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Unity University Faculty of Engineering

Department of Mining Engineering

GENERAL GEOLOGY (Geol 2081)

Chapter 4:

STRUCTURAL GEOLOGY AND TECTONICS

Tadesse Alemu Director

Basic Geoscience Mapping Directorate Geological Survey of Ethiopia (tadessealemu@yahoo.com)

November 2012 Addis Ababa

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4. STRUCTURAL GEOLOGY AND TECTONICS  The crust of the earth is deformed at many scales, locations, and times; this deformation produces identifiable structures in the crust such as fractures and folds. An appreciation of earth structures has both enormous practical value and profound intellectual implications for how we view this planet. This chapter deals with ways to recognize and characterize major and minor structures in the earth's crust and ways to gain insight into how these structures form. It develops skills in three-dimensional thinking that are essential for understanding crustal structures. It also explores techniques for determining the sequence in which structures form. The chapter will also focus on macroscopic structures but will also introduce to some of the fascinating structures that form at the microscopic scale. Our ability to understand geologic structures depends in large part on how we perceive them. Few geologic structures form by trivially simple processes, but depending on how we view geologic structures, they can appear horribly complicated or amenable to understanding; perspective is critically important. Structural geology and Tectonics is the branch of geology that deals with:

• Form, arrangement and internal architecture of rocks • Description, representation, and analysis of structures from the small to moderate

scale • Reconstruction of the motions of rocks • Structural geology provides information about the conditions during regional

deformation using structures • Both are concerned with the reconstruction of the motions that shape the outer

layers of earth • Both deal with motion and deformation in the Earth’s crust and upper mantle • Tectonic events at all scales produce deformation structures • These two disciplines are closely related and interdependent

Tectonics: Study of the origin and geologic evolution (history of motion and deformation) of large areas (regional to global) of the Earth’s lithosphere (e.g., origin of continents; building of mountain belts; formation of ocean floor)

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Structural Geology: Study of deformation in rocks at scales ranging from submicroscopic to regional (micro-, meso-, and macro-scale). It describes a geometric feature in a rock whose shape, form, and distribution, which can be described as:

Microstructure: The small-scale arrangement of geometric and mineralogical elements within a rock. Texture: Preferred orientation of crystallographic axes in the sample. Microfabric: Comprises the microstructure and the texture of a material.

Fundamental Structures are:

Contacts Primary Structures Secondary structures Fractures (Joints and Faults) Vein Fold

Structural Analyses consist of:-

Descriptive: Recognize, describe structures by measuring their locations,

geometries and orientations Break a structure into structural elements - physical & geometric

Kinematic: Interprets deformational movements that formed the structures

• Translation, Rotation, Distortion, Dilation Dynamic:

Interprets forces and stresses from interpreted deformational movements of structures

 

4.1. Introduction to Deformation, Stress and Strain  Deformation: changes in shape, position, and/or orientation of a body. It includes all changes in the original location, orientation or form of a crustal rock body. Homogeneous deformation: the displacement gradient is a constant throughout the deformed body. For a homogeneous deformation, initially straight lines remain straight, circles become ellipses and parallel lines remain parallel after deformation.

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Compressional stress:

• “Push-together” stress. • Shortens and thickens crust. • Associated with orogenesis (mountain building).

Tensional stress:

• “Pull-apart” stress. • Thins and stretches crust. • Associated with rifting.

Shear stress:

• Slippage of one rock mass past another. • In shallow crust, shear is often accommodated by bedding planes.

Strain:

• Changes in the shape or size of a rock body caused by stress. • Strain occurs when stresses exceed rock strength. • Strained rocks deform by folding, flowing, or fracturing.

Basic Measures of Strain I. Change in linear dimension: the relative changes of lines

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2. Quadratic elongation λ = ⎨L 1 ⎬ 2 = (1ε+)2 dimensionless!

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γ = tan ψ Note: for small angular changes, tan ψ → ψ III. Change in volume (dilation): Δv = (V1- Vo)/Vo Strain Ellipsoid Is the visualization of state of finite strain at a point. Principal axes are lines that remain perpendicular before and after strain. Their lengths define the major, intermediate, and minor semi-axes of the strain ellipse Axes x>y>z (λ1>λ2>λ3). Strain Ellipsoid is visualization of strain tensor (2nd rank). A final state of "finite" strain may be reached by a variety of strain paths. Finite strain is final state; "incremental strains" represent steps along path or strain increments that result in final finite state of strain. Coaxial and Non-coaxial Strain In coaxial strain axes remain in same positions in deforming material whereas in non-coaxial strain they rotate through the material. Note that the overall shape of the final strain ellipse is the same, although the orientation is different.

Figure 4.3. Total strain ellipse with lines of no total longitudinal strain.

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Figure 4.4. Coaxial total strain ellipse with four zones, each of which has a different deformation history (after Ramsay, 1967).

Zones Lines Structures formed Zone 1a Lines that have been elongated only Boudinaged Zone 1b Lines that underwent early shortening

followed by elongation (net lengthening) Remnants of disrupted folds and isolated fold hinges

Zone 2 Lines that underwent early shortening followed by elongation (net shortening)

Folds that are becoming unfolded and boudinaged

Zone 3 Lines that have been shortened only Folds with large amplitudes and short wavelengths

Some Important Strain Concepts:

• With coaxial deformation all planes and lines (except principal ones) rotate towards plane of maximum flattening or line of maximum extension, respectively

• In simple shear, a type of noncoaxial deformation, this also holds true, and the plane of maximum flattening and the line of maximum extension all rotate towards the shear direction. Thus in high strain deformation everything moves to the same plane and line.

• Progressive structural history from shortening to extension is possible through a single coaxial deformation

• As a strain ellipse progresses through incremental strains to a finite strain state lines that were in quadrants of compression migrate to quadrants of extension. Thus a single line may undergo folding (compression) and subsequent boudinage (extension) as the strain ellipse continues to be flattened.

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HOW ROCKS DEFORM Elastic deformation – The rock returns to original size and shape when stress removed. When the (strength) of a rock is surpassed, it either flows (ductile deformation) or fractures (brittle deformation). Brittle behavior occurs in the shallow crust; ductile in the deeper crust. Factors controlling rock strength and deformation style:

• Temperature and confining pressure - Low T and P = brittle deformation - High T and P = ductile deformation

• Rock type – Mineral composition controls strength. • Time – Stress applied for a long time generates change.

4.2. Primary Structures  Structures of rocks that are present before the onset of deformation are called primary structures. They are original features of sedimentary or igneous rocks, resulting from deposition or emplacement. Structures reflecting subsequent deformation or metamorphism, which are the subject of most of this chapter, are secondary structures. Primary structures play an important role in the interpretation of the structure of deformed areas. Less common, but of considerable value in areas where they occur, are primary features that can be used to analyze the strain of the deformed rock. These include pebbles and fossils in sedimentary rocks; and vesicles, lapilli, and crystals in rocks of igneous origin. As well as acting as markers, certain primary structures can also provide very valuable additional information. These structures indicate the direction in which the surrounding rocks get younger or, as it is generally expressed, the younging direction of the sequence. Some of the more reliable and commonly occurring structures, from the point of view of younging criteria are discussed below; Cross Bedding Cross bedding is defined as a structure confined to a single sedimentation unit and characterized by internal bedding or lamination, called foreset bedding, inclined to the principal surface of accumulation. The type of cross bedding used to determine the

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Sole Marking The name sole marking, is given to certain irregularities in the interface between a pelite and the coarser material (conglomerate, sandstone, or limestone) stratigraphy overlying it. The structure is referred to as sole mark because it is generally observed on the original lower surface of the sandstone after the pelite has disintegrated and fallen away. These structures are commonly preserved in deformed rocks and recognition of them, on the underside of the sandstone, gives the direction of younging. Dessication Cracks Dessication cracks are a fairly common feature of sediments that have been deposited on land. They are very commonly associated with the ephemeral lakes of arid regions. They form when the water that has deposited the sediment drains away or evaporate, leaving the sediment to dry out subaerially. Examination of this structure in present-day sediments reveals that the cracks that develop during drying have polygonal form in plan and that, in section, the individual polygons of sediment become turned up at the edges, so that they have a concave upper surface. This form is commonly preserved in the rock and the upward concavity indicates the direction of younging.

4.3. Folds 

4.3.1. Introduction  Folds form from curving, buckling, and bending of originally planar rock layers (e.g., beds, foliation) through ductile deformation. Practically, folds are defined by the attitude of their axis and/or hinge line, axial plane. Folds occur in any geologic layer such as bedding, lava flow layers, and foliation. Folds range in size from mm to km, and are manifestations of ductile deformation (i.e., form at depth where T, P are high and fracturing does not occur). Parts of folded Surface

• Hingeline: Lines joining points of maximum (tightest) curvature • Axial plane/surface: The surface joining the hinglines in adjacent folded surfaces • Fold axis: Parallel to the hingeline • Profile plane: is the section drawn perpendicular to the fold axis • Inflection lines: Lines connecting points of zero curvature • Wavelength: distance between two inflection points or crest to crest distance

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• Amplitude: distance from medial surface to hinge line

Figure 4.6. Parts of a fold.

4.3.2. Classification and nomenclature of fold  Based on geometry:

– Antiform: Folds with upward closure – Synform: Folds with downward closure – Anticline: A fold with the oldest rocks at the core (i.e., at the concave

side). – Syncline: A fold with the youngest rocks at the core.

*Note: In simply-folded areas, anticlines are generally antiformal, and synclines are synformal.

However, in refolded areas this is not generally the case. – Antiformal syncline – Synformal anticline

Fold shape in profile (perpendicular to hinge line): • Parallel folds: no change in layer thickness as measured perpendicular to layer.

Therefore layer not flowing--clue to mechanical properties.

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• Concentric folds: layers are in a sense don't continue for long distances along down

axial surface perpendicular to axis due to "room problem".

• Similar folds: folds in which the layer thickness does not change parallel to the axial

surface. Clearly layer must flow internally to change shape.

• Kink (m to microscale) and Chevron folds (several m to regional scale): Planar limbs, sharp hinges, composed of many thin layers.

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4.3.3. Types of Fold 

• Polyclinal Fold – Folds with more than two axial plane (rare)

• Conjugate fold: Has converging paired axial surfaces – Axial planes intersect along the axis (if cylindrical) – Axial plane may displace another axial plane

• Box Fold : Conjugate folds with round hinge zones

• Kink Fold: Conjugate fold with sharp hinge zones

• Monocline: a local steepening in otherwise uniformly dipping strata. • Isoclinal fold: limbs are parallel to the axial plane.

• Recumbent fold: fold with horizontal axial plane. Commonly isoclinal. • Symmetric folds: are with the median plane and the axial plane is perpendicular,

and the axial plane divides the fold into mirror quarter waves. They are commonly polyharmonic fold with fold waves with two or more orders of wavelengths and amplitude. Large polyharmonic folds have parasitic (smaller) fold.

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• Asymmetric folds: is defined as one in which the axial plane is inclined. The asymmetry of smaller folds is represented as Z-, S-, W-, or M-shaped. These asymmetries are defined looking down-plunge, or projected onto a horizontal plane. The asymmetries of the smaller folds can be used to determine the closure of the larger fold.

Figure 4.7. Characteristic fold vergence of parasitic fold across a large-scale antiform.

4.4. Foliations  Foliation: Any type of planar fabric in rock, including bedding, cleavage, schistosity. Foliations are penetrative (occur throughout) in samples at 10's of cm scale. Thus faults are not foliations, nor are fractures and joints because the latter are simply fractures and not related to internal structure of rock. Cleavage: Secondary fabric element (not bedding) formed under low grade metamorphic conditions (or less) that allows the rock to split along planes. Foliations commonly developed in plane of maximum flattening of strain ellipsoid or perpendicular to direction of maximum shortening: Strain Ellipsoid Axes: X>Y>Z, so foliation commonly in X-Y plane and perpendicular to Z. There are three common types of foliation. These are:-

• Axial plane foliations • Shear (Mylonitic) foliations • Transposed foliations

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Axial plane foliations Axial plane foliations are referred to as the surface generally parallel to the axial plane of the fold in the hinge area. However, it is important to realize that axial plane foliations commonly are not strictly parallel to the axial planes of folds. A number of morphological axial plane foliation types are generally recognized and named. Due to morphological gradation between the types of foliations, it appears difficult to have complete agreement in the literature as to the meaning of some of these names. However, for convenience the type of axial plane foliations are divided into; (i) continuous, (ii) spaced or disjunctive, and (iii) crenulation. Continuous

• Slaty and phyllitic cleavage: very closely spaced cleavage planes on < 0.1 to perhaps 1 mm scale at the upper limit. A smooth highly penetrative foliation consisting of parallel orientations of clays or micas. Penetrative to microscopic scale. It is a term that is used to describe the foliation in fine grained rocks belonging to the Greenschist facies.

• Schistosity: progressive more coarsely crystalline penetrative foliation. • Gneissic Layering: Strong compositional banding defines foliation. High grade

metamorphic rock. Spaced or Disjunctive Are broadly spaced cleavage domains (> 1 mm to perhaps10 of cm apart) with intervening uncleaved areas called microlithons. Various adjectives describe styles of surfaces including: sutured (stylolytic), wavy, anastomozing, and planar. Crenulation Cleavage Is formed due to microfolding of an existing foliation (commonly a slaty or phyllitic cleavage), typically due to shortening of foliation about parallel to its layering. Crenulations are either symmetrical or asymmetrical.

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Figure 4.8. Two basic categories of crenulation cleavage. (a) asymmetrical crenulation cleavage, and (b) asymmetrical crenulation cleavage

Shear (Mylonitic) foliations They are developed due to high shear strain, usually in a fault or shear zone, but sometimes applied to fabrics extending over large regions. Transposed foliations Another type of foliation that occurs parallel to the axial plane of folds, but it differs from most axial plane foliations, in that it is largely defined by a layering that predates the folding. An essential part of the transposition process is rotation of a preexisting foliation by folding into an orientation approximately parallel to the axial plane of the folds.

Figure 4.9. Transposition of quartz-rich layers in quartz-mica schist (Photograph by T. Alemu)

A B

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Strain Significance of Foliations Commonly in plane of maximum flattening of strain ellipsoid or perpendicular to direction of maximum shortening: Strain Ellipsoid Axes: X>Y>Z, so foliation commonly in X-Y plane and perpendicular to Z

4.5. Lineations  Lineation is a fabric element that can be represented by a line. Forms and Types of Lineations

• Fold hinges, • Mullions: cusps and bulges between contrasting lithologies due to mechanical

incompatibilities • Rods: preserved fold hinges • Boudins: lineations formed by stretching and necking of a layer.

Figure 4.10. Type of lineations: (a) Simple linear fabric defined by preferred orientation of linear bodies. (b) Combined lineation and foliation defined by preferred orientation of elongate tabular bodies. (c) Linear fabric defined by common axis of variably oriented, tabular bodies. (d) Linear fabric defined by penetrative folding. (e) Lineation defined by intersection of two foliations.

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4.6. Joints  Joints are fractures along which there has been no appreciable displacement parallel to the fracture and only slight movement normal to the fracture plane. Joints are most common of all structures present in all settings in all kind of rocks as well as consolidated and unconsolidated sediment. Characteristics of fractures according to Pollard and Aydin (1988):-

• They have two parallel surfaces that meet at the fracture front. These surfaces are approximately planar

• The relative displacement of originally adjacent points across the fractures is small compared to the fracture length.

• Cross cutting relationship and material filling the fractures can help in resolving the chronological order of deformation.

Most joints form by extensional fracturing of rock in the upper few kilometers of the Earth's crust. The limiting depth formation of extension fractures should be the ductile-brittle transition. Joints can be caused by a number of processes that create tensional effective stress in rock:

Uplift and erosion Residual stress Tectonic deformation Natural hydraulic fracturing

If joints are filled they are called veins. The filling range from quartz and feldspar (pegmatite and aplite) to quartz, calcite and dolomite. Importance of studying joints:

To understand the nature and sequence of deformation in an area. To find out relationship between joints and faults and or folds. Help to find out the brittle deformation in an area of construction (dams, bridges,

and power plants). In mineral exploration to find out the trend and type of fractures and joints that

host mineralization which will help in exploration. Joints and fractures serve as the plumping system for ground water flow in many area and they are the only routes by which ground water can move through igneous and metamorphic rocks.

Joints and fractures porosity and permeability is very important for water supplies and hydrocarbon reservoirs.

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Joints orientations in road cuts greatly affect both construction and maintenance. Those oriented parallel to or dip into a highway cut become hazardous during construction and later because they provide potential movement surfaces.

4.6.1. Terminology of Joints  Conjugate joints Conjugate fractures are paired fracture systems, formed in the same time, and produced by tension or shear. Curved joints Occur frequently and may be caused by the textural and compositional differences within a thick bed or large rock mass or they may a result of changes in stress direction or analysis. Tectonic joints Form at depth in response to abnormal fluid pressure and involve hydro fracturing. They form mainly by tectonic stress and the horizontal compaction of sediment at depth less than 3 km, where the escape of fluid is hindered by low permeability and abnormally high pore pressure is created. Hydraulic joints Form as tectonic fractures by the pore pressure created due to the confined pressed fluid during burial and vertical compaction of sediment at depth greater than 5 km. Filled veins in low metamorphic rocks are one of the best of examples of hydraulic fractures. Unloading joints Form near surface as erosion removes overburden and thermal elastic contraction occurs. They form when more than half of the original overburden has been removed. The present stress and tectonic activity may serve to orient these joints. Vertical unloading fractures occur during cooling and elastic contraction of rock mass and may occur at depths of 200 to 500 m. Release joints Similar to unloading fractures but they form by release of stress. Orientation of release joints is controlled by the rock fabric. Released joints form late in the history of an area and is oriented perpendicular to the original tectonic compression that formed the dominant fabric in the rock.

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Release joints may also develop parallel to the fold axes when erosion begins and rock mass that was under burial depth and lithification begins to cool and contract, these joints start to propagate parallel to an existing tectonic fabric. Sheared fractures may be straight or curved but usually can't be traced for long distance. Nontectonic joints Sheeting joints:

• Those joints form subparallel to the surface topography. These joints may be more observed in igneous rocks. Pacing within these fractures increases downward. These fractures thought that they form by unloading overlong time when erosion removes large quantities of the overburden rocks.

Columnar joints and Mud Cracks:

• Columnar joints form in flows, dikes, sills and volcanic necks in response to cooling and shrinking of the magma.

4.6.2. Genetic classification of Joints  From the point of view of fracture mechanics, crack tips have been related to three modes of displacement, namely extensional or Mode I displacement, and shear fractures of Modes II and III.

• Mode I joints: it is the extensional fractures and formed by opening with no displacement parallel to the fracture surface. In extensional fractures the fracture plane is oriented parallel to σ1 and σ2 and perpendicular to σ3.

• Mode II and Mode III joints. are faults like fractures one of them is strike -slip and the other is dip-slip. Same fracture can exhibit both mode II and mode III in different parts of the region.

Figure 4.11. Three modes of displacement of joints. Extensional or Mode I displacement (a), and shear fractures of Modes II and III (b and c).

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Systematic joints: have a subparallel orientation and regular spacing. Joint set: joints that share a similar orientation in same area. Joint system: two or more joints sets in the same area Nonsystematic joints: joints that do not share a common orientation and those

highly curved and irregular fracture surfaces. They occur in most area but are not easily related to a recognizable stress.

Plumose joints: those have feathered texture on their surfaces and from this texture the direction of propagation of joints can be determined.

**Some times both systematic and nonsystematic joints formed in the same area at the same time but nonsystematic joints usually terminate at systematic joints which indicate that nonsystematic joints formed later.

4.6.3. Relation of joints to other structures  Joints may form during brittle folding in a position related to the fold axis and axial surface as follows (Fig. 4.12)

parallel normal oblique

Figure 4.12.Position of joints related to the fold axis and axial surface

Joints in fold-thrust belts (orogens) seem to form at depth under high pore pressure. Many form parallel to σ1 and perpendicular to folds and strike, which is σ3. Joints are also formed adjacent to brittle faults, and movement along faults usually produces a series of systematic joints. Fractures form in pluton in response to cooling and later tectonic stress. Many of these joints are filled with hydrothermal minerals as late stage

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products. Different types of joints are present with pluton (i.e. longitudinal, and cross joints)

4.7. Faults  Faults are fractures that have appreciable movement parallel to their plane. Understanding faults is useful in design for long-term stability of dams, bridges, buildings and power plants. Faults may be hundred of meters or a few centimeters in length. Their outcrop may have as knife-sharp edges or fault/shear zone. Fault/shear zones may consist of a serious of interleaving anastomozing brittle faults and crushed rock or of ductile shear zones composed of mylonitic rocks.

4.7.1. Fault terminology   Parts of the fault consist of:- • Fault plane: Surface that the movement has taken place within the fault.

• Hanging wall: The rock mass resting on the fault plane.

• Footwall: The rock mass beneath the fault plane.

• Slip: Describes the movement parallel to the fault plane (fault displacement).

• Dip slip: Describes the up and down movement parallel to the dip direction of the

fault.

• Strike slip: Applies where movement is parallel to strike of the fault plane.

• Oblique slip: Is a combination of strike slip and dip slip.

• Net slip (true displacement): Is the total amount of motion measured parallel to the direction of motion

• Separation: The amount op apparent offset of a faulted surface, measured in specified direction. There are strike separation, dip separation, and net separation.

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• Heave: The horizontal component of dip separation measured perpendicular to strike of the fault.

• Throw: The vertical component measured in vertical plane containing the dip. • en-echelon faults: Faults that are approximately parallel one another but occur in

short unconnected segments, and sometimes overlapping. • Radial faults: faults that are converge toward one point. • Concentric faults: faults those are concentric to a point. • Bedding faults (bedding plane faults): follow bedding or occur parallel to the

orientation of bedding planes.

Figure 4.13. Fault terminologies.

Figure 4.14. Sketches illustrating differences between faults, fault zones, shear zones. (a) Fault, (b) Fault zone with inset showing the cataclastic deformation adjacent to the fault surface, (c) Sketch illustrating the relation between principal fault and fault splay, (d) Anastomozing faults in a fault zone, (e) A shear zone, showing the rock continuity across the zone.

Footwall Hanging wall

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4.7.2. Nature of movement along faults  Fault Slip (also called displacement): True displacement between two formerly continuous points (piercing points) on either side of fault. Piercing Point-linear feature providing tie across fault Fault Separation: Distance between displaced parts of a marker horizon as measured along a specified line (sometimes called apparent slip). Throw and Heave: Throw up and heave out respectively vertical and horizontal components of dip separation. Terms valuable because at times it is convenient to measure horizontal or vertical separation, for example from maps or boreholes. Criteria for faulting

• Repetition or omission of stratigraphic units asymmetrical repetition • Displacement of recognizable marker such as (fossils, color, composition, texture

.etc.). • Truncation of structures, beds or rock units. • Occurrence of fault rocks (mylonite or cataclastic or both) • Abundant veins, silicification or other mineralization along fracture may indicate

faulting. • Drag Units appear to be pulled into a fault during movement (usually within the

drag fold and the result is thrust fault) • Reverse drag occurs along listric normal faults. • Slickenside along a fault surface • Topographic characteristics such as drainages that are controlled by faults and

fault scarps.

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Figure 4.15. A. Thrust fault resulting in repeated section in a vertical drill hole. B. Normal fault resulting in missing section in a vertical drill hole.

Figure 4.16. Change in fault character with depth for a steeply dipping fault. Note the change in fault zone width and types of structures with depth.

Type of Faults Anderson 1942 defined three fundamental possibilities of stress regimes and stress orientation that produce the three types of faults (Normal, thrust, and strike-slip). Note that σ1> σ 2> σ 3.

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Figure 4.17. Anderson's theory of Fault Mechanics: (a) high-angle normal faults, (b) low-angle reverse (thrust) faults, (c) Strike-slip faults

4.7.3. Normal faults  The hanging wall has moved down relative to the footwall. σ1 is vertical and σ2 and σ3 are horizontal. Shear plane is oriented 45º or less to σ1 and // σ2. Form due to horizontal extension or vertical compression.

• Graben: consists of a block that has dropped down between two subparallel normal faults that dip towards each other.

• Horst: consists of two subparallel normal faults that dip away from each other so that the block between the two faults remains high.

• Listric: are normal faults that frequently exhibit (concave-up) geometry so that they exhibit steep dip near surface and flatten with depth.

Normal faults usually found in areas where extensional regime is present.

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Figure 4.18. Contrasting models of rifting. (a) Cross section illustrating the old concept of symmetric horsts and grabens. (b) Cross section illustrating the contemporary image of tilted fault blocks and half grabens, all above a detachment.

Many steeply dipping faults appear to intersect downward into a shallowly dipping detachment, or "low angle normal fault". This contrast with the simple view of steep, planar normal faults that disappear into the earth (cf. Fig. 4.18b). Faults my be listric or change dip (from steep to shallow)with depth (this is also true of thrust faults). Faults commonly dip in same direction towards the bedding is called synthetic and faults dip in the opposite direction called Antithetic.

4.7.4. Thrust faults  The hanging wall has moved up relative to the footwall (dip angle 30º or less). σ1 and σ2 are horizontal and σ3 is vertical. Thus a state of horizontal compression is defined for thrust faults. Shear plane is oriented to σ 1 with angle = or < 45º and // σ 2. Reverse Faults: Are similar to the thrust faults regarding the sense of motion but the dip angle of the fault plane is 45º or more. Thrust faults usually formed in areas of compressional regime. Why Study Thrust Faults

• Host the largest, and potentially most destructive earthquakes (at subduction zones). Low dip requires that faults have large surface area in brittle "seismogenic zone" and that this surface area is close to ground surface where we live.

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• Associated with mountain building and collisional tectonics. Low-angle faults with big displacements have been known since 1800's.

• Large displacements and mechanical paradoxes. • Influence positions of ore deposits and hydrocarbons.

Common Terminologies of Thrust faults:

• Bedding-step thrusting: Faults consist of flats parallel to bedding (surface of weakness) and ramps where the fault cuts across the bedding

• In sequence thrusting: As thrust system continues to form it converges on an adjacent sedimentary basin (foreland) and incorporates the basin by "in sequence" thrusting. The thrusts tend to break forward towards the basin because it is composed of sediment weaker than that incorporated in the thrust system.

• Out-of-sequence thrusts: Are thrusts that form later in the interior of the wedge of previously thrust-faulted material.

• Duplex: In thrust systems, thrust sheets bounded by a roof thrust and sole thrust, more generally a block bounded by faults.

Figure 4.19. (a) Single thrust system. (b) A duplex

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Tectonic Setting of Thrust Systems

• Occur in zones of plate convergence • Both subduction thrusts and "foreland thrusts" • Thrust systems especially well developed in collision zones • Also at the toes of landslides, whether on land or in the deep sea

Figure 4.20. (A). With the onset of convergence, an accretionary prism develops that verges towards the trench, and a back-arc fold-thrust belt forms cratonward of the volcanic arc and verges towards the upper-plate craton. (B). Eventually the two continents collide. A fold-thrust belt forms in the foreland of the orogen. Slivers of obducted ocean crust may separate lower-plate rocks from the metamorphic hinterland of the orogen and define the suture between the two plates.

4.7.5. Strikeslip faults  Faults that have movement along strikes. σ1 and σ3 are horizontal and σ2 is vertical. Shear plane is oriented to σ1 with angle 45º and // σ3. They Form also due to horizontal compression. There are two types of strike slip faults:

• Right lateral strike-slip fault (dextral): Where the side opposite the observer moves to the right.

• Left lateral strike-slip fault (sinistral): Where the side opposite the observer moves to the left.

A

B

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Figure 4.21. Strike-slip faults. (A). Sinistral strike-slip fault. (B). Dextral strike-slip fault. (C). Relation to principal stress axes.

Types of strike-slip faults Transform Faults: Are a type of strike-slip fault (defined by Wilson 1965). They form due to the differences in motion between lithospheric plates. They are basically occurring where type of plate boundary is transformed into another. Types of transform faults are:

• Ridge-Ridge • Ridge-Arc • Arc-Arc

Transcurrent Faults: are types of strike-slip faults, which are confined to the crust • Indent-linked faults • Tear faults • Transfer faults

Features of Strike-Slip Faults

• Restraining or compressional bends-- folds and thrusts • Releasing or extensional relays--- depression • Pull-apart basins

A B

C

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Figure 4.22. Stepover along strike-slip fault. (a) Sketch map showing how regional dextral strike-slip fault can be distributed along fault segment that are not coplanar. Slip is relayed from one segment to another at a stepover. (b) At a restraining stepover compression and thrusting occur. (c) At a releasing stepover, extension and subsidence occur.

Figure 4.23. Map view models of fault bending along strike-slip faults. The “edges” of the crustal blocks are provided for references. (a) Releasing bend at which normal faults and pull -apart basin have formed. (b) Restraining bend at which thrust faults have formed.

4.7.6. Shear zones  Shear zones are produced by both homogeneous and inhomogeneous simple shear or oblique motion and are thought of as zones of ductile shear. Shear zones are classified by Ramsay (1980) as:

1) Brittle 2) Brittle-ductile 3) Ductile

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Figure 4.24. Brittle (a) to Brittle-ductile (b-c) and ductile (d) deformation, reflecting the general subdivisions of shear zones (after Ramsay, 1980).

• Shear zones on all scales are zones of weakness. • Associate with the formation of mylonite. • Presence of sheath folds. • Shear zones may act both as closed and open geochemical systems with respect to

fluids and elements. • Shear zones generally have parallel sides. • Displacement profiles along any cross section through shear zone should be

identical. INDICATORS OF SHEAR SENSE OF MOVEMENT

Rotated porphyroblasts and porphyroclasts. Pressure shadows Fractured grains. Boudins Presence of C- C’ and S- surfaces (parallel alignment of platy mineral) Riedel shears.

Figure 4.25. Characteristic geometry of (a) C-S and (b) C-C’ structures in dextral shear zones. The C-surface is parallel to the shear zone boundary and is a surface of shear accumulation (i.e., not parallel to a plane of principal finite strain). The S-foliation is oblique to the shear zone boundary and may approximate the XY-plane of the finite strain ellipsoid. The C’-foliation in (b) displaces earlier foliations (C or composite C/S).

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Figure 4.26. C-S Fabrics (foliation and shear surfaces). Note inclined quartz grains (S fabric) with flat-lying mica-rich shear zones (C- fabric). Quartz grains are undergoing dynamic recrystallization at edges. Interiors show undulose extinction showing crystal plastic deformation due to dislocations movement. Left lateral offset is deduced from the C-S relationship.

4.8. Introduction to Plate Tectonics  Plate tectonics is a unifying theory that attempts to explain natural phenomena such as earthquakes and volcanoes. The earth's surface had been mapped into a series of plates (Fig. 4.27). The seven major plates are: Eurasian, Pacific, Australian, North American, South American, African and Antarctic - all comprise both oceanic and continental crust. For example, the North America Plate includes most of North America plus half of the northern part of the Atlantic Ocean. (The Pacific Plate is almost entirely oceanic, but it does include the part of California which lies to the west of the Sand Andreas Fault.) There are also numerous small plates (e.g., Jaun de Fuca, Nazca, Scotia, Philippine, Caribbean). Boundaries between these plates are of three types: divergent (i.e., spreading), convergent, and transform.

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Figure 4.27. Map showing the distributions of the Earth’s plates.

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PLATE BOUNDARIES The geological processes that take place at different boundaries are described below. Before going there, however, it is important to recognize that plates are not just pieces of continental or oceanic crust, but that, along with the crustal rock, they include a considerable thickness of the rigid upper part of the mantle. Together, the crust and the rigid part of the mantle make up the lithosphere, which has a total thickness of approximately 100 km. At spreading centers, the lithospheric mantle may be very thin because the upward convective motion of hot mantle material generates temperatures that are too high for the existence of a significant thickness of rigid lithosphere. The fact that the plates include both crustal material and lithospheric mantle material makes it possible for a single plate to be comprised of both oceanic and continental crust. For example, the North American Plate includes most of North America, plus half of the northern Atlantic Ocean. Similarly the South American plate extends across the western part of the southern Atlantic Ocean, while the European and African plates each comprise part the eastern Atlantic Ocean. Immediately beneath the base of the lithosphere lies the partial melting zone (the low velocity zone) of the upper mantle - which is part of the asthenosphere. It is thought that the relative lack of strength and rigidity of the partial melting zone facilitates the sliding of the lithospheric plates. Divergent Boundaries Divergent boundaries (4.28) are spreading boundaries, where new oceanic crust is created from molten mantle material. Most are associated with the oceanic-ridges, and the crustal material created at a spreading boundary is always oceanic in character (Fig. 4.29).

• Spreading is caused by the convective movement within the mantle, which has the effect of pulling the plates apart.

• Magma from the mantle pushes up to fill the voids left by spreading.

• A variety of volcanic rocks (all of similar composition) are created in the upper

part, including pillow lavas which are formed where magma is pushed out into sea-water. Beneath that are vertical dykes intruded into cracks resulting from the spreading. The base of the oceanic crust is comprised of gabbro (i.e., mafic intrusive rock) (Fig.4.29). By oceanic we mean that it is mafic igneous rock (e.g., basalt or gabbro, rich in ferro-magnesian minerals) as opposed to the felsic igneous rocks (such as granite, which is dominated by quartz and feldspar) which

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are typical of continental areas. Another term for mafic igneous rock is SIMA (silicon and magnesium rich), and another term for felsic igneous rock is SIAL (silica and aluminum rich).

• Spreading rates vary quite considerable, from 2 to 4 cm/y in the Atlantic, to between

6 and 18 cm/y in the Pacific.

Spreading starts within a continental area with up-warping or doming, fracturing in a radial pattern - with three arms, and formation a rift valley (such as the Rift Valley in eastern Africa). It is suggested that this type of valley eventually develops into a linear sea (such as the present day Red Sea), and finally into an ocean (such as the Atlantic). A major continental rift is assumed to be initiated by a series of hot spots. Each hot spot has an associated three-arm rift, but in most cases only two of these arms will continue to separate - the third one being termed a "failed arm". Some of these failed arms become major river channels. Rifting along a series of hot spots will then lead to continental rifting. It is thought that some 20 hot spots were responsible for the initiation of spreading along the mid-Atlantic ridge

Figure 4.28. Divergent boundary of two continental plates. Creates a rift valley (Example: East

African Rift).

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A)

B)

Figure 4.29. (A) Mechanism of sea floor spreading at Mid Oceanic Ridge. (B) Stratigraphy of oceanic crust.

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Convergent Boundaries Convergent boundaries, where two plates move towards each other, are of three types depending on what type of crust is present on either side of the boundary (i.e., ocean-ocean, ocean-continent or continent-continent) (Fig. 4.30). a) Ocean-Ocean (Fig. 4.30a): At an ocean-ocean convergent boundary one of the plates (ocean crust and lithospheric mantle) is pushed under, or subducted under the other. There is commonly an oceanic trench along the boundary. The subducted lithosphere descends into the hot mantle at a relatively shallow angle close to the subduction zone, but at steeper angles (up to about 45º) farther down. The significant volume of water within the subducting material (that includes ocean-floor sedimentary rock) mixes with the surrounding mantle. The addition of water to hot mantle lowers its melting point, and leads to the formation of magma. The magma, which is lighter than the surrounding mantle material, rises through the mantle and through the overlying oceanic crust to the ocean floor, to create a chain of volcanic islands known as an island arc. A mature island arc will develop into a chain of relatively large islands (such as Japan, or Indonesia) as more and more volcanic material is extruded and sedimentary rocks accumulate around the islands. Examples of ocean-ocean convergent zones are: subduction of the Pacific plate south of Alaska (Aleutian Islands), west of Kamchatka and Japan, west of the Philippines and in the northern part of New Zealand; subduction of the India-Australian plate south of Indonesia; and subduction of the Atlantic Plate beneath the Caribbean Plate. b) Ocean-continent (Fig. 4.30b): At an ocean-continent convergent boundary the oceanic plate is pushed under the continental plate in the same manner as an ocean-ocean collision. Similar geological features apply, and an offshore oceanic trench will normally be present. The mafic magma produced adjacent to the subduction zone will rise to the base of the continental crust and lead to partial melting of the crustal rock. The resulting magma will ascend through the crust producing a chain of largely volcanic mountains. Examples are: subduction of the Nazca plate under South America (which has created the Andes Range), and subduction of the Juan de Fuca plate under North America (creating mountains like Garibaldi, Baker, St. Helens, Ranier, Hood and Shasta – collectively known as the Cascade Range).

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Figure 4.30. Convergent plate boundaries. (A) Convergent boundary of two oceanic plates. Creates

an island arc and a trench (Example: Japan). (B) Convergent boundary of an oceanic plate and a continental plate. Form a volcanic mountain range and a trench (Examples: Cascades or Andes Mts). (C) Convergent boundary of two continental plates. Forms a folded mountain range (Examples: Appalachians, Alps, and Himalayas).

A

B

C

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c) Continent-continent Fig. 4.30c): A continent-continent collision occurs when a continent or large island has been moved along with oceanic crust (which was being subducted under another continent), and then collides with that other continent. The colliding continental material will not be subducted because it is too light (i.e., because it is composed largely of SIAL rocks), but the mantle convection system continues to operate, so the root of the oceanic plate breaks off and is absorbed into the mantle. There is tremendous deformation of the pre-existing continental rocks, and creation of mountains from that rock, from any sediments which had accumulated along the shores (i.e., within geosynclines) of both continental masses, and commonly also from some ocean crust and upper mantle material. Examples are: the collision of the Indo-Australian plate into the Eurasian plate, to create the Himalaya Mountains, and the collision of the African plate into the Eurasian plate, to create the Alps in Europe and the Zagros Mts. in Iran). Transform Boundaries Transform boundaries exist where one plate slides past another, without production or destruction of crustal material (Fig. 4.31). Most transform faults connect segments of mid-ocean ridges and are thus ocean-ocean boundaries. Some transform faults connect continental parts of plates. An example is the San Andreas Fault, which connects the Juan de Fuca ridge with the Gulf of California ridge. Transform faults do not just connect divergent boundaries. For instance the convergent boundary beneath the Himalayas is connected to the subduction zone beneath Indonesia by a transform fault, and the Queen Charlotte Fault connects the Juan de Fuca divergent boundary to the Aleutian subduction zone. Plate Boundary Summary

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Figure 4.31. Transform-fault boundary where the North American and Pacific plates are moving

past each other (Example: San Andreas Fault in California).

THE THEORY OF PLATE TECTONICS

It was following developments in the exploration of the ocean floor in the 1950s that new evidence was found to revive continental drift theory and led to the development of a theory to explain the movement of the continents - plate tectonic theory. The theory of Plate Tectonics is based around the idea that the crust is broken up into a series of large crustal plates which "float" on the asthenosphere below. Motions in the asthenosphere, called convection currents, cause plates to move away from each other at the rising limb of a convection current, forming a constructive plate boundary where new oceanic crust is formed. As plates continue to move outwards, eventually the oceanic plate may be subducted at a destructive plate boundary.

Supporting evidence for Plate Tectonics Theory:

1. Discovery of the Mid-Atlantic Ridge - Ocean floor mapping led to the discovery of a global mid-oceanic ridge mountain chain zigzagging around the continents. 2. Magnetic Variations on the Ocean Floor (Palaeomagnetism) - during cooling, minerals in the Basaltic rock, align themselves along the Earth's magnetic filed - forming a permanent record of magnetic field in the rocks. Periodic variations in the earth's

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magnetic field have produced almost symmetrical magnetic patterns in the rocks either side of the Mid-Atlantic ridge (alternating stripes of magnetically different rocks).

Figure 4.32. The theory of plate tectonics.

3. Theory of Sea-Floor Spreading - Hess, put forward an idea that mid-ocean ridges are a structurally weak point where magma is able to rise to the surface and where due to the upwelling and eruption of this material, new crust is created. This helps, to support the continental drift theory as it helps to explain how the continents may be moving, as they are carried on the 'spreading' ocean floor. Hess's theory was supported by the fact that the youngest rocks are nearest to the ridge (showing the present day magnetic polarity in their mineral alignment) and the oldest rocks (showing reversed polarity) are further away from the ridge. The problem with Hess's Theory was that "sea floor spreading" and the associated development of new oceanic crust, suggests that the earth is increasing in size, although we know this is not the case. Hess therefore realized that whilst in some areas, new crust was being formed; elsewhere old crust must be being destroyed. Examples of where this occurs were found around the Pacific Ocean with crust being subducted and destroyed at deep-sea ocean trenches.

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Therefore the theory suggests that newly formed oceanic crust moves away from a ridge rather like a conveyor belt and descends (millions of years later) into a deep-oceanic trench.

CONTINENTAL DRIFT AND PLATE TECTONICS

Since it was created, the earth has undergone many changes, including landform creation and destruction and changes in the relative position of the continental landmasses. Geologists have put forward many theories to explain earth surface changes but none have succeeded in providing universal explanation. The theory of plate tectonics which builds upon the concept of continental drift, however is a unifying theory that explains many of Earth's major geologic features. CONTINENTAL DRIFT Alfred Wegener proposed the theory of continental drift back in 1912. The theory suggests that there has been large-scale movement of continents across the globe and that during the Permian period, 225 million years ago, all the continents were joined as one super continent Pangaea. Around 200 million years ago, Pangaea split into Laurasia and Gondwanaland. The continents have continued to move and today's configuration of continents represents the most recent stage in their movement. Continental Drift theory is based on the following evidence

• THE JIGSAW FIT OF THE CONTINENTS - there is a noticeable jigsaw fit between many of our continents - for example, between the East Coast of South America and the West Coast of Africa, which suggests that at some point in time the continents were once assembled together.

• PLANT / ANIMAL FOSSILS - A number of identical fossils have been found distributed across the southern continents, again suggesting that they were once joined. Fossils of the Mesosauras dating back 280 million years ago have been found in South America and Africa (and nowhere else) - it is known from the fossil that this animal could not swim. Plant Fossils, such as Glossopteris (a tree) have been found in South America, Africa, India and Australia.

• GEOLOGICAL SEQUENCES - A number of continents show evidence of matching geological sequences with rocks of similar age, type, formation and structure occurring in different countries - e.g. the Appalachians (E USA) show a geological match with mountains in NW Europe and if they were fitted together would form a single continuous mountain belt.

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• CLIMATOLOGICAL ANOMALIES - A number of climatic anomalies have been discovered which suggest that continents must once have been in a different position and therefore have experienced a different climate. Coal which only forms under wet / warm conditions have been found beneath the Antarctica ice cap and there is evidence of glaciations in Brazil (which now has a equatorial climate) Whilst continental drift theory is very important, at the time it was not widely accepted due to the lack of explanation of how / what forces would be capable of resulting in moving the huge rock masses.

A)

B) Figure 4.33. Evidence for Continental Drift. (A) The Jigsaw Fit of the Continents. (B) The

distribution of plants/animals.

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THE RULES OF PLATE TECTONICS

1. Continental crust is less dense, or lighter, than Oceanic crust so it doesn't sink. It is never destroyed and is considered permanent. 2. Oceanic crust is heavier so it can sink below Continental crust. It is constantly being formed and destroyed at ocean ridges and trenches. 3. Continental crust can carry on beyond the edges of the land and finally end far below the sea. This explains why the edges of all the continents don't have deep trenches right up against their coastlines. 4. Plates can never overlap. This means that they must either collide and both be pushed up to form mountains, or one of the plates must be pushed down into the mantle and be destroyed. 5. There can never be gaps between plates, so if two plates move apart, as in the middle of the Atlantic, new rock will be formed to fill the space. 6. We know the Earth isn't getting bigger or smaller, so the amount of new crust being formed must be the same as the amount being destroyed. 7. Plate movement is very slow. This is partly why Wegener's original ideas were ignored. Nobody could 'see' the continents moving. When the plates make a sudden movement we call it an Earthquake, and it's the only time we are directly aware of the plates moving.