These notes are principally the contents of thecourse textbook

Lecture 1- Introduction to Course and Primary and Nontectonic structures

Lectures 2 and 3- Force and Stress

Lectures 4 and 5 - Strain and Rheology

Lecture 6  - Introduction to Brittle Deformation

Lecture 7 - Joints and Veins

Lecture 8 -  Faults and Faulting

Lecture 9- 

Structural Geology is one of the key courses taken by geologists.

Structural Geology is the study of deformed earthmaterialsat all scales. All deformation can be ultimately traced to platetectonics.Test yourself on this!
 

Structural Geology is  important because:

(1) You will begin to do somefieldwork. Geology is grounded in observations which are made in the field. For many this course will decide whether youstayin Geology or go into another field.

(2) Structural Geology carriesaunique strength different to other fields and at the same time difficult.Themost difficult part for many of you will be visualizing geologicalstructuresin 3-dimensions and changing in time (4th dimension).

How much deformation has taken placeisan aspect of Structural Geology known as Kinematics

(3) Structural Geology emphasizes 3Dviewing.We have to employ clever ways to reduce 3-dimensional problemsinto 2-dimensionalmaps and diagrams (e.g., stereonets, block diagrams,geological maps) Eventuallyyou should have developed a framework in yourminds that will allow youto see better in three dimensions. These aretools for answering how muchor what the kinematics are of a problem?

(4) Structural Geology explains why rocks deform over TIME? Is it because of their composition? Thesurroundingtemperature? Fluid content? Pressure? Rate of deformation? What are theforces involved?

 
(Click hereforchaptersummary)

Primary Structure: " those structuresthatform during or shortly after the deposition of rocks, and nontectonicstructures;i.e., those structures that are not an immediate consequenceof plate interactions".

" Strictly speaking, however, this usage is misleading, becausemost structures, if not directly, are indirectly a consequenceof tectonic activity. 

For example, the creation of slopes down whichsedimentsslide, the occurrence of volcanic activity leading to the flowof basalt.All of these phenomena are ultimately a manifestation of movements in the Earth. 

Under the broad masthead of "primary and nontectonicstructures"we discuss depositional, penecontemporaneous, intrusive, andgravity-slide structures for both sedimentary and igneous rocks. 

This heading also includes impact structures,whichare discussed at the end of the
chapter.

Interesting Primary Structures
I am going to highlight the following primary structures because 

(1) they are most relevant to the field exampleswewill see, e.g.
 

AngularUnconformitiesare the result of erosion are indicative of deformation.

Stylolites(animatedfilm) indicate the direction of shortening the sediment or rockunderwent

 
Younging direction indicators can serve to tell us the original "way up" in a rock. They are of many types of which the following are just a few:
 

Mudcracks

 

Cross-bedding

We will see crossbedding on the March 10 field trip in Pliocene-aged sediments of the Peoria Formation. Click on this link to go the the Web Page of the Field Trip. Here is an exampleof cross-bedding from Arizona 
 

GradedBeddingis one indicator of polarity inorder to determine the originalyoungingdirection (which way is up) in a bed 

 

Loadcastsor ball and pillow structures can also tell you which way is up

 

CoolingFractures- Fig. 2.23 

 

Pillowlavas-Fig. 2.22 and a detail of a pillow lava from an ophiolite comple in Oman slide

 

ImpactCratersaremuch less common on earththan on other terrestrial planets but can be examinedfor their unique structures which indicate short-lived high-pressures(shock metamorphism)

 

ShatterCones:Shatter cones are formed bythe passage of the explosion shock wave throughthe rock.

" They are best preserved and identified infine-grainedrocks. These Shatter conesoccur mostlyin shale:

           Chapter 3:  Figures 3.2, 3.3 and 3.12, pages 35-51

In geology, you will always hearofthe use of stress instead of force. Why? 

In the earth as rocks break, faults arecreatedby the concentration of force over surfaces. The deformation is more readilyexplained by stresses. 

For example: 
Great stress can be achieved with very small forces. Stresses from near the core/mantle boundary can be reproducedinthe laboratory. 

Even the greatest forces may not breakrocksif the stress is insufficient. 
 
 
 

 a heeled shoe can exert more stressthanan elephant. 

115 lb/.125 in. sq.vs. 10000lb/12x12 in sq.. 

football player: 122kg/4cm2 = (270 lb) / 1insq.

 Your car tires are in p.s.i (~36 p.s.i)eachp.s.i. is = 1 kg/cm2  or 1 p.s.i. = 0.07 kg/cm2 

 
 

Units of stress (stress) sigma = F/A i.e. kg m /s2 m2 or Pa,units of pressure  100000 Pa  ~ 1atm of pressure = 1bar 10 kbar = 1000 MPa = 1 GPa For a  density of 2.1 g/cc lithostaticpressureis 21 MPa at a depth of 1km For a density of 2.7 g/cc lithostaticpressureis 27  MPa at a depth of 1km At the core-mantle boundary the pressureis136 GPa

 

Lithostatic stress is the stress inducedbythe overlying weight of rock. 

Let's calculate the lithostatic stressat a depth of 10 m

 WHAT'S  the lithostatic stressata depth of 1km in granite? 

Hydrostatic stress conditions are achievedwhenthe principal stress values are the same in all directions

We can represent stress at a point by drawinglineswhose length is proportional to the value of stress at the centerof the stress ellipse 

At depth we have stress from all directions.However,we can show that the three stresses that are at right angles toeach othercan always be found with which to describe the general stresscondition: 

The strength of the crust at any depthcanbe regarded as the maximum differential stress it can withstand beforefailingeither in plastic mode or in brittle mode.  A plot of thestrengthof the crust with depth gives the  yield-stressenvelope.and Jelly sandwich analogy

The way a rock deforms internallyis known as the strain.  In addition to internal deformationthe rock may rotate in space, say within a fold and the rock mayshoft when it is translated in space aboard lithospheric plates. Actually on a spherical shell as are the lithospheric plates, translationson a regional scale are not straight-line paths but curved paths on a sphere. On a local scale those curved paths may appear as straight (translation)lines.  Click here to see a graphical representation of the paths certain points on a deforming

So, strain is the distortiondetermined INSIDE a body. 

Linear strain
In order to describe strain we can discuss stain measurements:
strain is described as theratio of a deformed length compared to the original length of the materialline.
We can measure the extension(e) of something, say by 33%.  THis means that the change in length(delta l) is 33% the size of the original length.  When the sign ispositive we take it to mean that the material line has lengthened.

Angular strain
Angular strain is measurein degrees as is indiacted by alpha in example b of figure 4.9 in yourbook.
 
 

Rheology depends on many environmentalvariables, which can be studied in a laboratory setting using a rock press.  These variables are: 
 

• therate (de/dt) at which thedeformation is carried out 
• Rock type
• Temperature,which also varies as a function of the rock type
• Confining pressure
• Time
• Confining Pressure
• pore-fluid pressure

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Introduction to Brittle Deformation

Fractures, joints and faults Are they the same?
In this introductory class to structural geology we will look at faults fractures.  A fracture is a break in a rock.  If there is no perceptible displaement ALONG the fracture we normally call this type of fracture a JOINT.  If there is a VERY SMALL amount of movement ALONG the joint we use the term shear fracture, although in a general sense when the movement is measurable, then that is the case of a fault.  Dykes and veins are cases where the joint opening has been filled in by some precipitate or magma.

Brittle deformation such as joints and faults occur essentially within the realm of elastic behavior or before too much plastic deformation has taken place, i.e. in the colder shallower portions of the upper crust (< 15 km in continental crust).

Why do need to look at joints and faults?
It is important to distinguish between joints and faults (i.e., when there is significant shear movement along the fault) because these can tell us the conditions and orientations of stress with respect to the fracture.  The original direction and amount of stress can for example tell us the direction in which plates were converging, slipping or separating.
 

• Brittle deformation is the mechanism that causes earthquakes. 
• Fracture patterns control the flow of contaminants in the subsurface.  Expansion and contraction joints developed during shrinking and swelling of clays in our part of the country create hard-to-detect arrays of vertical fractures which speed the velocity at which contaminant liquids can travel through the near surface sediments. 
• For example, the Baton Rouge fault divides fresh waters from saltier waters to the south.  The drinking supply to a million people is in the hands of a fracture system. 

 

 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 

In all the cases we can see in your book  (figure 6.4 ) with the exception of the tensile crack or joint, the maximum principal compressive stress direction is at about 30 degrees to the plane.  (Note that this is different to the 45 degree value for which there is a maximum shear stress and which we estimated when we were considering the difference between forces and stress.)
 

Griffith Crack Theory versus othe empirical rules regarding conditions needed to fault rocks
There are two ways we can predict the theory of fracture development: one method involves quantitatively predicting the conditions of failure with a mathematical relation (Griffith Crack Theory) and another through a mathematical law that describes the results of experiments (Mohr-Coulomb failure envelop and Byerlee's Law)  over the last two hundred years.

Griffith Crack Theory
This theory assumes that all rocks have imperfections which consist of tiny microscopic openings oriented in all directions.  When stress is applied the stress redistributes itself around those small cracks and concentrates itself on the tips of the cracks causing the cracks to spread along their major axes in a preferred direction with respect to the principal stress directions.
 

Joints
Tensile joints can be created by in spite of confining pressures by introducing a stress that can overcome the minimum principal stress.  High pore fluid pressures can break a rock in this manner.

Under very low confining stresses a shortening can open tension joints.
 

Faults
Under greater confining pressures differential stresses force microcracks to grow and concentrate in large numbers along planes at a low angle to the principal stress direction. (See figure 6.15) These planes ebcome the zones of slippage or faulting.
 

Experimental Laws-
Over time and through experience it has been observed that there is a practical linear relation between the normal and shear stresses a rock can sustain across a potential fault surface.

The part of the Mohr-Coulomb yield or failure envelope Coulomb developed has a linear realtionship.   You can think of the constant of the C as the cohesion of the rock or the internal bonds that must be broken to create the shear fracture surface and then you can think of mu as the constrant of proportionality that represents the degree of friction or roughness across the surface once it is created. (yield shear stress = C + mu x normal stress across the shear fracture)

Having said that, the relation is not exactly linear because the yield stress envelope curves to higher failure angles at greater confining pressures (greater depths) so that a fault should be somewhat listric in nature.

Byerlee's Law   is a remarkable result because he showed that once a shear fracture is created the degree of friction across the fracture surface is virtually the same for all rock types and only changes as we increase the confining pressure.  In practice, as Byerlee's Law is taken to mean, we assume that at least in old continental crust there always exist pre-existing potential shear fracture surfaces which can slip once the coefficient of friction across the surface is overcome. (yield shear stress = 50 MPa + 0.6 normal stress across the shear farcture)
 
 

Joints and Veins Why study joints?   Tectonicists may want to know what the past stress field look liked (paleostress) Geomorphologists want to know whether the drainage patterns are controlled by joints? Engineers may want to know what to know the natural directions of weakness in the subsurface Environmental geologists may need to understand the potential pathways for contaminants What do joints look like? On the surface an ideal tension joint has "feather-like" features which are referred to as plumose structures.  Hackles or the ridges that form on a joint surface diverge in the direction in which the joint originally opened.  If the joint openend up in several stages "arrest lines" perpendicular to the hackles are apparent. In the field, the appearance is not always very clear because the hackles only raise the surface of the joint a few millimeters at most Joints are never alone (i.e. they occur in groups or sets) often they have a regular spacing (systematic sets) Joint spacing and spatial frequency IN SEDIMENTARY varies as a function of bed THICKNESS and composition (among other factors) Normally thinner beds have joints spaced more closely than thicker beds because once a large joint is formed it releases a stress around a larger surrounding area than a small stress.  New joints in the area that has expienced the stress release are less apt to form. Stiffer material builds up more stress and tends to have a greater density of joints.   How are joints formed/ what can they tell us about structural history? Joints are formed at right angles to the minimum principal stress direction.  Changes in the orientation of the stress field LOCALLY may cause a joint to  change direction. Joints can form by the stress concentraion of fluid pressure in microcracks (hydrofracturing) Joints may form related to extensional stresses inside folding beds of rock Joints may be related to the stress-release of rocks as they near the surface by erosional unroofing (sheeting joints).  They expand in the vertical direction as the lithostatic stress drops.   Simultaneously they shrink in the horizontal direction.  Joints also cool and contract vertically as they cool on their way to the surface. "Gashes" or joints that were once filled with chemical precipitates such as calcite or quartz record by the overall shape and crystal growth patterns whether there has been non-coaxial strain during their formation (sigmoidal tension gashes) or coaxial strain. The direction of growth of joi What are different examples of joints? For examples of: liesengang structure tension gash sigmoidal tension gash systematic joints  columnar joints click here Try to estimate the direction of the principal stresses from the example in Oman A well-known example of columnar joints is found in northern Ireland at the Giant's Causeway

Faults and Faulting Introductory Pop Question:   Image you are a well geologist and you are told that the same rock layer has been pierced twice by drilling in the vertical?  What kind of scenario would you use to explain the phenomenon to your impatient boss?   What do faults look like?  .... Examine this profile  through a generic fault. As we descend through a fault zone from top to bottom you will observe a change in the deformed rock types as they display brittle (breccia and cataclasite) through ductile behavior (mylonite).  During plastic deformation the rock shears without losing continuity because the surrounding conditions permit the rock to heal itself microstructurally during deformation. At a shallow depth faults will tend to be composed of several related faults which together form a fault zone composed of a main fault and fault splays At depth, discrete offsets become less clear in the zone of ductile deformation. Faults can occur on differents scales, from the microscopic scale to a regional scale, e.g. San Andreas strike-slip fault system.     Relation between stress and faulting - Andersonian Theory We have seen that in general the maximum principal stress direction is at a shallow angle to the plane of shearing.  Rocks fail more readily under normal fault conditions Fault types Faults are named according to the orientation of the dominant slip with respect to the fault attitude. How many different types of faults are there?  Look here. For each of the normal and thrust type faults we can distinguish faults based on their steepness.  Normal faults dip steeply, (60 degrees) but reactivation can move faults at shallower or even steeper angles. Thrust faults are low-angle faults (10-30 degress) but are called reverse faults when they have greater dips (45 or more), again probably because of reactivation of an old fault of opportunity. Strike-slip faults can be dextral (right-lateral) or sinistral (left-lateral) whether  the opposite block moves to the right or left respectively. Each of the different plate margins will tend to have a dominant type of fault system (groups of faults).  Contractional margins (collisional margins) will tend to develop thrust faults in special arrangements.  Normal fault systems will tend to form at extensional divergent plate margins.     How to recognize a fault in the field... Fault types based on the dominant slip on their surface CAN NOT BE distinguished unequivocably by ONLY a 2-D view;  We usually need a full 3-D representation of the fault.  For example we can appear to have a stgrike-slip fault but have it really be a normal fault. In field mapping of thrusted terrains we can find features such as tectonic windows and kilppen. Fault motion can be determined in part by such surface features as the following: slickensides,  fault striae,  chatter marks/ releasing step Faults and folds It seems logical that EXACTLY the same rock deformation conditions CAN NOT EXIST at the same time and in the same place. However, as a function of different rates of strain a fold can progressively develop into a fault, so that what initially starts off as a fold eventually gets faulted, e.g., fault-propagation fold.   There are also other settings in which folds can form, creating very interesting features such as sheath folds, fault-bend folds.  Click here for a complete range of examples.