension from contraction, typical for many buckle folds, is not found in pure flexural-flow folds. Flexural flow produces identical strain in the inner and outer part of a fold, but strain increases away from the hinge. Note that evidence for a combination of orthogonal flexuring and some flexural flow or slip is commonly found in buckle folds, in which case a neutral surface may well exist.

Pure flexural folds have no neutral surface, and strain increases away from the hinge zone.

Pure flexural folds are perfect Class 1B folds. We can estimate the amount of layer-parallel shortening for such folds by measuring the length of any one of the folded layers. This layer has maintained its original length because it was the shear plane throughout the folding history. Constant layer length and thickness are assumptions that simplify restoration of cross-sections.

Orthogonal flexure (also Class 1B) 

Orthogonal flexure, also called tangential longitudinal strain, is a deformation type with its own specific conditions:

All lines originally orthogonal to the layering remain so throughout the deformation history.

The result is stretching of the outer part and shortening of the inner part of the folded layer. The long axis of the strain ellipse is therefore orthogonal to bedding in the inner part of the layer and parallel to bedding in the outer part.

Layer-parallel shortening resulting in orthogonal flexure and flexural flow. Note what happens to the originally orthogonal lines. Strain ellipses are indicated.
Orthogonal flexure and flexural flow have in common that they produce parallel (Class 1B) folds. But the two models produce quite different strain patterns: The neutral (no strain) surface separating the outer-extended and inner-contracted part of the folded layer does not exist in flexural flow, where strain is identical across the fold along dip isogons. During the folding history, the neutral surface moves inward toward the core of the fold, which can result in contraction structures overprinted by extension structures.

Orthogonal flexure produces parallel folds with a neutral surface.

Pure orthogonal flexure is only possible for open folds. When folds get tighter, the conditions for orthogonal flexure become harder and harder to maintain, and flexural slip or flow will gradually take over. Evidence for orthogonal flexure is typically found in stiff, competent layers that resist ductile deformation. Some have simplified the definition of orthogonal flexure to a mechanism resulting in outer-arc contraction and inner-arc extension. By getting rid of the requirement of orthogonality, the model becomes more general and embraces many more natural examples.

Kinking and chevron folding 

(a) Conjugate kink bands in mylonitized anorthosite gabbro, Bergen, Norway. (b) Kink folds related to Laramide thrusting in north Wyoming (Dead Indian Summit) (c) Kink-like folds in oceanic sediments in Oman.
Kink bands are common in well-laminated and anisotropic rocks rich in phyllosilicate minerals, and some field occurrences. Kink bands are centimetre to decimetre wide zones or bands with sharp boundaries across which the foliation is abruptly rotated. Wider zones are sometimes referred to as kink folds. Kink bands and kink folds are characterized by their strong asymmetry and their Class 2 fold geometry. They are closely related to chevron folds, which also are Class 2 folds, but differ in terms of symmetry. Both are relatively low-temperature (low metamorphic
grade) deformation structures where there is a significant mechanical anisotropy represented by lamination or repeated competent–incompetent layers, and both imply layer shortening. Classic kink bands have very angular hinges and lack even the narrow hinge zone found in the outer arc of chevron folds. There is another important difference between the two. While chevron folds initiate with their axial surface perpendicular to the shortening direction, kink bands form oblique to this direction, typically in conjugate pairs.

(a) The orientation of s1 can be determined from the orientation of conjugate sets of kink bands.
(b) Continued kink band growth can produce chevron folds.
When conjugate sets of low-strain kink bands are observed, s1 or ISA1 is commonly assumed to bisect the sets. Going from strain to stress is not straightforward, but the smaller the strain the better the correlation. When a single set of kink bands occurs, we know that s1 is oblique to the band, but its precise orientation is unknown because kink bands may rotate during progressive deformation. Kink folds generated by bending do not directly reveal the orientation of stress. Such kink folds have orientations that are controlled by the local geometries of ramps or fault bends. Hence, in such cases the bisecting axis between two kink zones does not in general represent s1 or ISA1. Experiments have shown that conjugate sets can nicely merge to form chevron folds if strain is high enough (around 50%). However, 50% shortening is not commonly achieved by kinking in naturally deformed rocks, so this way of forming chevron folds may not be the most common one after all. Classic chevron folds with beds on the centimetre scale are more likely to form by flexural slip of multi-layered rocks during layer-parallel shortening. The typical setting is where competent beds are separated by thin incompetent layers, for instance quartzite or chert separated by shale or phyllite. Flexural slip then occurs between the competent layers, which are strained only in the thin hinge zones. Just like buckle folds, the hinges have to stretch in the outer arc and shorten in their inner parts. Furthermore, geometric problems in the hinge zone require flow of the incompetent rock into the hinge, or alternatively inward hinge collapse of the competent bed. Hinge collapse is particularly common in relatively thick competent layers that occur between thinner ones. Another way of resolving hinge compatibility problems is by reverse faulting.
Credits: Haakon Fossen (Structural Geology)