Kinematic analysis includes four components of deformation

- translation (change in position)
- rotation (change in orientation)
- dilation (change in size) and
- distortion (change in shape)

Translation and rotation are grouped as** rigid body deformation**, whereas dilation and distortion constitute **strain.**

In practice, dilation is very difficult to measure in most rocks, and so normally when we speak of strain we are speaking of distortion.

Many of the features we see in deformed rocks are products of strain. It is important to have some understanding of strain to answer questions like:

- How much shortening has there been in the Cordillera?
- How is fabric produced by deformation?
- How can we figure out original thicknesses of strata affected by folds and cleavage.
- When we see a major mylonite zone, how much movement has there been
- Can we trust sedimentary structures and paleocurrent directions measured in deformed rocks?
- How much distortion accompanied folding in a foothills oilfield, and how did this affect porosity?

All these are questions that can only be answered with an understanding of strain.

Many rocks contain evidence for strain. To actually measure strain we need **strain markers**: features that had an original shape or size that we can guess, so we can use them to estimate distortion.

**Homogeneous strain** is strain that produces the same distortion everywhere. As a result, straight lines stay straight, and parallel lines stay parallel.

**Heterogeneous strain** is involves distortion or dilation that varies from place to place. Straight lines do not in general stay straight; parallel lines do not in general stay parallel.

When strain is systematically heterogeneous between different components of a rock (e.g. garnets are rotated but quartz is flattened) we describe the strain as **partitioned**.

Homogeneous strain is much easier to deal with and we will start our study of strain with homogeneous strain. Where we encounter heterogeneous strain, we will still be able to use the concepts and terminology of homogeneous strain by either

- looking at a single point or very small element, or
- by averaging out all the variations and looking at the
__bulk strain__

Another phenomenon that helps us understand deformation is strain compatibillity. When rocks deform by ductile processes, it is unusual for large empty spaces to open up within rocks, because of overburden pressure. Hence, the strained portions of a rock have to fit together without leaving any spaces. This limits the types of strain heterogeneity that can occur.

Clearly, not all deformed rocks satisfy strain compatibility. A fault breccia, for example, may contain large open spaces. On the other hand, at large scale, strain compatibility is an important concept in the balancing of cross-sections through thrust belts, so it has applications in both brittle and ductile deformed rocks.

Everything we have said so far compares a final state to an initial state. Typically, if we look at deformed objects as guides to strain, that's what we are doing. The strain involved is called the finite strain.

In getting from starting state A to final state B, a rock passes through an infinite series of very small steps. There are infinitely many routes that could be followed.

Like most physical processes it is sometimes helpful to look at a very small part of the strain history, which is referred to as the **incremental strain**. Incremental strains may be represented by symbols such as δe or δγ

In strain theory, we sometimes consider an infinitesimally small increment of strain, which is referred to as **instantaneous strain **or **infinitesimal strain**, effectively the derivative of the finite strain.

For example, if finite strain is represented by an extension *e* or a shear strain* γ* then an instantaneous strain is *de* or *dγ*.

The finite strain is the integral of the the infinitesimal strain, from the time of the formation of the rock to the present day.

*de/dt* and *dγ/dt* (also written
$\dot{e}$,
$\dot{\gamma}$
) are** strain rates** - measures of the speed at which rocks flow. Strain itself is a dimensionless quantity, a ratio of two lengths. Therefore the units of strain rate are just "per second" or *s ^{-1}*. Sometimes structural geologists will say "strains per second". Typical geological strain rates are of the order of 10

In dealing with strain in one dimension we usually focus on **material lines**. A material line can be thought of as a row of sand grains, or maybe a row of calcium atoms in a calcite crystal. It is a line of material particles (as opposed to a line that is just part of a geometrical construction.)

**Longitudinal strain** means change in length. Suppose a material line starts off with length *l _{0 }*and ends up with length

Extension *e* is also known as fractional change in length *e = (l-l _{0})/l_{0}*

The stretch is the new length as a proportion of the old *s = l/l _{0} = 1+e*

*λ = s ^{2}*

*ε = ln(s)*

There is a second way that strain affects material lines. A line is subject to **shear strain** if adjacent particles of rock flowed past it in a parallel direction. For example, shear strain is occurring at present along the western margin of the North American plate as the Pacific Plate moves northward relative to the North American Plate

The easiest way to represent the shear strain along a line is to look for an** ****originally perpendicular material marker line** (e.g. in a deformed fossil__)__

After deformation we look at how much that marker has rotated out of perpendicular. Again there are several ways to represent the shear strain:

The angle is the **angle of shear** represented by the Greek letter psi *ψ*

**Shear strain** (sometimes called engineering shear strain, is represented by the Greek letter gamma *γ* *=* tan *ψ*

For some purposes it is easier to use the **tensor shear strain** represented by *γ*/2

Once we move into two dimensions, things are a bit more complicated. There is an infinite number of differently oriented material lines in a surface. Lines with different orientations will show different values of *s* and *γ*

The most intuitive way to represent strain in 2D is to imagine a circle before deformation, and to look at its shape after deformation. In general (for **homogeneous strain**) the circle will become an ellipse - the **strain ellipse.**

The strain ellipse is the product of a finite strain applied to a circle of unit radius. It is an ellipse whose radius is proportional to the **stretch** *s* in any direction.

A deformed circular object has the same shape (though not, strictly, the same size) as the strain ellipse.

If we superimpose the strain ellipse on the original unit circle, we can separate a **field of finite extension** from a **field of finite shortening**. These fields are separated by two **lines of no finite extension** or **LNFE**

Depending on the dilation and the shape of the strain ellipse the LNFE may be closer to the long axis or the short axis. If the strain involves a lot of dilation (negative or positive) then all lines may have positive or negative *s*, and there is no LNFE.

Like any other ellipse, there are two other special lines, the longest and shortest radii, known as **strain axes**, labelled X and Z or alternatively *s _{1}* and

The strain axes have another special property - the are perpendicular before and after deformation. So, they must be lines of zero shear strain. In general, they are the only two lines along which the shear strain is zero.

Describing a strain in 2D

Note that it takes **3 numbers** to describe a finite strain in 2D: *s _{1, }s_{3}*, and the orientation of a strain axis (typically represented

The **strain ratio*** R _{s} = s_{1}/ s_{3}* is called the strain ratio and is a measure of the overall intensity of distortion.

**Dilation** *Δ* is the fractional change in area; it is positive if the area increases, negative if it decreases.

The product* s _{1} x s_{3}*

1 + *Δ* = *s _{1} x s_{3}*

Notice that if we want to extend our description to include the rotational part of deformation (relative to some material reference line like the trace of a horizontal bedding plane or the edge of a shear zone) then we need another measurement - this might be obtained from the orientation of the strain axis before and after deformation (*φ, φ'*).

The difference *ω = φ'- φ* is a measure of how much overall rotation has occurred during deformation.

Under some circumstances, we may envisage that the strain axes have the same orientation (relative to a reference direction) before and after deformation. Such a deformation is known as a** pure strain** because the rotation is zero. It is also sometimes called a **non-rotational** or **irrotational deformation.** There is a special type of pure strain, known (confusingly) as **pure shear** in which the dilation is also zero. Then we can say that

1 + *Δ* = *s _{1} x s_{3}* = 1

and therefore

All other types of deformation are termed **rotational**. In such deformations the strain axes rotate during deformation.

In general, it is always possible to think of a finite rotational strain as a combination of a non-rotational, pure strain, combined with a simple rotation. (Remember though, that because we are dealing with finite strain, it's not possible using the techniques of this section to say whether the rotation occurred before, during, or after the strain. Often, we assume that rotation and distortion were occurring at the same time, but this is not necessarily the case.) Also note that if we choose our reference direction to coincide with a strain axis, then any finite strain appears non-rotational.

One special case of rotational deformation, **simple shear,** is of common interest in the study of shear zones and folds. This is a non-dilational strain in which one LNFE is not rotated. Simple shear can easily be simulated by drawing on the side of a deck of cards and then shearing the deck of cards. It is also easily simulated using the shear transformation in CorelDraw, Illustrator, or other drawing program.

In strain analysis we often are able to measure *s* or *γ* in a few directions, and would like to reconstruct the whole strain ellipse from these data.

[Examples shown in the lecture may include: belemnites; fold and boudinage of veins; veins and stylolites; deformed trilobites; deformed graptolites]

To do this, we need to know how *s* and *γ* vary with direction in the deformed rock. We measure the direction *φ'* clockwise from the *s _{1}* (X) axis, in the deformed rock.

For simplicity, let's deal with a pure strain, so that we don't have to worry about rotation.

The graph of *s* or *e* plotted against *φ'* looks like a distorted cosine curve.

We can also plot a graph of shear strain against *φ'*

This also looks a bit like a sine wave at low strains, but becomes more distorted at higher strain ratio *R*.

[These graphs are known as Breddin graphs: examples will be shown in the lecture]

In general the mathematical form of these curves makes them rather cumbersome to deal with.

It turns out that there's a simplification that can be achieved if we define two new quantities, based on the square of the stretch (also known as the quadratic elongation) *λ*.

The values of *λ* along the strain axes are *λ _{1}* and

*λ' *is defined as *1/λ* and *γ'* is defined as *γ/λ*

Then

$${\lambda}^{\prime}=\frac{{{\lambda}^{\prime}}_{3}+{{\lambda}^{\prime}}_{1}}{2}-\frac{{{\lambda}^{\prime}}_{3}-{{\lambda}^{\prime}}_{1}}{2}\mathrm{cos}(2\theta )$$

and

$${\gamma}^{\prime}=\frac{{{\lambda}^{\prime}}_{3}-{{\lambda}^{\prime}}_{1}}{2}\mathrm{sin}(2\theta )$$

These are equations of a circle with radius *(λ' _{3 }- λ'_{1})/2* with centre at

Notice how any point on the Mohr circle describes the state of strain on a line in the strained rock. It is easy to convert from *λ'* and *γ'* to the more useful quantities s and γ, and vice versa. Notice also that angles φ' measured in the rock are represented by points on the circle that are at an angle *2φ'* around the Mohr circle. The two strain axes, 90° apart in the rock, are located 180° apart on the circle. Notice that *γ'*, and therefore *γ*, is zero (as it should be) at both these axes.

In addition, the angle marked ψ on the diagram has to be tan^{-1}*(γ'/λ')* = tan^{-1}*γ*, or the angle of shear *ψ*.

Because it can be easily constructed with a compass, the Mohr circle construction is very useful for determining longitudinal strain and shear strain in any given direction, provided the principal strains are known.

So far we have met two methods for representation of finite strain in 2D. These are the Mohr circle and the strain ellipse. We now introduce a third representation, that is actually the basis for all the theory of the strain ellipse and the Mohr circle. This is a 2 x 2 matrix of 4 numbers known as the **deformation gradient tensor,** or more simply as the **deformation matrix** and commonly represented with a bold **F**.

- We can draw a unit square, and set up some axes so that it is at the origin of a
__graph__*.*The square will have two diagonally opposite corners at coordinates (1,0) and (0,1).

- After deformation the square will be a parallelogram. The coordinates of the two diagonally opposite corners are used to fill the
**deformation**, also known as the**matrix**__deformation gradient tensor__*.*

*Uses of the deformation matrix - as in other sections of the course, we will not make extensive use of the tensor representations of deformation because not everyone has done linear algebra. However, this link is provided for anyone interested in seeing how a little bit of linear algebra simplifies all these ellipses and circles!*

We now have enough information to deal with practical measurement of strain from rocks. It's important to remember that any strain determination is dependent on some assumptions about the starting state of the rock.

Because the starting rock can no longer be examined, these assumptions are always difficult to test in real rocks.

In strain analysis the best strategy is to use 2 or more different techniques, based on different assumptions, and see if they give the same answer.

This is the simplest case. If we have an object that was perfectly circular before deformation, we can measure its shape to obtain the shape of the strain ellipse.

The starting assumptions are

- The objects were initially spherical
- They deformed homogeneously with the matrix

The most likely objects are volcanic lapilli, and reduction spots. (Limestone ooids are sometimes very close to circular. However, they are less likely to deform homogenenously with their matrix, because they often have slightly different texture or composition than the surrounding material.)

Because we typically have no control over original size, this method can only tell us about distortion, not dilation.

Linear objects are sometimes available, in circumstances where we can reconstruct the length before deformation. For example, very rigid elongated objects such as belemnite fossils and toumaline crustals may undergo boudinage during strain. It is possible to determine the original length of the object by adding up the lengths of all the boudins. The final length can be measured directly, and the stretch can be calculated (*l/l _{0}*).

Once several objects have been measured, their orientations and lengths can be plotted (in radial coordinates). A best fit ellipse may be estimated using a set of elliptical templates. In theory a minimum of three points may constrain the ellipse but usually more are advisable.

The assumptions for this method are:

- There is no distortion of the boudins
- The separation of the boudins represents the whole of the longitudinal strain.

If these assumptions are satisfied, this method can in theory yield both the dilational and distortional components of the strain ellipse.

Often, we can measure shear strains more easily than longitudinal strains. This is particularly the case with bilaterally symmetrical fossils, which have been distorted by strain. A minimum of two, differently oriented fossils is necessary. If only two or three fossils are available, a Mohr circle construction is typically used. For multiple measurements we use a construction called a Wellman plot.

For the Mohr construction, plot the circular part of a Mohr diagram at arbitrary size on tracing paper. Draw two radii representing the lines along which shear strain is known; the angle between the radii must be double the angle in the rock (because angles are doubled in the Mohr construction). Next, plot the axes on graph paper, and draw lines from the origin inclined at angles ψ to the horizontal axis. Then superimpose the tracing paper on the graph paper to complete the construction, making sure that corresponding lines intersect.

For the **Wellman plot**, first draw on tracing paper a horizontal reference line. For each fossil, measure the orientations of two originally perpendicular lines. Draw lines with the same orientation, starting at __each end__ of the reference line (4 lines in all, forming a parallelogram). Mark the corners of the parallelogram. Repeat for each fossil. The marked points should fall on a strain ellipse that has the reference line as a radius.

Assumptions for this method:

- Original shape of fossils is known
- Fossils have deformed homogeneously with their matrix

This method yields only the shape (distortion component) and not the size (dilation component) of the strain ellipse.

A common problem in strain analysis may be that the only available markers were elliptical, not circular, even before deformation. We can attempt to overcome this with a plot known as an R_{f}/φ plot.

R_{f} is the **ellipticity** of a measured object (e.g. the ratio of long axis to short axis of a deformed pebble.). R_{f} can be thought of as a combination of an *initial ellipticity* R_{i} and a *strain ellipticity *R_{s}.

Phi is the orientation of the long axis of the object, with respect to some reference direction (e.g. north, or in a vertical section, horizontal.)

In the plot, each point is plotted on axes R_{f} and φ, typically producing a scattered cluster. The densest part of the cluster, somewhere in the centre, represents the φ value of the S1 axis. The value of R_{s} is represented by R_{f} at the densest part of the cluster. Sets of standardized curves are available for estimating R_{s}.

[Alternatively, the two can be plotted on a hyperbolic net, in which φ is measured as an angle and R_{f} is plotted as a distance from the centre. The net is rotated so that its line of symmetry bisects the cluster. The net also shows a set of hyperbolic lines; the line which bisects the cluster is chosen, and the value of R at its closest approach to the origin is an estimate of R_{s}.]

There are many examples of rocks in which the only strain markers are objects which are expected not to have deformed homogeneously with their matrix. Examples include pebbles in conglomerates with a soft matrix, sand grains or ooids that have become sutured by stylolytic contacts, etc.

In these cases, the shapes of the objects cannot be used to determine strain. Instead, it may be possible to use the *spacing* of the objects. This method works if the objects were uniformly spaced in all directions before deformation, or if they were at least, statistically uniform in their spacing; we say the objects were **anticlustere**. (What this means is that it's possible to identify an average or typical spacing of nearest objects, measured centre to centre, in the undeformed rock.)

The procedure involves tracing paper. Mark an x or a set of axes on tracing paper. Place the origin over one object. Mark the centres of all nearby objects. Move the tracing paper to the next object (without rotating it). Repeat the procedure.

After a sufficient number of objects has been traced (typically >50), an elliptical *void* may appear in the centre of the diagram. The shape of this void is the shape of the strain ellipse.

Assumptions for this method

- Objects were initially uniformly and isotropically distributed
- If objects were not spherical, they did not have a preferred original orientation
- Objects have not slid past each other during deformation

Unless the original size of objects can be estimated (rarely possible) this method gives only the shape (distortion component) of the strain ellipse.

Finite strain is generally the end result of a **strain history** in which innumerable **incremental strains** (or, in the limit as t-> 0) **infinitesimal or instantaneous strains **are combined.

Defining the complete strain history is a much larger challenge than measuring the finite strain. To keep thing simple we will continue to deal with only two dimensions - nothing flows in or out of our cross-section - 'plane strain'.

To describe progressive strain and flow, it's sometimes helpful to draw a field of small arrows representing the **velocity field**. We can also join these up to make longer arrows representing **particle paths**.

Particle paths are usually curved. Often there are one or two special particle paths that are straight. These have a special role in the generation of fabric in metamorphic rocks. They are known as **flow asymptotes** or **flow apophyses.** In linear algebra terms, they coincide with the eigenvectors of the deformation matrix.

At each moment in a strain history, the ongoing deformation is defined by **instantaneous strain axes **(or **ISA:** these are the** instantaneous stretching** and **instantaneous shortening** axes). Between them, there are usually **lines of no instantaneous extension** (LNIE).

All these lines may change orientation during the strain history. Except for special types of strain, the instantaneous directions of these lines do not coincide with their finite strain equivalents.

During deformation, material lines may move from the field of shortening to the field of extension or vice versa.

If rocks, and strain, were perfectly homogeneous, we would never be able to figure out strain histories because we would only see the end result, the finite strain. Fortunately, **strain partitioning** is common. In strain partitioning, different components of a rock record different parts of the strain history.

In **coaxial strain** the strain axes remain along the same material lines throughout the strain history: ie the strain axes do not rotate relative to the rock. As a result, the **finite strain axes coincide with the instantaneous strain axes, and with the flow apophyses.**

Note that although the strain axes don't rotate, other lines tend to rotate towards the stretching direction and away from the shortening direction.

Every increment of strain is **non-rotational** or **pure strain**. The **vorticity** of the deformation is zero.

The deformation matrix at every stage is a symmetric matrix. The flow apophyses, which are the same as the strain axes, coincide with the **eigenvectors** of the deformation matrix.

A diagram will show the effect of progressive pure shear on a rock with competent layers oriented initially perpendicular to s_{1}

As strain starts, competent layers buckle to produce folds. Then, with increasing strain, limbs of the folds rotate from the field of incremental shortening into the field of incremental extension. Initial folds may become unfolded and undergo boudinage.

Geometries like these are extremely common in folded rocks - limbs that are boudinaged while the hinges are tightly buckled.

In** noncoaxial flow**, the strain axes and the rock rotate relative to each other. The instantaneous strain axes do not lie along the same material lines througout deformation.

In noncoaxial flow the deformation matrix is **not symmetric**.

If we think of non-coaxial flow as being broken down into an infinite number of tiny increments, we can mathematically consider it as consisting of separate increments of distortion (pure strain) and increments of rigid-body rotation

In dealing with non-coaxial deformations, it would be helpful to have a measure of the amount of rotation compared with distortion. To do this we define a quantity called **vorticity**. For a non-dilational deformation, the **kinematic vorticity number** is the ratio of the rotation rate (in radians per second) to the distortion rate (in strains per second).

It's possible to distinguish different categories of noncoaxial flow, depending on different amounts of vorticity.

Progressive pure shear is characterized by a kinematic vorticity number of zero.

**Simple shear **is a special case of noncoaxial flow, but is important because an ideal shear zone develops by progressive simple shear.

In simple shear:

- All the particle paths are parallel to the shear zone boundary (SZB).
- There is only one flow apophysis, also parallel to the SZB.
- The kinematic vorticity number is 1
- The instantaneous stretching direction is at 45° to the SZB.
- The LNIE are parallel and perpendicular to the SZB

However, as *finite* strain increases:

- The finite stretching axis, and all the planar and linear fabric elements, tend to approach the SZB.
- Material lines rotate with the vorticity and may pass from the shortening field to the extension field.
- There is a relationship between the amount of shear and the orientation of finite s
_{1}. If the angle of s_{1}to the shear zone is θ and the amount of simple shear is γ, then γ = 2 / tan(2θ).

Subsimple shear combines distortion with gradual rotation of the strain axes, and is intermediate between pure shear and simple shear. There are two flow asymptotes or **apophyses**, which coincide with the eigenvectors of the deformation matrix, but they are not at right angles and do not coincide with the strain axes.

The kinematic vorticity number equal to the cosine of the angle between the flow apophyses, and is therefore between zero and 1.

If the rate of rotation exceeds the rate of extension, there are no flow apophyses. Everything rotates.

The vorticity number is greater than 1.

Supersimple shear describes the behaviour of relatively rigid particles in a shear zone, which tend to cause strain partitioning. The rigid particles capture the vorticity while the softer matrix captures the distortion.

Supersimple shear tends to produce features like **delta porphyroclasts** and **helicitic (snowball) porphyroblasts.**

....is now covered in a separate section