Structural geology Flashcards
Joints
Most common type of tensile fracture.
Form near Earth’s surface.
Control the bulk strength of the rock (coal cleat)
Important fluid conduits (groundwater and hydrocarbons)
Sigma 1 parallel to joint.
Often control weathering and subsurface fluid flow.
Systematic joint sets due to both regional and local stresses during deformation.
Joint spacing controlled by bed thickness.
Orthogonal joint sets require flip of principal stresses.
Rapid unroofing causes joints
Cooling and contraction causes joints
Fracture modes
Mode I - opening and tension [fractures]
Mode II - sliding and shear, strike slip
Mode III - tearing and shear
[Veins and Faults all three - veins must have mode I]
Vein systems
Array (collections) of fractures filled by mineralisation. Leads to bulk volume increase. Common in low grade metamorphic rocks (PT conditions). Contain many important mineral deposits (quartz reef - thick vein with gold).
Brittle failure, sigma 1 parallel to vein (like fractures)
Shape depends on fractures - can be folded (with at least 2 stages of deformation [1 - extension -> fracture, 2 - compression -> folding]
Veins - opening mode fractures filled with new minerals. Crystal growth often controlled by the progressive opening of the vein.
Opening perpendicular to crack
Mineralisation perpendicular to extension direction.
Oblique opening
Mineralisation at an angle to the opening (due to shearing from faults).
Episodes of cracking and opening
Syntaxial veins - grow inwards from wallrock
Antitaxial veins - appear to grow from a median suture line towards walls (outward from walls)
Ataxial veins - formed by repeating opening and sealing of fractures
En Echelon Veins in Shear Zone
Veins that step down - can determine how the rock was sheared. Shearing in opposite direction to what you would assume, in order to keep the steps. In class exercise.
Faults
Fractures with shear displacements.
Brittle failure, sigma 1 must have some obliquity to fault plane (sigma 1 bisects the acute angle)
Dip - angle from horizontal to fault plane
Hade - angle from fault plane to 90
Dextral = right lateral
Normal fault - horizontal component called heave and vertical component called throw
Reverse fault - horizontal component called contraction and vertical component called throw
Thrust - low angled fault <45
Block diagrams - beware
Often misleading, as they give a false impression of the 3D form of a fault. Faults do not extend laterally forever - they are discontinuous with a finite length. Somewhere along this length they have a max displacement and at ends have 0 displacement.
True offset
Oblique-slip fault - has dip-slip (DS) and strike-slip (SS) components
As rocks are often horizontally layered, DS usually easier to measure than SS.
To measure SS, need either steeply inclined beds or dyke.
To accurately calculate the total displacement vector of a fault, need different dipping beds/layers/intrusions to be faulted. Knowing the slip direction can be helpful.
In map view, offsets can look strike-slip, but in reality, are normal.
Faults - movement indicators from outcrop
1) slickensides (fibrous mineral growth) and striations (scarpes) on fault planes parallel to the direction of slip.
2) fault drag folds (normal drag)
3) fault plane striations
Fault rocks
Faults often grow by the superimposition of many slip events on the fault surface. Each slip causes an earthquake. Depending on PT, different types of rocks form.
Shallow depths - forms by fragmentation; fault breccia (>30% rock fragments visible cohesive/incohesive) and fault gouge (<30% rock fragments visible and rock incohesive)
Deep - rocks become cohesive (lithified with some recrystallization). Matrix formed by tectonic reduction of grain size. Protocataclasite - cataclasite - ultracataclasite (grain size decreasing)
-cataclasties still considered to be formed under brittle conditions.
Deeper - ductile (recrystallization dominates) -> mylonites (strong foliation)
-very rapid movement -> partial melting of rock -> pseudotachylite (black glass)
Folds
Folds form in already-layered rock masses (bedded sediments, layered/foliated metamorphic rocks or in igneous rocks in discrete layers). They may occur on any scale (scale usually depends on thickness of layers they deform). Thick layers produce larger folds.
Fold geometries
Straight hinge line - cylindrical fold (doesn’t match reality)
Bowed up/down hinge line - non-cylindrical (reality)
Synform - downward closing fold (concave up)
Antiform - upward closing fold (convex up)
Cline - need to know ages
-anticline - oldest in middle
-syncline - youngest in middle
Monocline - one limb
Fold hinge - joins points of max curvature
Axial surface - contains fold axes of folded layers. Defines plane of flattening (xy plane of strain ellipsoid)
Crest line - the line which lies along the highest points in a folded layer
Trough line - the line which lies along the lowest points in a folded layer
Plunging folds - axis plane dipping
Fold axial trace - intersection of axial plane with present land surface
Fold classification
Links to Fleuty fold classification (axial plane dip vs plunge dip)
Upright fold - axial plane is vertical
Inclined fold
Overfold - one limb is overturned
Recumbent fold - axial plane is horizontal
Symmetric folds - limbs have same dip, same shape
Asymmetric folds - verge to one direction, one limbs is steep and one is shallow
Overturned folds - verge to one direction, shallow limb and overturned limb
Vergence direction suggests shear direction (right angles to limb)
Fold tightness
Interlimb angle - draw tangents to folded surface at points of inflection Flat lying - 180 (least strain) Gentle - 120-180 Open - 70-120 Tight - 30-70 Isoclinal - 0-30 (most strain)
Fold hinge shapes
Folds can have varying hinge geometries due to the nature of the faulted sequence or the conditions of folding. Ex: chevron and kinds often formed in thin multi-layered sequences and more typical of relatively shallow depths of formations
- Chevron folds - straight limbs and angled hinges
- Kink folds - horizontal, angled, horizontal
- Cuspate folds - go up to a point (soft rocks)
- Box folds - more rounded than kink
- Disharmonic folds - beds have different fold shapes, common
Facing direction
If beds are upright (beds young up the axial plane) we have upward facing folds (Anticline: concave up and oldest in middle; Syncline: concave down and younger in middle).
If beds are overturned, then have downward facing folds (beds young down the axial plane). Antiformal syncline (concave up, youngest in middle) and Synformal anticline (concave down, oldest in middle)
-Syncline = younging direction. Antiformal = shape.
Fold description method
Axial surface dip Hinge dip Fleuty description Interlimb angle and fold tightness Symmetry Harmonicity Hinge shape Fold class Depth of formation
Non-cylindrical folds/pericline
Reality of folds - they die out along their length (periclines). Amplitude changes across the fold.
Often developed as en-echelon structures, as one dies out laterally, another grows laterally. The zone between two en-echelon periclines, where strain transfers from one fold onto another, is a relay
Sheath folds
Periclinal folds have curvilinear fold axes. If a fold axis curves by >90 it is called a sheath fold. Found in very highly deformed zones (geometry tells us about their deformation).
Classes of folding - quantitative
Based on fold profiles (cross-sections) of a folded layer. Classes of folds distinguished by:
-relative curvature of upper and lower bounding surfaces of layer
-relative thickness of the folded layer in the hinge vs limbs
Class 1 - isogons converge (fan outwards); shallow/brittle
-a = strongly convergent (bottom bed tighter than top)
-b = parallel
-c = weakly convergent
Class 2 - isogons parallel; deep/ductile
-similar
Class 3 - isogons diverge (fan inwards); deep/ductile
-divergent (top tighter than bottom)
Dip isogons - a line that connects a point on top and bottom bed with the same dip
Parallel folds (class 1b)
No thickness change in hinge or limb, uniform orthogonal layer thickness. Dip isogons perpendicular to fold limbs.
Typical of strong layers within a weaker matrix. Spacially inefficient and cannot extend for great distances. Shallow in crust.
Similar folds (class 2)
Thickened hinges and thinned limbs, uniform thickness parallel to axial plane. Dip isogons parallel to fold trace.
Typical of weaker rocks/greater depths. Spacially very efficient (exactly equal curvature) so the rock can extend for great distances.
Harmony
The extend to which folding of multilayer sequence is consistent through the sequence.
- disharmonic fold - dies out within a couple of half wavelengths
- harmonic fold - continuous along its axial trace for many multiples of its half wavelength (each bed doing the same thing)
- polyharmonic fold - harmony on various multiples of half wavelengths within the multilayer sequence.
Parasitic folds
Polyharmonic folding generates parasitic folds. Occurs on many scales. Can describe the collective geometry of parasitic folds with respect to their enveloping surface.
Parasitic fold geometries are predictable in relation to associated main fold (map vergence)
Fold vergence
Characteristic of parasitic folds
Indicates the direction of movement and rotation during deformation.
Can identify the location of major folds using minor folds.
Z (LHS of fold)
M (top of fold)
S (RHS of fold)
Powerful field mapping tool, especially in areas that are relatively poorly exposed.
Fold-cleavage relationships
Axial planar cleavage - cleavage that is parallel to the axial plane
Planes defined by fractures or mineral alignment parallel to a folds axial plane
-bedding steeper than cleavage = overturned limb
-cleavage steeper than bedding = upright limb
Assume:
-bedding folded and right way up
-cleavage parallel to axial plane
-axial plane bisects interlimb angle
Foliation
Alignment of minerals, forming cleavage or schistosity. Usually the plane of flattening xy plane of strain ellipsoid.
Folding mechanisms
Brittle deformation - buckling, bending, flexural slip and kinking
Ductile deformation - passive folding, flexural shear/flow, oblique shear
Buckling
Horizontal forces
Sinusoidal (smooth wave) folding of a single strong layer within a weaker matrix by lateral compression (layer maintains thickness). Tends to form parallel folds with uniform orthogonal layer thickness. Shallow crust.
Strain accommodated parallel to layering (tangential longitudinal strain). Most strain accommodated in hinge (nothing happening in limbs).
Outer arc extended; normal faults, extension veins, boudinage, horizontal cleavage
Inner arc compressed and shortened; thrust faults, stylolites, vertical cleavage.
Between them is a neutral surface where no deformation is taking place.
Bending
Often hard to separate from buckling
Vertical forces
Achieved when F act across the layer (at 90) and may involve more than one mechanism.
Look at context - diapir (salt or intrusion rises into layers and pushes them up), faults (beds respond to vertical motion - reverse and normal), boudinage *layers drop into folds)
Flexural slip
Where many thick, strong and still layer in between which is some material with low cohesion (weak boundary). Length of beds does not change, slip between (eg: bending a book).
All action in limbs and nothing in hinge.
Sense of shear changes across hinge zones - consistent between anticlinal and synclinal limbs.
Produces parallel folds - constant bed thickness.
Slickensides and striations are lineations and oriented at right angles to fold axis.
Slip at layer boundaries form bedding-plane faults (thrust) characterised by grooving and mineral growth parallel to movement direction (slickensides or striations).
Within the competent layer, veins often open at 45 angle to the shear directions of individual beds. Sigmoidal shape of cracks due to progressive lengthening of veins with rotation.
Kinking
Produces folds with straight limbs and very sharp hinges - chevron folds or kink bands.
Occurs in strongly layered or laminated sequences where have strong mechanical anisotropy between very competent layers and thin incompetent layers (associated with thin bedded sandstones interbedded with shales; slates and schists). Resultant folds show flexural slip and are parallel folds. MIX of buckle and flexural.
Straight limbs, angular hinge
Usually parallel folds - shallow crust.
Chevron folds
Kinking
See very intense layer-parallel slip on incompetent layers (discontinuities between layers)
gaps in hinge zones = saddle reefs (can fill with mineralisation)
lock up when inter-limb angle ~60
Kink bands
Kinking Smaller scale (mm) than chevron Two parallel sided, very well defined axial surface. Layering deflects suddenly across the 1st kink plane then returns to its original orientation across the 2nd kink plane Ideally, each kink has an inter-limb angle of 120 and is symmetric around axial plane
Passive folding
Produces harmonic folds where the layering plays no mechanical role and has no influence on the fold shape.
Can form in response to any kind of ductile strain (simple shear, subsimple shear, transpression, pure shear).
Oblique shear/flow
Distortion of layering by shear on an infinite number of closely spaced parallel planes that are inclined perpendicular/at a high oblique angle to layering. Displacement on shear planes varies along the layer so it folds. Deformation = heterogeneous simple shear -> class 2 (similar) folds. Rocks must be very ductile and have low viscosity contrast - deep in crust or in particularly weaker rocks (salt).
Flexural flow/shear
Modification of folding layer by simple shear that occurs parallel to the layer and is evenly distributed within the layer.
Occurs in layers with an initial uniform fabric (shale) and is geometrically similar to flexural slip folding thousands of infinitesimally thin layers.
Distortion in limbs, no strain in hinge.
Strain becomes more evenly distributed in limbs, shear increases down limbs (change in orientation), pure flexural folds have no neutral surface and strain increases away from hinge. Do not see individual slip surface.
Controls on fold wavelength - buckle folds
Two main controls - viscosity and thickness (both proportional to wavelength but t more significant due to 2t vs cube rooted n)
Viscosity = resistance to flow, high viscosity (most competent rocks).
Increased viscosity ratio = longer wavelength (proportional)
-high viscosity ratio = high viscosity layer encased in low viscosity rock
-low viscosity ratio - folds are poorly developed (short wavelength), layer thickens by pure shear
Thicker layers = longer wavelength
Controls on fold wavelength - parasitic folds
Interplay between layer thickness and viscosity contrast on intensity of folding (wavelength, A, number of folds) explains how parasitic folds might form.
Thinner beds with lower viscosity contrasts -> short wavelength; incorporated within larger wavelength folds defined by thicker layers
Controls on fold - Harmony
To form parasitic folds, polyharmonic folding must occur.
Controls on fold harmony relate to layer competency and thickness, but also to the relative abundance of and distance between competent layers.
-closely spaced with same properties -> harmonic folds
-competent layers are at intermediate distances get polyharmonic folds
-few similarary competent layers which are far apart -> disharmonic folding
Superimposed folding (rocks folded twice)
Multiple deformations can produce their own fold sets (F1, F2...). Superimposition can produce complex geometries (complex to distinguish in 2D). Ramsay and Huber - orientation of fold axis and axial surface of first generation folds Type 0 (redundant superimposition) - exact same orientation, add together to get greater amplitude Type 1 (dome-basin or egg-carton) - folds perpendicular to each other, forming dome-basin interference pattern. -dome = anticline + anticline -basin = syncline + syncline Type 2 (dome-crescent or mushroom) - axial surface has low dip fold add to upright fold Type 3 (convergent-divergent or rippled fold) - isoclinal recumbent fold + upright fold (produces synformal anticline etc)
Crustal loading and unloading
Gravity and lithostatic stress
-unloading - removal of overburden and corresponding lateral stresses, allowing expansion of rock by fracturing.
Plain strain
Deformation in which one of the principal strain axes does not change length.
Impact
By meteorites, catastrophic event responsible for large local stresses (craters).
Thermal effects (+/- phase change)
Changes in T can change stress conditions by
1) increasing P of fluids in rock
2) causing expansion (heating) or contraction (cooling)
3) causing mineralogical phase changes
Plate tectonics
The most important forces arise from plate motions - drives geological change. Structures usually at plate boundaries (high stress).
Stress (σ)
Related to P. Force per unit area (planar surface within rock or a surface partly containing volume of rock).
Stress is a pair of equal and opposite pressures acting on a surface = a pair of equal and opposite forces acting on a unit area of a body.
Stress is not seen
1Pa=1N/m^2
1bar = 0.1 MPa
Stress gradient with depth is 250bar/km or 25MPa/km
Strain (=deformation = change in shape)
Measure of the response of a rock to stress. Relates to change of shape and/or volume.
Elastic strain
The removal of the deforming stress causes an immediate return of the body to its original shape.
Strain is temporary and recoverable.
Straight line
Hooke’s law: σ=Ee
-linear relationship between stress (σ) and strain (e).
Not seen in rocks as rocks only record the permanent strain history since they formed. Eventually get an earthquake.
Shallow crust
Temporary distortions of the lattice of individual crystals. Atomic spacing changed proportional to stress (relation to inter-actomic bonding force which is characteristic of each mineral -> different E).
Viscous strain
The rock does not fracture and there is no return to original body shape after removal of deforming stress
Strain is permanent and unrecoverable.
Curved line
Viscosity = material’s resistance to deform
Viscous strain is seen in rocks as it is permanent.
Viscous deformation = material flow without failure
Ductile deformation = permanent strain without visible fracturing
Occurs at depths (high PT) or in weak rocks with lots of clay close to the surface
Accommodated by 3 principal deformation mechanisms (microscopic scale)
1) cataclasis - fracture and sliding of solid particles, individual crystals fractured
2) Intracrystalline Plasticity - dislocation creep and dislocation glide, crystals change shape
3) Diffusive mass transfer - molecular scale diffusion of material in solid state or in solution.
Brittle strain
The rock fractures and there is no return to original body shape after removal of deforming stress.
Strain permanent and unrecoverable.
Associated with joints, veins and faults.
Elastic limit
Beyond this point, the sample will change shape permanently.
Young’s modulus (E)
Elasticity of the material
E=σ/e
Different rocks have different elasticities which can change at different conditions.
Newtonian fluid
Ideal viscous fluid.
σ=ne.
n=viscosity (steeper line = more viscous)
e.=strain rate
Relationship between stress and strain rate: with constant stress, strain increases linearly with time: e=σt/n
-higher the applied stress, the faster the material deforms
Total strain depends on the magnitude of the stress and the time over which stress acts.
A shearing stress must be constantly applied to maintain movement
More strain = lower viscosity
Ideal types of strain
Perfectly elastic materials show linear stress-strain curves until they reach their failure strength and fracture - brittle failure
Perfectly viscous materials never fails, continue to flow as long as a constant stress is applied; flat-line curve - ductile failure
Viscoelastic behaviour
A material that behaves elastically overall, but takes time to reach limiting value.
Material instantly behaves elastically with applied stress, but the straining continues after this, decreasing exponentially with time until reaching elastic limit.
Stress below failure strength -> instantaneous elastic response and remaining strain decays to 0.
Real rock behaviour
Elastic + visco-elastic + viscous. Elastic immediately recovered, viscoelastic recovers over time. Viscous is permanent.
Factors influencing type of deformation
P T Rock properties Strain rate (e.) Fluid pressure
Depth and deformation
Effects of depth and T on deformation give rise to the brittle-ductile transition (BDT), below which all rocks deform in a ductile fashion.
BDT moves up for higher T, faster rate of strain and more fluids
Typical depth ~10-12km
Brittle upper zone - faulting and earthquakes
BDT
Ductile lower zone - shear zone and fault creep
Hydrostatic stress (loading)
Due to the density of overlying water at rest
Lithostatic stress
Due to the density of overlying rock
Pressure and rock deformation
With increasing confining P a rock gets stronger (yield stress/strength increases)
Behaviour has viscoelastic properties and becomes more ductile (inhibiting failure)
T and rock deformation
Increase in T decreases yield strength (stress) and Young’s modulus.
Increased T introduces viscoelastic behaviour and becomes more viscous.
BDT ~ 300 degrees
Rock properties and deformation
Such as strength of rock (cohesion) - related to mineralogy, texture, degree of fracturing…
Mudstones/salts weak and more prone to ductile flow
Granites strong and more prone to fracturing
Strain rate and rock deformation
Lower strain rates, decreasing yield strength and increasing ductility.
Long-term strain behaviour is creep. On long time scales, viscous strain can arise under low stresses that would produce only elastic effects on short time scales.
Pore fluid pressure and rock deformation
Pore fluids reduce the resistance to slip along internal planes/grains.
Usually water-based, the elastic properties of rocks are reduced (water is incompressible).
Fluids weaken a rock and decrease Young’s modulus, causing ductile deformation.
Brittle deformation - structures
Joints - fractures in rocks with no shear displacement
Veins - fracture which is filled with precipitated material
Faults - rock deformed brittlely with shear displacement
Ductile deformation effects
1) deformation of objects (3D). Sphere -> ellipsoid
2) Deformation of layers -> folding
3) Foliation: alignment of minerals by rotation, recrystallization of both.
Physical changes in deformation
Rigid body translation - object shape constant, movement in 1D
Rigid body rotation - object shape constant, body rotated about an axis
Distortion - shape change by squeezing or squashing
Volume change - deforming object to change V
Homogeneous vs heterogenous strain
Strain usually heterogeneous, but on some scales (
Strain ellipsoid
Measures the amount of ductile deformation in 3D. Matches the changed shape.
X-axis - long axis; Z-axis - short axis.
-flattening strain - y-axis extended
-constrictional strain - y-axis shortened
-plane strain - Y-axis unchanged, all deformation in x-z plane.
Pure shear (co-axial strain)
Irrotational strain. Major axes of strain ellipse unchanged through deformation (ie. axes remain where they are and get longer/shorter) - coaxial strain.
2D coaxial strain measured by aspect ratio of strain ellipse (X/Z).
All deformation in this dimension
No change in volume
Simple shear (non co-axial strain)
Rotational strain. The major axes of strain ellipse rotate through deformation.
2D non-coaxial strain measured by aspect ratio of strain ellipse (X vs Z) and by angular shear and shear strain (Y=tan(a))
No change in volume
Measuring strain
1) measure the length changes of lines
2) measure changes of the angle between lines
Shear strain (Y) = tan (angle) = change X/I
Extensional/contractional strain
e=I-I0/I0
-shortening with be negative strain
Strain (%) = e x 100
Stretch = I/I0
Flinn diagram
Quantitative classification of strain ellipsoid shapes. y-axis - ratio of x/y axes x-axis - ratio of y/z axes k=(a01)/(b-1) = tan (q) 3 end members types of 3D homogeneous strain. -uniaxial prolate, k=infinity -plane strain, k=1 -uniaxial oblate, k=0
Measuring strain
Using change in angles and length
requires you know the original shapes and/or dimensions of the deformed objects = strain markers
i) originally circular - oolite, reduction spots, lapilli tuffs
ii) originally elliptical - conglomerate pebbles
iii) original known size - belemnites, acicular minerals, boudinage. e=I-I0/I
iv) objects originally possessed bilateral symmetry - fossils (tribolites)
Initially spherical objects
strain ratio: x/y=tan(a)
Measure long and short axes dimension and plot. Angle (a) is angle from line of best fit to short axis (x-axis)
Spatial distribution - originally elliptical
Centre-to-centre method
Higher strain occurs in matrix and strain reflected by changes in distance between randomly arranged particles (original shape required).
Measure distance between 2 centres and the angle between that line and a reference line.
Undeformed rock - horizontal line
More deformation - more extreme sine curve
Initial bilateral symmetry
If the principal direction unknown, use Wellman graphical method.
i) transfer hinges and symmetry lines
ii) draw arbitrary line A-B
iii) draw lines parallel to hinge and symmetry line from points A and B
iv) draw best fit strain ellipse through all corners
Stresses
Stress ellipsoid: σ1>σ2>σ3, all perpendicular to one another. Perpendicular to strain ellipse.
Origins of stress - crustal loading, crustal unloading, thermal effect and plate tectonics
Geostatic stresses - stresses caused by water or rock
pgz=σv=σ1=σ2=σ3
Hydrostatic stress (pressure) or confining stress
σ1=σ2=σ3 - spherical stress ellipse
Newly solidified magma chambers; young, undeformed sediments
Uniaxial compressional stress
2 axes = 0.
σ1>0, σ2=σ3=0.
Condition of mine pillars.
Biaxial stress
One axis = 0
σ1>0 and σ3>0, σ2=0.
Implicit in many 2D cross-sections
2D stress ellipse
Triaxial stress
Most common
All stresses >0.
Represented by strain ellipsoid
Faults
Usually results from extension or compression and if have moderate confining pressure (σ3 +ve (compressional))
Conjugate faults, where sigma 1 bisects the acute angle
Andesonian faulting
Normal faults - σ1 vertical and σ3 horizontal
Reverse faults - σ1 horizontal and σ3 vertical
-folds
Strike-slip faults - σ1 horizontal and σ3 horizontal
Joints
Usually result from extension and if have low confining pressure (σ3 low or negative (tensile)).
σ1 parallel to joints
Mohr Diagram
Represents the stress state at the point of failure (relationship between shear stress, normal stress, principal stresses and angle of failure).
Made from experimental data.
Plot s1 and 3
Midpoint = mean stress (s1+s3/2)
Radius = Deviatoric stress (s1-s3/2)
Differential stress = s1-s3 (drives failures - larger circle = more likely failure will occur)
P - point of failure
B is the angle of internal friction = 90-2a
C is the cohesion/strength = y-int (minimum shear stress value required to cause failure when normal stress is 0)
Failure envelope - by subjecting rock to different s1 and s3 values and noting the shear and normal stresses at which failure occurred. Divided fields of stable stress from unstable stress. Mohr circles toughing failure envelope = critical states of stress
Curved line
Increase confining P, increase shear stress needed for failure (bigger mohr circle)
Pre-existing fractures
weaken rocks
envelope moves - less stress required to fail.
Veins may strengthen a rock
Coloumb criterion
T=c+uσ(n)
Positive linear relationship - predicts linear Mohr failure
-c=cohesion
u=coefficent of internal frication
Griffith Criterion
Rocks in experiments do not fail according to Coulomb criterion when σ3 is low/negative (tensile)
Griffiths failure criterion - idea that failure occurs from the opening and linkage of micro cracks.
Rounded end to failure envelope, failure closer to origin than predicted. Rocks are stronger in compression than tension.
Failure criterion
Low confining pressure - Griffiths criterion
Higher confining pressure - Coulomb criterion
-ve numbers - tension failure (steep dip) - surface
+ve numbers - shear failure
very +ve numbers - ductile (30-40) - depth
Faults are steep at the surface and shallow out with depth - Mohr circle
Pore fluid pressure on Mohr failure envelope
Pore fluid P’s decrease values of σ1/3. Thus, elevated pore fluid pressures make it possible to have movement on very low angle faults.
Shifts Mohr circle to less differential stress needed to overcome shear resistance on plane.
More fluid = more chance of failure (closer to vertical axis)
Fracking - reservoir stimulation
Use pore fluids in oil/gas industry to weaken rocks are force Mohr circle to move.
Joint sets
Joints often systematic - parallel, similar spacings; discrete sets
Can be 100s m in length
Extensional and have no measurable shear displacement
Common
Tensile failure (steeply dipping), low stress and strain (low confining P, -ve)
Joints are planes. Shape controlled by bed thickness (thicker units more circular, thinner units start and stop at edge of bed)
Joints often occur in association with folds. Timing of joints can be synchronous with folding or can be uplift and unloading related structures which form much later than faults.
Transverse joints
Diagonal joints
Longitudinal joints
Joint systems
Master/systematic joints - long, pronounced, older than others, regular spacing
Cross-joints - shorter and occur between systematic joints, less regular
Orthogonal join formation - strain during development of set 1 joints causes σ1 and σ3 to flip. Only possible if differential stress is low (shallow)
Joints can be dissolved out by meteoric water to provide fissures (limestone). They can have major impacts on flow within flows.
Joint spacings
Decreases with layer thickness and strain
More joints in stiffer layers
Joints often confined to beds
Each joint relieves part of total stress acting on rock. Each joint acts within its own bed.
Joint in thicker bed has longer length, relieves more stress
Join in thinner bed has thinner length, so relieves less stress
Thus, to relieve total stress evenly, more joints with closer spacing required in thin beds relative to thick beds.
Hackle joints
Joints start at origin and propagate away.
Ribbing if front stops, before more stress applied and propagates out again
Sheeted joint
Like unpeeling onion
Pluton forms, lifted to surface, stresses relieved and form parallel to pluton
Removal of overburden allows expansion of rock by fracturing
Cooling joints
Hot magma hits air and cools slowly from outside in. Become stressed and fracture.
