Fault Inter

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INTRODUCTION

Today, in mature petroleum provinces, we seldom or never find classical anticlineswith simple four-way closure. All those structures have been found long ago.Nowadays, our standard fare consists of structures bounded in part by dip and inpart by faults. We must become expert in recognizing such structures, and indeciding whether or not they have potential as hydrocarbon traps.

Conduits and Seals

Where a crack occurs in the earth, we usually call it a fracture. Where the rock layersare displaced across the fracture, we usually call it a fault. The importance of thefault, as distinct from the fracture, is that the fault has greater potential to form atrap ( Figure 1 ).

Figure 1


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At the time of cracking, the fault is likely to provide an escape for water (andpossibly hydrocarbons) contained in the rocks below; it provides a conduit. For thefault to form a trap, however, the fault must provide a seal.

An attractive situation arises where a fault is a conduit in its lower reaches, and aseal in its upper reaches. This may happen, for example, if the conduit formed by the

crack allows the escape of mineral-rich waters from below, and the conditions ofchemistry and temperature cause the minerals to precipitate out of solution as thewater rises in the fault. If a seal forms just above a faulted reservoir we have a trap,while if the conduit remains open below the reservoir we have a path by which thereservoir may be charged from deep source rock.

The danger in a fault trap, of course, is that the seal may be broken by laterreactivation of the fault, so that the hydrocarbons are lost.

Sometimes, therefore, we need faults to act as conduits, allowing the charge of areservoir from a deep source. In other situations, we need faults to act as seals. Andour constant enemy is the breach of a fault seal by reactivation of the fault.

In Figure 1, we depend on the fault to act as a conduit, to charge the reservoir fromthe deep source rock. We depend on the offset cap rock at A to provide an updipbarrier. But we also depend on a fault seal at B. If this seal does not exist, we havenothing. if it does, the spill point, C, depends on the thickness and offset of the caprock.

In Figure 2, with a thicker cap rock, we do not need the fault seal at B, but it mustbe present higher in the section, perhaps at D.


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Figure 2

If it is not present by E, the trap cannot be charged unless both the unconformitysurface and its overlying rocks are impermeable-and then only after the developmentof this impermeability. If the fault has been reactivated as suggested at F, anycharge depends on the subsequent sealing of this breach, and can represent onlyhydrocarbons generated since the resealing.

We see, then, the critical decisions we must make in fault interpretation. Is itpossible for the fault trap to be sourced? Does the fault plane seal? Did the sealdevelop before or after the generation and migration of the hydrocarbons? Wherethe thickness and offset of the cap rock imply a spill point, is the trapped volumesufficient to warrant drilling? Has the fault been reactivated? Is it likely to haveresealed? When? Were hydrocarbons still migrating after this time? Having regard tothe burial history, should we expect oil or gas?

In general, there are uncertainties in our answers, and fault traps always carry anadditional element of risk. Some decision-makers, as a matter of principle, will notdrill downthrown fault traps, or traps where faults can be seen extending to thesurface. Others point to the large number of proven fields having just theseconditions. As always, the drilling decision should be based on a reasonedassessment of potential reward and potential risk.


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One other seismic decision is very relevant to the assessment of risk: is thereevidence of any mobile rock above the trap, in which any fault displacement could beabsorbed? Thus, the presence of thick salt above the unconformity in Figure 2 wouldvirtually remove any risk from fault reactivation (though in Figure 1 it might not,because of the silt layer). Similarly, a thick overpressured shale is often seen toabsorb quite major faulting. Where the salt or shale layer is thin, however, the fault

may be transmitted; then we face the judgment of whether it is likely to be sealedby flowage of the mobile material into the fault zone.

All in all, therefore, we cannot afford to dismiss fault traps, but working with themwill often intensify our problems-and occasionally cause us sorrow.

Interpretation, Past & Present

In the past, a great deal of seismic fault interpretation has been grossly in error.Some of the reasons for this are understandable.

With the poorer quality of old data, reflection continuity was often interrupted by seismic

artifacts-changes in surface conditions, refraction blind spots, incorrect statics. To aperson raised among mountains, every such break was a fault.

In the regrettable days when geologists and geophysicists were kept apart,many geophysicists did not even know what the geologists knew about thenature of faulting.

What explorationists know about basin dynamics and fault generation hasincreased enormously in the last decade.

There is always a temptation to give the seismic data the weight of ameasurement, while regarding the geological input as mere generalization.

Today, we resist this, and insist on an interpretation that satisfies both theseismic measurement and the geological generalizations. Perhaps the clearestexample of this is that we would no longer accept any fault interpretation,however "obvious," if it did not satisfy conservation of mass before and afterfaulting. One of the cheapest ways to find oil today is to reinterpret the oldfault maps over producing oil fields in the light of present geologicalknowledge, and to discover pockets of oil never tapped by the drill.

Round and Round

In modern practice, therefore, we are much concerned with establishing the tectonicregime of the area-and how that regime has changed over time-before we attempt

final picks on the faults. Here we put ourselves into an interpretive loop: we cannotpick the faults until we know the tectonic history, but we learn the tectonic history,very largely, from the faults. So no fault can be picked in isolation; we search firstfor faults that show the tectonic message most clearly, we make regional and localinferences about the history, and we go back into all the faults until a coherentpattern emerges.

2-D versus 3-D Seismic Data


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Fault interpretation on 2-D seismic data is much less safe than on 3-D data. Onereason for this is that 2-D lines are seldom sufficiently close to give definite faultcorrelations from line to line. Another is that fault interpretation is jeopardizedwhenever a 2-D line is not perpendicular to the fault. However, we must still beready to do the best we can on 2-D data.

Fault interpretation on 3-D data has given us new insight into the geometry,distribution and behavior of faulting. Once, where we mapped a few long continuousfaults using 2-D data, we now see more numerous short, discontinuous faultsseparated by some sort of transfer zone.

Similarly, mapping around faults has changed. The use of 3-D data has allowed us tosee the details of structure along and near the fault planes. We have learned thatstructural contours at the faults are not necessarily the smooth and flowing artisticlines that so many of us were taught in the past.

General Issues in Interpretation

Figure 1 illustrates the fault-picking operation at a simple level.

Figure 1

We join the apparent breaks or steps in the reflections with a geologically plausiblefault trace, continuing the pick upwards and downwards until the faulting becomesmere flexure, or disappears. Simple enough. But even in this simple case, we arelooking for additional messages. Above the fault we are always vigilant for evidence


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of reactivation (possibly in the opposite direction), because such reactivation mayhave broken a seal. However, we also remember the possibility that merecompaction, particularly in shales, can suggest minor reactivation of a fault wherenone has occurred. Further, the figure illustrates that either type of movement canproduce significant local changes in stratigraphy above a fault zone.

Figure 2 goes to an opposite extreme; it shows us that, even with fair record quality,fault interpretation can sometimes be extremely problematical.

Figure 2

In the face of the changing intervals between major reflections, the requirement for

geological plausibility becomes the key; whatever picks we make, we must have anexplanation for them. Just joining the reflection breaks in some arbitrary manner willnot do.

Even when the fault interpretation appears simple, we have to satisfy ourselves thatother faulting is compatible with that interpretation. Are the other faults ofcompatible type and the same age (acceptable), or of different type and differentage (acceptable), or of different type and the same age (problematical)?


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As we have agreed earlier, many of the problems of fault interpretation on 2-D dataarise because the lines are not at right angles to the faults. This problem is madeworse by the fact that many areas have been subjected to several tectonic episodes,each generating a system of subparallel faults-but at different orientations. Figure 3is a fairly simple illustration of this, with a system of curved extensional faultscrossed by a system of vertical strike-slip faults.

Figure 3

If we lay out a rectangular grid of 2-D lines over such an area, some or all of thelines must be at unfavorable angles.

Figure 4 illustrates fault picking at a basin edge.


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Figure 4

We can see very readily the likely geological development, and the picking isstraightforward. However, the earth is also capable of some exceedingly grotesquecontortions, and sometimes our courage is taxed ( Figure 5 ,


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Figure 5

Figure 6 , and Figure 7 ).


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Figure 7


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Figure 6

In some cases, situations that would otherwise tax our courage are eased when wecan recognize the presence of mobile rocks such as salt or overpressured shale.Thus, the major salt-induced fault in Figure 8 gives us few problems.


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Figure 8

An essential step in fault interpretation is the recognition of episodes of extension, ofneutral subsidence, and of compression. Thus, in Figure 9, we recognize an orderlysubsidence at the edge of a rift basin, and the repeated reactivation and adjustmentof the main basin-bounding fault.


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Figure 9

On the other side of the basin, Figure 10 shows the formation of a major left-tiltedfaulted block as a consequence of extension at rift time, before the filling of thebasin; such blocks, of course, may represent attractive petroleum targets.


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Figure 10

We would also read extension (followed by stable subsidence) into the tilted faultblocks ofFigure 11.


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Figure 11.

It is equally important that we recognize the evidence for compression. At a regionalscale, this may be fairly clear, and confirmed by the mountainous nature of thesurface. Occasionally we may even see such thrust situations at prospect scale (Figure 12 ); the associated rollover can yield a very important structural trap.


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Figure 12

A more complex example, involving both compression and extension, is given inFigure 13.


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Figure 13.

Many petroleum provinces have been subject to episodes of both extension andcompression. When these do not act in the same direction, some sidewaysmovements and adjustments occur, and these create distinctive fault patterns, bothin plan and in section. Figure 14 braces us for the difficult decisions that we shallhave to make, when we are working in such provinces.


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Figure 14

Stress

A rock is said to be in a state of stress when a force is applied to it. Earth stress isnot simple, but involves total force-per-unit-area for a particular point. What makesthe measurement of stress complex is the variety of magnitudes and directions offorce that occur within a single volume of rock. Basically, however, we can think ofstress as the three-dimensional intensity of force acting at a specified point.

At any moment in its total history, a rock has structure. It is also true that for any

such moment, it is subject to stresses that tend to alter its properties. In reality, onlyin deep space would this same rock be relatively free of stress. On earth, the forcesthat compose stress can be divided into two main typesbody forces, which act atevery point within the crust and surface forces (also called applied forces), which actonly at interfaces between objects and, therefore, are defined only along surfaces.

We can better understand how these forces act by considering a single clastic grainburied within the earth's crust. Two types of body forces act on this grain at alltimesgravity and inertia. We usually ignore the effects of inertia. The force due to


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gravity is called the lithostatic pressure. It results from the simple weight ofoverburden transferred by grain-to-grain contact. If our particular grain becomesinvolved in deformation, two types of surface forces will also exert their influence onit normal forces, which act perpendicular to the surface of the clast, and shearforces, which act parallel to these same surfaces.

Figure 1 (Components of stress on a single grain), shows diagrammatically thebreakdown of all these forces that compose the total stress on this single grain.

Figure 1

Since gravity and inertia, by definition, act on every particle at every point in the

crust, it is the surface forces that are primarily responsible for the creation ofgeologic structure and with which we as explorationists are concerned.

A rock undergoes deformation when stress causes the displacement of particleswithin it. This stress can result from body or surface forces, or both. In nature,deformation is almost never simple, since it results from a complex interactionbetween the chemical and physical properties of a rock mass, its immediateenvironment, and the rate and intensity at which the displacing forces are applied.


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We commonly make use of the concept of a stress field which refers to thedistribution of stress acting within a defined body. Such a body can be a single foldedlayer or a sizable portion of a continent. A stress field is described as homogeneous ifthe stress at each point is equal. The only situation that normally approximates thisnear the earth's surface is when stress is almost totally due to lithostatic pressure."Approximates" is the appropriate term to use here, since there is nearly always

some component, however small, of lateral stress in the crust.

As a measure of force-per-unit-area, stress is a vector quantity and may beexpressed as the sum of various components. If we trade our clastic grain for ahypothetical cube of uniform composition and subject this cube to a progressivedeformation ( Figure 2 , Resolution of unidirectional force (F) acting on a cube faceinto the basic normal (N) and shear (S) components of stress.

Figure 2

Also shown is the idealized physical effect of each component), we see that the totalforce (F) can be resolved into contributions that act perpendicular to the faces of thecube (N, or normal stress), and parallel to them (S, or shear stress). Normal stressescan be either compressional and tend to compact, or tensional and tend to separateparticles within a body. The effect of shear stresses, on the other hand, is to move


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particles past one another (as in the bottom right ofFigure 3 , Generalizedorientation of the principal axes of stress for three basic geologic structures.

Figure 3

In (a) and (b), horizontal compression can be thought of as the actual applied force,while (c) shows the effect of horizontal extension). During deformation, the relativemagnitudes of these two stress types will change, such that one type may decreaseas the other increases.

Principal Axes of Stress

Material behavior science, through its detailed analysis of stress conditions, hasderived an important conclusion that has proven very useful in structural geology:for any point in a homogeneous stress field, there exist three mutually orthogonalplanes along which all shear stresses vanish and only the components of normalstress exist. These three planes are known as the principal planes of stress, and theaxes of their intersection are thus the principal axes of stress. These axes are usedto describe what are referred to as the three principal stresses. This ideal triaxialsystem makes everything simpler, since it allows us to speak in terms of only normal


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stresses, i.e., compression, or "squeezing," and extension, or "pulling apart" ( Figure3 ).

In geology, principal stress is usually spoken of in terms of compression, which istaken as positive, and tension as negative. (For materials and engineering science,the opposite is true-i.e., tension is positive, compression is negative.) The three

principal axes, or stress directions, are correspondingly written as s1 (maximumprincipal stress), s2 (intermediate principal stress), and s3 (least, or minimum,principal stress). For our purposes, it is useful to understand four special states ofstress:

1. uniaxial stress, where two principal stresses are zero and the other is nonzero

2. biaxial stress, where two principal stresses are nonzero and the other iszero

3. triaxial stress, where all three principal stresses are nonzero

4. pure shear stress, where s1 equals s3 and is nonzero, while s2is zero. This is actually a special case of biaxial stress.

Within the earth's crust, the most common stress situation is triaxial, withs1 >s2>s3> 0.

In terms of actual physical effects, al can be thought of as representing the relativecompression that acts to deform a rock body by compacting particles into each other,while a represents the relative tension and the stretching of particles away from eachother. We say "relative" because these stresses are often defined only in relation toeach other. For example, a strong extensional event in the crust will have a alassociated with it that is secondary i.e., the compression is passive, related mainly tooverburden pressure in this case.

The relationship between the three principal axes of stress can be pictured byconsidering the triaxial examples shown in Figure 3 , which place our hypotheticalcube in evolving fold and fault structures. Again, s1 represents the direction ofmaximum effective compression; s3, extension. At this point, we should not beconcerned with the specific angle that the faults make with respect to thesedirections of stress. Note, however, that these faults and s2 are normal to the planeof the paper. In physical terms, s2 is parallel to the surface of the earth and basicallyrepresents the presence of essentially infinite neighboring material. Thus, with thesethree axes, the conditions of stress at any point or as geologists often apply it at acertain location within a defined plate tectonic regime, can be described and relatedto actual geologic structures.

Though there are no shear stresses acting on the surfaces of the cube in ourexample, the fact that s1, s2, and s3 differ in magnitude and direction means thatshear stress can be resolved within the cube. In fact, the quantity s1 -s3 (s1 minuss3), called the differential stress, is sometimes used as a general indicator ofshearing stress. If we were to measure the shearing stresses generated within ourcube, we would find that they reach a maximum along planes that are inclined at 45degrees to s1 and s3 and intersect in the plane ofs2 ( Figure 4 ) This represents theideal case, and (as we will see later on) it helps explain why rocks do not fracture


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randomly when stressed to the point of rupture.

Figure 4

Strain

Strain is said to exist when particles within a body have undergone displacement.Specifically, strain, usually denoted as 1, compares the change in length, DL, of anyline element in a rock body to its original length, L, which increases in somedirections and decreases in others. Strain is a measure (qualitative or numerical) of

this displacement, which, if permanent, is called deformation. Changes in shaperesulting from strain are called distortion; those in volume are described asdilatation, which can be either positive (in expansion) or negative (in contraction orshrinkage). Strain can result from both body and surface forces.

Although strain is commonly conceived of as the effect of stress, stress and strainare inseparable during actual deformation. Furthermore, if we make a closecomparison betweenFigure 1(Computer-generated stress field for a hypothetical


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fold in a less viscous matrix.

Figure 1

The short lines represents 3 axes) and Figure 2 (Computer-generated strain patternforFigure 1 .


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Figure 2

Short lines are drawn perpendicular to the axes of maximum shortening), we can seethat the specific relationship between stress direction and resulting strain is complex,even in simple situations. We can see, for example, that the orientations of the shortlines in these figures coincide least within the flanks of the folded layer, whereshearing stresses are greatest.

As we move up or down the flanks, toward the "hinges" of the fold, we notice thatthe attitude of the strain lines must be explained in terms of a growing combinationof normal and shearing stresses. Thus, we would need to understand precisely themechanisms involved in the transformation of stress into strain before we couldaccurately predict how these two vector quantities are geometrically related. As aresult, we often introduce helpful simplifications into the analysis of strain in nature.The more important of these are based on the assumption that the strain involved indeformation has been relatively homogeneous.

Much of the terminology derived for understanding stress has also been applied tostrain. For example, our hypothetical cube is said to have suffered homogeneousstrain, also called uniform strain, when the strain is the same at all points within it.This means that originally straight lines remain straight after deformation (Figure 3


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Homogeneous).

Figure 3

Thus, for example, a cube becomes a rhombohedron, while a sphere inscribed withinit becomes an ellipsoid. In fact, when we describe strain, we often do so withreference to an imaginary object called the strain ellipsoid. In this strain ellipsoid,the principle axes of strain are denoted as l1, l2 and l3 where the maximumlengthening is designated l1 and the maximum shortening as l3.

This should also help make clear the basic concept of inhomogeneous strain ( Figure3 , Inhomogeneous). Inhomogeneous strain is by far the most common in naturallydeformed rock. This type of strain involves some amount of rotation in the positionof particles, which means that originally straight lines become warped and detailedanalysis becomes impracticable. It is, therefore, almost always useful to find some

way in which natural deformation can be approximated as homogeneous. The mostcommon approach is to consider geologic structures as the summation of manylocalized homogeneous strain fields. This method has proved especially helpful in theexplanation of secondary rock fabrics, such as mineral alignment and fracturing.

Pure and Simple Shear

There are a number of basic ways in which deformation by homogeneous strain atconstant volume occurs. Those that involve simple flattening and stretching areshown in Figure 4 (Uniform extension),


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Figure 4

Figure 5 (Unifrom flattening),


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Figure 5

and Figure 6 (Plane strain: Each are basic types of homogeneous strain imposed on acube of ideally uniform composition.


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Figure 6

In each case, the inscribed circle and ellipse represent cross sections through thestrain wllipsoid before and after deformation). To understand how we treat naturaldeformation, however, it is also necessary that we look at the two basic types ofshear strain-pure shearand simple shear( Figure 7 , Hypothetical cross section anddiagram to illustrate domains of pure and simple shear in a series of folds that show

progressively greater strain.


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Figure 7

The shape of the strain ellipse can be the same for either type of shear and cannotbe used to derive detailed strain history). Both types help us explain a great manylarge and small scale features seen in rocks.

Pure shearis a form of strain in which no rotation of the strain axes takes place. It isalso often referred to as an irrotational deformation. It results from uniformextension in one direction and contraction perpendicular to it ( Figure 8 ).


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Figure 8

In simple shear, all particles within a body are displaced in one direction. This is ourcube pressed into a rhombohedron again; this time, however, we need to take noteof the rotation in the strain ellipsoid. Simple shear can be visualized by imagining theresult of placing our cube with its inscribed sphere between the two surfaces of anactive fault. The shearing motion created by these two surfaces stretches andflattens the sphere into a strain ellipsoid whose long axis is progressively rotateduntil it is nearly parallel to the fault plane itself. Displacement within such a bodytakes place by slippage along closely spaced planes ( Figure 8 ).

In actual materials, this can be accomplished in a number of ways for example, byslippage between grains or crystals, or by actual flow at elevated temperatures and

pressures. As we shall see, this style of deformation has widespread application togeologic structure.

Rock Strength

A substantial amount of literature exists with regard to experimental rock mechanicsin the laboratory, the intent of which is to simulate and analyze the process andeffects of deformation. Studies are usually performed on cylindrical, core-likesamples, which are subjected to compressive or tensile stresses in a chamber whose


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pressure and temperature are regulated. From these studies, scientists havedetermined that deformation generally progresses through three main stages (Nadai,1950). These stages are defined by the behavior of the deformed material, and aremost clearly and simply shown by the use of stress-strain diagrams ( Figure 1 ).

Figure 1

In order, the three stages are as follows:

1. ElasticDuring this initial stage, the stress-strain relationship is linear. If stress isremoved, the body reverts to its original dimensions. No deformation (permanent strain)results; strain is said to be completely reversible ( Figure 1 , segment A).

2. PlasticAs stress continues to increase, it will eventually reach some limitbeyond which the body suffers permanent strain. This is termed the elasticlimit, or yield stress. Beyond this limit, material is said to behave plastically:any increase in stress brings a corresponding increase in deformation ( Figure1 , segment B).

3. Rupture With continued increase of stress, the body will eventuallyfracture, rupture or fault ( Figure 1 , point C).


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Elastic strain can also end in rupture; a material simply fails without having undergone anymeasurable deformation.

Models of Deformation

These three stages through which deformation normally progresses are idealized as

the behavior of three hypothetical "bodies" that are subjected to stress, as shown inFigure 2 ,

Figure 2

Figure 3 , and Figure 4(Stress/strain relationships for several ideal materials: (a)Hookean (elastic) body; (b) St.


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Figure 3

Venant (plastic) body; (c) Newtonian (viscous) body).

Figure 4

Each style of behavior is denoted by the name of a well-known mathematician andshould be apparent from the stress-strain graphs shown. A Hookean body knowsonly elastic strain before rupture, and is approximated by a simple elastic springattached to a fixed body. A St. Venant body, in contrast, shows elastic strain up to ayield stress and then deforms indefinitely by shear strain at that same stress. Thistype of behavior is approximated by a weight that is pulled across some surface by


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an attached spring; the spring stretches elastically up to the point where friction onthe table top is overcome and the weight begins to slide. The last example, called aNewtonian body or, sometimes, Newtonian liquid, has no shear strength at all andtherefore exhibits no elastic strain. It deforms by what is called viscous strain. InFigure 4 , this type of body is represented as a porous piston pulled through a fluid-filled cylinder. A Newtonian body, then, will deform indefinitely in response to any

shear stress, with the total strain being directly proportional to the amount ofelapsed time.

Strain Hardening

These three types of ideal bodies help describe the components in progressivedeformation when it is caused by increasing amounts of stress. We might comparethem to the graphs shown in Figure 5 , which describe two generalized stress/strainrelationships for known materials.

Figure 5

The principal difference between the ideal


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( Figure 2 , Figure 3 , and Figure 4 ) and reality ( Figure 5 ) lies in the behaviorknown as strain hardening.

Such hardening is often the result of complex readjustment, or evenrecrystallization, of the material suffering strain. We can envision it on one level asbeing due to the compaction and realignment of particles during progressive

deformation. In sandstones, for example, stress concentrations along grainboundaries result in extensive grain fracturing and dissolution. This causes theprogressive filling of pore space with grain fragments and recrystallized quartz. Allother conditions (e.g., temperature and pressure) being equal, this will increase rockstrength-that is, greater and greater amounts of stress become necessary to imposethe same increment of deformation ( Figure 5 , upper curve). We expect certainclastic lithologies to respond in this manner.

On the other hand, strain hardening can occur up to some ultimate strength, afterwhich the stress necessary to cause a given strain decreases continually ( Figure 5 ,lower curve). Lithologies such as salt and gypsum-anhydrite may behave this wayunder certain conditions.

Brittleness and Ductility

On the basis of the differences between elastic and plastic behavior, we are able tocharacterize the general response of materials as either brittle or ductile.

Brittle materials rupture before any significant plastic deformation occurs. Suchbehavior in rocks is marked by the development of breakage discontinuities alongthe planes that represent maximum shear strain. These are not necessarily faults orfractures visible to the unaided eye, but may take place between individual grains orwithin crystal lattices. In contrast, material is described as ductile if it is able toundergo a large amount of plastic deformation before failing.

It should not be assumed that brittleness guarantees faulting, or that ductilityinevitably leads to folding. In every situation, whether a rock responds in a brittle orductile manner depends on several parameters-composition, effective confiningpressure, temperature, strain rate and anisotropy. It is, in fact, relativelymeaningless to speak of the true "strength" of a rock without reference to theseparameters.

Figure 6 ,


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Figure 6

Figure 7 ,


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Figure 7

and Figure 8 ,


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Figure 8

(A series of stress/strain diagrams showing the effects of confining pressure (Figure6 ), temperature (Figure 7 ), and strain rate (Figure 8 ), on rock deformation onhomogeneous (Solenhofen) limestone. Note for (Figure 8 ): a strain rate of 10-7persec is roughly equivalent to 1 year of time; most geologic deformation occurs atstrain rates of 10-14per sec) indicate how laboratory analysis has shown deformationto vary as a function of pressure (a), temperature (b), and strain rate (c) onhomogeneous limestone.

These laboratory tests tell us that near the surface, at low temperatures andpressures, rock will tend to act in a more brittle manner. With growing overburden,which increases both pressure and temperature, ductility generally increases. At thesame time, however, the element of time is crucial. If small amounts of stress areapplied over sufficiently long periods of time, almost any rock will deform plastically;only those lithologies with very low resistance to shear, such as basalt, may not. Butgiven normal rates of tectonic deformation, there is some transitional depth rangeover which the response of a particular rock type will grade from dominantly brittleto ductile.


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Figure 9 is a schematic representation of the spectrum from brittle to ductilebehavior in limestone, as determined by laboratory tests.

Figure 9

Such testing applies a uniaxial stress (either compression or tension) to a cylindricalsample that is sealed within a chamber whose temperature and confining pressurecan be regulated. Limestone and marble have been favored as samples, since theselithologies are more isotropic than clastic rock types.

In addition to pressure, temperature and strain rate, the other factor that determinesa rock's response to stress is, in one sense, the most obvious-its composition. Figure

10 is a diagram derived by Handin et al.


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Figure 10

(1963) from extensive lab experiments on natural rock. Shown are measuredductilities for four common sedimentary lithologies in a water-saturated condition.

One interesting point made by this diagram is that, with only one kilometer of burial,marked differences in ductility already exist between limestone and other lithologies.This is partly the result of pore pressure effects that strongly reduce ultimatestrength. It is also, however, directly related to the mineral structure of calcite,which allows for significant intracrystalline gliding, even at moderate pressures andtemperatures.

The same trend of ductility increase is true for sandstones, and is often aided by

pore pressure effects. We might also note that dolomite, while slightly more ductilethan quartzite at depths below two kilometers, shows the tightest range ofpermanent strain before rupture. Due to its mineral structure, dolomite does notdeform readily by intracrystalline gliding.

What is normally referred to as "rock strength," then, is not a fixed property, but arelative response, determined by a specific set of immediate environmentalconditions. Since both engineers and geoscientists make use of the termscompressive strength, tensile strength and shear strength, two major points should


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be made in relation to these qualities. First, almost all materials are weaker undertension than compression ( Table 1). And second, it is very often the shearingstrength of materials that determines when and how they will fail.

Average crushing

strength

Tensile

strength

Shearing

strengthSandstone740 10-30 100-200Limestone960 30-60 150-250Quartzite 2020 30-90 150-250Granite 1480 30-50 150-300Basalt 2500 - 50-150

Table 1: Measured strength of various rock types at standard temperature andpressure

Principal Stress Directions and Faulting

Figure 1 ,

Figure 1

Figure 2 ,


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Figure 2

and Figure 3 display the idealized orientation of the principal stress axes duringreverse (a), strike-slip (b) and normal (c) faulting. Because faulting and fracturing

both represent the brittle rupture of rocks, they are often discussed in relation tosimilar stress systems.


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Figure 3

Figure 4 shows the generalized orientation of actual fractures formed in aexperimental triaxial test, in which a block of Solenhofen limestone was shortened by1 percent at room temperature.


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Figure 4

The two categories of fractures that explorationists usually speak of, shear fracturesand extension fractures, were generated by this test and are indicated. This is howrock generally ruptures in the laboratory. Note that a broad correspondence existsbetween ideal faulting and shear fracturing with relation to stress orientation. As wehave previously noted, rocks are weakest in shear.

Two mutually orthogonal types of extensional ruptures are shown in Figure 4 . Thesehave been related to the stresses generated during loading and unloading, and, ineach case, they develop normal to s3. Notice that during unloading, s1 and s3exchange the orientation that they had during loading. This tells us that faults canconceivably develop at very low angles to s1, and that we should expect therelaxation of tectonic stress to generate late-stage features, especially fractures, in arock body.

More generally, it can be understood from Figure 1 , Figure 2 , and Figure 3 that aregion undergoing a high degree of faulting is characterized by numerous local stressfields that change in both orientation and magnitude as diastrophism progresses.The geometry and attitude of particular fault planes, therefore, may not alwaysappear directly related to the stress axes that apply to large-scale structural trends.The stress orientations shown in these figures do not, by themselves, always explain


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the specific angles at which faults develop or the shape of fault planes. For example,many fault planes are curved and cannot be explained simply in terms of one stressaxis alone we need several at varying depths.

At smaller and smaller levels of scale, many faults reveal increasingly complexcomponents of displacement, usually involving shear. couples. Figure 5 shows the

detailed resolution of shear fractures in an evolving monocline. The fractures shownare those which immediately precede propagation of a high-angle reverse fault frombasement into the overlying strata.

Figure 5

While in their basic geometry monoclines represent one of the simpler geologicstructures, the details of the deformation associated with them are not simple at all.

We should expect, then, that rocks in the vicinity of a fault or fault zone will becomplexly sheared for some distance on either side of the fault plane itself. Thisdistance may be measurable in centimeters to kilometers, depending on thelithologies, the type of fault, and the amount and environment of displacement.

General Terminology


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Much of the basic nomenclature relating to faults was derived from coal mining in theBritish Midlands during the late eighteenth and early nineteenth centuries. In fact,the word "fault" itself was originally used by miners to describe the sudden,unexpected, and vexatious termination of a coal seam. Thus, "fault" then carriedmuch of its vernacular sense-some sort of mistake had been made.

Early geologists like Murchison and Lyell, however, were quick to realize that a faultwas, in reality, a fracture where displacement had occurred, and that simplegeometric methods could be used to predict where a particular seam might be foundagain, either above or below. The bounding surface of a fault presented a "wall" tothe disgruntled miner, who was normally forced to continue his heading a short wayinto solid rock and then start a new shaft in order to relocate the seam. The wall ofthe fault plane was almost always inclined, which meant that the miner could hanghis lamp from the rocks on one side above the fault and rest his foot on those on theother side below it ( Figure 1 ).

Figure 1

Thus, the terms hangingwalland footwallthen, as now, simply label the two sides ofthe fault and imply nothing about displacement. Due to its more common occurrencein the coal-mining area of Britain, a fault was normalto the miner's experience when


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it was inclined toward the hanging wall. A coal seam that ended against such anormal fault could be found again if the miner continued his tunnel a short distancein the same direction and then sunk a shaft downward. At times, the reverse wastrue, and the corresponding fault was designated as a reverse fault. These termsremain in use today.

By modern definition, rocks are said to be faulted when they have sufferedobservable displacement along a plane or interval of rupture. As noted previously,such rupture occurs mainly by shear. The fault plane can be relatively simple ( Figure2 )

Figure 2

or it may consist of a large number of individual offset surfaces and thus be moreaccurately described as a fault zone ( Figure 3 ).


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Figure 3

(In the true mathematical sense, the term "plane" is incorrectly used, for rarely doesa fault exist without horizontal and vertical kinks, bends and changes of direction.Because of this, some of our colleagues protest the use of "fault plane" and insteadsupport the term fault surface. Although technically incorrect, fault plane is morecommonly used, so that is the term we will use in this text.)

Less frequently, rocks may be displaced by a form of shearing that causes loss ofinternal cohesion but not actual rupture of lithological layers ( Figure 4 ). We use theterm shear zone to label fault zones in which the individual planes of displacementare extremely closely spaced.


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Figure 4

The great majority of faults, however, more closely approximate planes of slip alongwhich shearing has taken place as a result of movement.

Faults are generally classified on the basis of their relative sense of displacement.Yet, as for most of geology, certain settings have encouraged the development anduse of more specific terminology, often related to proposed mechanisms. Each majortype of fault discussed immediately below is given a more specific nomenclaturewhen discussed in relation to a particular structural style.

Displacement varies along a fault, being greatest near the middle of the fault anddecreasing to zero at both ends. If we contour the amount of displacement, the

resulting map describes a displacement ellipse. We can construct displacementellipses in both plan view ( Figure 5 ) and along the fault trace.


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Figure 5

Net slip along a fault is measured by a vector that traces the displacement between

originally adjacent points ( Figure 6 , Basic terminology for fault offset, showingstrike-slip (ss) and dip-slip (ns) ).


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Figure 6

It is most often resolved into dip-slip (measured in the dip direction) and strike-slip(measured in the strike direction) components. Faults in which one or the other ofthese components is dominant are correspondingly named dip-slip faults and strike-slip faults.

For nonvertical dip-slip faults, we find it useful to again divide the displacement intovertical and horizontal components ( Figure 7 , Frequently used terminology forfaults in cross section: (a) definition of throw, heave, and hade.


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Figure 7

(Note: the latter two terms are more common in mining geology); and (b) definitionof normal stratigraphic separation). From coal mining terminology, the former issometimes called the throw, the latter, the heave, and the angle between the faultplane and the vertical, the hade. We frequently refer to the vertical and horizontalseparation across a dip-slip fault, while the inclination of the fault plane is simplysaid to be either high-angle or low-angle with reference to a horizontal datum. Also,we often use log data to determine the normal stratigraphic separation, as shown inFigure 7.

When "Throw" is Not Throw

The definitions noted in Figure 7 are fine when the beds are horizontal. However,when they are dipping, we must introduce an additional term, called verticalseparation, and distinguish it from throw.

Figure 8 shows a block diagram of a normal fault in which the formations are dippingfrom right to left.


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Figure 8

In this figure, we can see that when we project dipping beds across the fault plane,we can measure the displacement in two ways-we can project the bed across thefault perpendicularly to the vertical plane, or we can project the bed in the same dipangle across the fault. The former is the throw, the latter is defined as the verticalseparation.

There are only two cases where the vertical separation equals the fault throw. Thefirst is the case as shown in Figure 7 in which the stratigraphic units are horizontal-i.e., zero dip. The second is when the fault is vertical.

When we measure the fault "throw" using well logs, we are actually measuring the

vertical separation. However, when we measure the fault "throw" using a seismicsection, we are usually measuring the throw, although we are actually measuring theapparent throw.

The apparent throwis a measure of the fault displacement as a function of theangles between the direction of the seismic line and the strike of the fault as well asthe dip of the beds. Thus, if several seismic lines are oriented at different anglesrelative to the strike of a fault, each line would yield a different value of the faultdisplacement. The fault throw," as measured on a seismic line, is the true fault throw


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only if the direction of the seismic line is perpendicular to the fault strike and thebedding horizontal.

Ignoring this question of relative orientation, we can use seismic data to determinethe vertical separation (actually apparent vertical separation) to allow us to integrateseismic measurements with well log measurements. Figure 9 (Determination of

vertical separation by projecting dipping reflections across the fault plane), showsthat we can do this by projecting the interpreted horizons across the faults.

Figure 9

Back to Displacement

Figure 10(Major types of dip-slip faults.


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Figure 10

H and F refer to hanging wall and footwall) shows the four types of dip-slip faults. In

most cases, it is assumed that simple shear has acted as the principal strain withinthe fault plane ( Figure 11 , Diagram illustrating the progressive simple shear withinan ideal fault plane ).

Figure 11

We should note here that the fault type may not always be obvious. Whereseparation cannot be determined or appears to change along strike, determining thenature of the fault can be difficult. In some regions, a diversity of interrelatedfaulting styles exists in close juxtaposition; differentiating specific structures canprove very difficult.

In addition, many intracontinental basins are characterized by near-vertical, high-angle faults of small displacement that cannot be easily identified as normal orreverse. Furthermore, multiple episodes of tectonism often affect a single region.


57/349

These episodes may involve contrasting stress regimes, and displacement alongearly faults can be reversed. One of the more important occurrences of thisphenomenon involves provinces initially dominated by normal faulting beingsubjected to compressive stresses as a result of changing plate boundaryinteractions.

Normal Faults

Normal faults can either be planar or listric (concave upward). On seismic sections,planar faults often appear curved, due to the effect of increasing velocity with depth.In general, however, normal faults are more easily identified on seismic sectionsthan other types of faults. This is because of their frequent occurrence in deep,dominantly marine basins characterized by otherwise relatively undeformedsediments.

Normal faulting on a regional scale is most often referred to as block faulting, since itresults in the creation of horst (high) and graben (trench) structural topography (Figure 1 ).

Figure 1

Refer to Figure 2 ,


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Figure 2

Figure 3 , Figure 4


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Figure 4

, and Figure 5


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Figure 5

.


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Figure 3

The normal faults in these figures are clearly visible. In these seismic sections, it isrelatively easy to correlate the reflection horizons across the faults and distinguishthe hanging wall from the footwall, as well as the (apparent) throw in time. (Thethrow in depth may be a little more difficult.)

Thrust Faults

A thrust faulthas often been defined as a reverse fault dipping less than 45 degrees.However, today we use "thrust" as a generic term to imply near-horizontal,tangential compression and a zone of movement dominated by simple shear strain.

In Figure 1 and Figure 2 , we see seismic examples of thrust faulting.


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Figure 1

In Figure 1, the effects of compression are documented by the presence of olderrocks overlying younger ones.


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Figure 2

We noticed the "rollover" of beds caused by the thrusting in Figure 2.

Thrust faults can (and do) dip at any angle, and individual thrust planes often showcomplex, curving geometries. They may be essentially horizontal (along incompetentbedding planes), or flat, for kilometers, then "ramp" to a higher structural level incompetent rock and, finally, when reaching another incompetent bed, flatten outagain .

Thrust faults may curl up at their termini to become nearly vertical or evenoverturned. In the latter case, thrusts become apparent "normal faults,"geometrically speaking.

Often, due to complexity, it is very difficult or even impossible to distinguish betweenunderthrusting, in which the lower, relatively undeformed "block" is active, andoverthrusting, in which the upper, deformed block actually moves. We may onlyknow that shortening is involved, and this shortening can be as great as tens ofkilometers or more for single faults.

Rocks that have been transported from their original location (or root zone) are saidto be allochthonous ("other earth"), while those that remain in place are called


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autochthonous ("same earth"). The allochthon is also variously known as a thrustsheetorplate (not lithospheric), or, in certain circumstances, a nappe. Some thrustsdefine a listric plane that flattens with depth while others are of a ramp-and-flatgeometry. Several or more subsidiary thrusts commonly occur within a singleallochthon, either as splays off the sole fault (the deepest and controlling majorthrust fault in an area) or as local ruptures in the cores of anticlines. In general, the

thrusts in foreland belts are very rarely simple listric planes, but are themselvesfolded and, at times, truncated by younger thrusts.

As discussed by Dahlstrom (1970) and Elliot (1976), the geometry and location ofthrust faults in thick sedimentary sequences is largely determined by the distributionof competent and incompetent layers. Four simple rules summarize this influence:

1. Thrusting cuts up-section in the direction of displacement. (This is often called thedirection oftectonic transportor, in older nomenclature, the facingdirection.)

2. Thrusting tends to parallel bedding in incompetent layers, occurring nearcontacts with competent units, and to cut obliquely up-section in thicker,

more brittle units.

3. The age of major faults is younger in the direction of thrusting.

4. Major thrust faults do not overlap appreciably.

Evolution of Thrusting

The overall evolution, therefore, is for thrusting to begin at deeper levels and toprogress upward and outward (i.e., away from the root zone) from a majorsedimentary basin. Ductility contrast in the stratigraphic section encourages a stair-step evolution of thrusts, which we know as ramping. A relatively flat portion of a

thrust plane, particularly where it remains parallel to bedding within a singlelithology, is referred to as decollement. If this flattening of the thrust plane occurs atthe boundary between the sedimentary section and basement, it is called basaldecollement.

The progressive stacking of thrust faults may develop inpiggybackfashion, whereyounger thrusts form in the footwall or, alternatively, in overstep fashion, wherethrusts become younger toward the root zone. Piggybacking seems prevalent on aregional scale and is by far the more significant progression. At the same time, bothpiggyback and overstep thrusting occur on a more local level, often as a result ofimbrication.

Imbricates occur most often in two structural positions of high stress concentration-near the toe of a major thrust and above ramps in a thrust plane. They dip steeplyas they approach the surface, and stack slice after slice of the same stratigraphicsection along faults, which sole out into a major thrust plane. Continued movementalong this plane after the imbricates have formed will rotate them, so they canbecome vertical and overturned.

As discussed by Dahlstrom (1970), imbrication actually offers a basic model forforeland thrusting-as the scale of a cross section is increased to become more


65/349

regional, major thrusts themselves become imbricates of the largest faults (i.e.,those with the greatest displacement). These, in turn, can be thought of assubsidiary faults to a basal detachment or decollement plane that marks thestructural boundary between basement, usually crystalline metamorphic or plutonicrocks, and sedimentary cover.

Thin- versus Thick-Skinned Tectonics

The concept of basal decollement is the fundamental structural principle in thehypothesis known as thin-skinned tectonics. This theory is often contrasted with thethick-skinned hypothesis, which postulates no sole fault and, there-fore, directinvolvement of basement in each major thrust.

The debate between these two schools of thought is a historical one that continuestoday. On the basis of drilling and seismic data, both of which have proved theflattening of thrusts at depth in foreland areas, most explorationists now favor thethin-skinned hypothesis for at least the more medial and distal portions of thrustbelts. However, toward the metamorphic core of many such belts, basement rocks

are known to be heavily involved in thrusting. This involvement can be veryextensive, as in the case of the Himalayas, and may, in fact, control the overall styleand evolution of thrusting.

Yet some recent studies based on deep-reflection seismic profiles (Cook et al., 1979;Cook, 1982) have strongly favored the thin-skinned hypothesis for basementthrusting as well. Thus, the debate has expanded to focus on two major questionswhether thick-skinned faulting occurs at all in foreland belts, and, if it does, what isthe nature of the transition between it and the decollement tectonics thatcharacterize the sedimentary cover. The controversy represents one of the majorareas of research in contemporary structural geology.

Strike-Slip Faults

Strike-slip faults can be simply classified as left-lateral (also called sinistral) or right-lateral (also called dextral) on the basis of the displacement sense ( Figure 1 ).


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Figure 1

When facing the fault, the direction in which the opposite side appears to havemoved indicates this sense. In cross section, the symbols "A" and "T" are used tomark which side of the fault has been displaced away from and which side towardthe observer.

The San Andreas ( Figure 2 ) is one of the best examples of an active right-lateralstrike-slip fault.


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Figure 2

Others include the Alpine Fault of New Zealand and the Atacama system of Chile.Left-lateral faults of this scale are well-represented by the Philippine Fault ( Figure 3). These large strike-slip faults are often called wrench faults in general, andtranscurrent faults more specifically if they cut across regional structural trends.


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Figure 3

Transform faults are strike-slip faults that connect convergent and divergent plateboundaries. Basically, they serve to "transform" the interaction between plates intostrike-slip motion. The motion along these faults often includes components ofcompression or tension. Transform plate boundaries, therefore, are the links thatunify the world's spreading centers, subduction zones and collision zones into asingle mosaic of movement.

In Figure 4, Figure 5


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Figure 5

, and Figure 6 , we see the variety of structural types associated with large strike-slip faults.


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Figure 4

Later in this text, when we discuss strike-slip faults and their relationship to

structural styles, we will see several examples of these structures in seismic section.

Figure 6

Simple Fault Traps

More often than not, structural traps containing hydrocarbons result from or areassociated with some form of faulting. Although there are numerous, and sometimesunique, types of fault traps, most generally have an appearance similar to one of thesimple models shown in Figure 1 and


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Figure 2.


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Figure 1

General Terminology

Much of the basic nomenclature relating to faults was derived from coal mining in theBritish Midlands during the late eighteenth and early nineteenth centuries. In fact,the word "fault" itself was originally used by miners to describe the sudden,unexpected, and vexatious termination of a coal seam. Thus, "fault" then carriedmuch of its vernacular sense-some sort of mistake had been made.

Early geologists like Murchison and Lyell, however, were quick to realize that a fault

was, in reality, a fracture where displacement had occurred, and that simplegeometric methods could be used to predict where a particular seam might be foundagain, either above or below. The bounding surface of a fault presented a "wall" tothe disgruntled miner, who was normally forced to continue his heading a short wayinto solid rock and then start a new shaft in order to relocate the seam. The wall ofthe fault plane was almost always inclined, which meant that the miner could hanghis lamp from the rocks on one side above the fault and rest his foot on those on theother side below it ( Figure 1 ).


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Figure 1

Thus, the terms hangingwalland footwallthen, as now, simply label the two sides ofthe fault and imply nothing about displacement. Due to its more common occurrencein the coal-mining area of Britain, a fault was normalto the miner's experience whenit was inclined toward the hanging wall. A coal seam that ended against such anormal fault could be found again if the miner continued his tunnel a short distancein the same direction and then sunk a shaft downward. At times, the reverse wastrue, and the corresponding fault was designated as a reverse fault. These termsremain in use today.

By modern definition, rocks are said to be faulted when they have sufferedobservable displacement along a plane or interval of rupture. As noted previously,such rupture occurs mainly by shear. The fault plane can be relatively simple ( Figure2 )


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Figure 2

or it may consist of a large number of individual offset surfaces and thus be moreaccurately described as a fault zone ( Figure 3 ).


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Figure 3

(In the true mathematical sense, the term "plane" is incorrectly used, for rarely doesa fault exist without horizontal and vertical kinks, bends and changes of direction.Because of this, some of our colleagues protest the use of "fault plane" and insteadsupport the term fault surface. Although technically incorrect, fault plane is morecommonly used, so that is the term we will use in this text.)

Less frequently, rocks may be displaced by a form of shearing that causes loss ofinternal cohesion but not actual rupture of lithological layers ( Figure 4 ). We use theterm shear zone to label fault zones in which the individual planes of displacementare extremely closely spaced.


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Figure 4

The great majority of faults, however, more closely approximate planes of slip alongwhich shearing has taken place as a result of movement.

Faults are generally classified on the basis of their relative sense of displacement.Yet, as for most of geology, certain settings have encouraged the development anduse of more specific terminology, often related to proposed mechanisms. Each majortype of fault discussed immediately below is given a more specific nomenclaturewhen discussed in relation to a particular structural style.

Displacement varies along a fault, being greatest near the middle of the fault anddecreasing to zero at both ends. If we contour the amount of displacement, the

resulting map describes a displacement ellipse. We can construct displacementellipses in both plan view ( Figure 5 ) and along the fault trace.


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Figure 5

Net slip along a fault is measured by a vector that traces the displacement between

originally adjacent points ( Figure 6 , Basic terminology for fault offset, showingstrike-slip (ss) and dip-slip (ns) ).


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Figure 6

It is most often resolved into dip-slip (measured in the dip direction) and strike-slip(measured in the strike direction) components. Faults in which one or the other ofthese components is dominant are correspondingly named dip-slip faults and strike-slip faults.

For nonvertical dip-slip faults, we find it useful to again divide the displacement intovertical and horizontal components ( Figure 7 , Frequently used terminology forfaults in cross section: (a) definition of throw, heave, and hade.


79/349

Figure 7

(Note: the latter two terms are more common in mining geology); and (b) definitionof normal stratigraphic separation). From coal mining terminology, the former issometimes called the throw, the latter, the heave, and the angle between the faultplane and the vertical, the hade. We frequently refer to the vertical and horizontalseparation across a dip-slip fault, while the inclination of the fault plane is simplysaid to be either high-angle or low-angle with reference to a horizontal datum. Also,we often use log data to determine the normal stratigraphic separation, as shown inFigure 7.

When "Throw" is Not Throw

The definitions noted in Figure 7 are fine when the beds are horizontal. However,when they are dipping, we must introduce an additional term, called verticalseparation, and distinguish it from throw.

Figure 8 shows a block diagram of a normal fault in which the formations are dippingfrom right to left.


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Figure 8

In this figure, we can see that when we project dipping beds across the fault plane,we can measure the displacement in two ways-we can project the bed across thefault perpendicularly to the vertical plane, or we can project the bed in the same dipangle across the fault. The former is the throw, the latter is defined as the verticalseparation.

There are only two cases where the vertical separation equals the fault throw. Thefirst is the case as shown in Figure 7 in which the stratigraphic units are horizontal-i.e., zero dip. The second is when the fault is vertical.

When we measure the fault "throw" using well logs, we are actually measuring the

vertical separation. However, when we measure the fault "throw" using a seismicsection, we are usually measuring the throw, although we are actually measuring theapparent throw.

The apparent throwis a measure of the fault displacement as a function of theangles between the direction of the seismic line and the strike of the fault as well asthe dip of the beds. Thus, if several seismic lines are oriented at different anglesrelative to the strike of a fault, each line would yield a different value of the faultdisplacement. The fault throw," as measured on a seismic line, is the true fault throw


81/349

only if the direction of the seismic line is perpendicular to the fault strike and thebedding horizontal.

Ignoring this question of relative orientation, we can use seismic data to determinethe vertical separation (actually apparent vertical separation) to allow us to integrateseismic measurements with well log measurements. Figure 9 (Determination of

vertical separation by projecting dipping reflections across the fault plane), showsthat we can do this by projecting the interpreted horizons across the faults.

Figure 9

Back to Displacement

Figure 10(Major types of dip-slip faults.


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Figure 10

H and F refer to hanging wall and footwall) shows the four types of dip-slip faults. In

most cases, it is assumed that simple shear has acted as the principal strain withinthe fault plane ( Figure 11 , Diagram illustrating the progressive simple shear withinan ideal fault plane ).

Figure 11

We should note here that the fault type may not always be obvious. Whereseparation cannot be determined or appears to change along strike, determining thenature of the fault can be difficult. In some regions, a diversity of interrelatedfaulting styles exists in close juxtaposition; differentiating specific structures canprove very difficult.

In addition, many intracontinental basins are characterized by near-vertical, high-angle faults of small displacement that cannot be easily identified as normal orreverse. Furthermore, multiple episodes of tectonism often affect a single region.


83/349

These episodes may involve contrasting stress regimes, and displacement alongearly faults can be reversed. One of the more important occurrences of thisphenomenon involves provinces initially dominated by normal faulting beingsubjected to compressive stresses as a result of changing plate boundaryinteractions.

Normal Faults

Normal faults can either be planar or listric (concave upward). On seismic sections,planar faults often appear curved, due to the effect of increasing velocity with depth.In general, however, normal faults are more easily identified on seismic sectionsthan other types of faults. This is because of their frequent occurrence in deep,dominantly marine basins characterized by otherwise relatively undeformedsediments.

Normal faulting on a regional scale is most often referred to as block faulting, since itresults in the creation of horst (high) and graben (trench) structural topography (Figure 1 ).

Figure 1

Refer to Figure 2 ,


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Figure 2

Figure 3 , Figure 4


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Figure 4

, and Figure 5


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Figure 5

.


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Figure 3

The normal faults in these figures are clearly visible. In these seismic sections, it isrelatively easy to correlate the reflection horizons across the faults and distinguishthe hanging wall from the footwall, as well as the (apparent) throw in time. (Thethrow in depth may be a little more difficult.)

Thrust Faults

A thrust faulthas often been defined as a reverse fault dipping less than 45 degrees.However, today we use "thrust" as a generic term to imply near-horizontal,tangential compression and a zone of movement dominated by simple shear strain.

In Figure 1 and Figure 2 , we see seismic examples of thrust faulting.


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Figure 1

In Figure 1, the effects of compression are documented by the presence of olderrocks overlying younger ones.


89/349

Figure 2

We noticed the "rollover" of beds caused by the thrusting in Figure 2.

Thrust faults can (and do) dip at any angle, and individual thrust planes often showcomplex, curving geometries. They may be essentially horizontal (along incompetentbedding planes), or flat, for kilometers, then "ramp" to a higher structural level incompetent rock and, finally, when reaching another incompetent bed, flatten outagain .

Thrust faults may curl up at their termini to become nearly vertical or evenoverturned. In the latter case, thrusts become apparent "normal faults,"geometrically speaking.

Often, due to complexity, it is very difficult or even impossible to distinguish betweenunderthrusting, in which the lower, relatively undeformed "block" is active, andoverthrusting, in which the upper, deformed block actually moves. We may onlyknow that shortening is involved, and this shortening can be as great as tens ofkilometers or more for single faults.

Rocks that have been transported from their original location (or root zone) are saidto be allochthonous ("other earth"), while those that remain in place are called


90/349

autochthonous ("same earth"). The allochthon is also variously known as a thrustsheetorplate (not lithospheric), or, in certain circumstances, a nappe. Some thrustsdefine a listric plane that flattens with depth while others are of a ramp-and-flatgeometry. Several or more subsidiary thrusts commonly occur within a singleallochthon, either as splays off the sole fault (the deepest and controlling majorthrust fault in an area) or as local ruptures in the cores of anticlines. In general, the

thrusts in foreland belts are very rarely simple listric planes, but are themselvesfolded and, at times, truncated by younger thrusts.

As discussed by Dahlstrom (1970) and Elliot (1976), the geometry and location ofthrust faults in thick sedimentary sequences is largely determined by the distributionof competent and incompetent layers. Four simple rules summarize this influence:

1. Thrusting cuts up-section in the direction of displacement. (This is often called thedirection oftectonic transportor, in older nomenclature, the facingdirection.)

2. Thrusting tends to parallel bedding in incompetent layers, occurring nearcontacts with competent units, and to cut obliquely up-section in thicker,

more brittle units.

3. The age of major faults is younger in the direction of thrusting.

4. Major thrust faults do not overlap appreciably.

Evolution of Thrusting

The overall evolution, therefore, is for thrusting to begin at deeper levels and toprogress upward and outward (i.e., away from the root zone) from a majorsedimentary basin. Ductility contrast in the stratigraphic section encourages a stair-step evolution of thrusts, which we know as ramping. A relatively flat portion of a

thrust plane, particularly where it remains parallel to bedding within a singlelithology, is referred to as decollement. If this flattening of the thrust plane occurs atthe boundary between the sedimentary section and basement, it is called basaldecollement.

The progressive stacking of thrust faults may develop inpiggybackfashion, whereyounger thrusts form in the footwall or, alternatively, in overstep fashion, wherethrusts become younger toward the root zone. Piggybacking seems prevalent on aregional scale and is by far the more significant progression. At the same time, bothpiggyback and overstep thrusting occur on a more local level, often as a result ofimbrication.

Imbricates occur most often in two structural positions of high stress concentration-near the toe of a major thrust and above ramps in a thrust plane. They dip steeplyas they approach the surface, and stack slice after slice of the same stratigraphicsection along faults, which sole out into a major thrust plane. Continued movementalong this plane after the imbricates have formed will rotate them, so they canbecome vertical and overturned.

As discussed by Dahlstrom (1970), imbrication actually offers a basic model forforeland thrusting-as the scale of a cross section is increased to become more


91/349

regional, major thrusts themselves become imbricates of the largest faults (i.e.,those with the greatest displacement). These, in turn, can be thought of assubsidiary faults to a basal detachment or decollement plane that marks thestructural boundary between basement, usually crystalline metamorphic or plutonicrocks, and sedimentary cover.

Thin- versus Thick-Skinned Tectonics

The concept of basal decollement is the fundamental structural principle in thehypothesis known as thin-skinned tectonics. This theory is often contrasted with thethick-skinned hypothesis, which postulates no sole fault and, there-fore, directinvolvement of basement in each major thrust.

The debate between these two schools of thought is a historical one that continuestoday. On the basis of drilling and seismic data, both of which have proved theflattening of thrusts at depth in foreland areas, most explorationists now favor thethin-skinned hypothesis for at least the more medial and distal portions of thrustbelts. However, toward the metamorphic core of many such belts, basement rocks

are known to be heavily involved in thrusting. This involvement can be veryextensive, as in the case of the Himalayas, and may, in fact, control the overall styleand evolution of thrusting.

Yet some recent studies based on deep-reflection seismic profiles (Cook et al., 1979;Cook, 1982) have strongly favored the thin-skinned hypothesis for basementthrusting as well. Thus, the debate has expanded to focus on two major questionswhether thick-skinned faulting occurs at all in foreland belts, and, if it does, what isthe nature of the transition between it and the decollement tectonics thatcharacterize the sedimentary cover. The controversy represents one of the majorareas of research in contemporary structural geology.

Strike-Slip Faults

Strike-slip faults can be simply classified as left-lateral (also called sinistral) or right-lateral (also called dextral) on the basis of the displacement sense ( Figure 1 ).


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Figure 1

When facing the fault, the direction in which the opposite side appears to havemoved indicates this sense. In cross section, the symbols "A" and "T" are used tomark which side of the fault has been displaced away from and which side towardthe observer.

The San Andreas ( Figure 2 ) is one of the best examples of an active right-lateralstrike-slip fault.


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Figure 2

Others include the Alpine Fault of New Zealand and the Atacama system of Chile.Left-lateral faults of this scale are well-represented by the Philippine Fault ( Figure 3). These large strike-slip faults are often called wrench faults in general, andtranscurrent faults more specifically if they cut across regional structural trends.


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Figure 3

Transform faults are strike-slip faults that connect convergent and divergent plateboundaries. Basically, they serve to "transform" the interaction between plates intostrike-slip motion. The motion along these faults often includes components ofcompression or tension. Transform plate boundaries, therefore, are the links thatunify the world's spreading centers, subduction zones and collision zones into asingle mosaic of movement.

In Figure 4, Figure 5


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Figure 5

, and Figure 6 , we see the variety of structural types associated with large strike-slip faults.


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Figure 4

Later in this text, when we discuss strike-slip faults and their relationship to

structural styles, we will see several examples of these structures in seismic section.

Figure 6

Simple Fault Traps

More often than not, structural traps containing hydrocarbons result from or areassociated with some form of faulting. Although there are numerous, and sometimesunique, types of fault traps, most generally have an appearance similar to one of thesimple models shown in Figure 1 and


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Figure 2.


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Figure 1

Faults and Diffractions We need only Huygens' principle to tell us to expect a diffraction from a fault. In Figure1 (Looking at diffractions in terms of Huygens sources), at point A on a continuousreflector, one Huygens source interacts with those each side of it to leave only the

reflected signal at normal incidence.


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Figure 1

At B, the presence of the last Huygens source at the edge tells us that a signal must beobserved in the direction C, while the absence of further Huygens sources to the righttells us that the "backward" radiation in direction D will not be "canceled."

Exactly the same message is conveyed, though in a different form, byconsidering the circular reflection zone on the reflector ( Figure 2, Looking atdiffractions in terms of the relfection zone).


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Figure 2

The reflected signal at surface point P includes contributions from all of the

reflection zone R. That at point Q includes contributions from only asemicircular zone S, so that we would expect only half the amplitude. Thediffracted signal at point T includes contributions from a zone that graduallydecreases in size as the offset increases.

The classical seismic expression of a fault is obtained by summing thesecontributions. For the assumed model of a terminating reflector, the result isas shown in Figure 3 (Seismic response to a terminating reflactor)-a planereflection that decreases in amplitude to one-half at the fault, and ahyperbolic diffraction with its apex at the fault.


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Figure 3

The diffraction comprises a forward branch of the same polarity as thereflector and a backward branch of the opposite polarity.

The physical necessity for the two branches of opposite polarity is easilyshown by bringing together two such models, as in Figure 4.

Figure 4.


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In the limit, when the reflector becomes continuous, the diffractions mustcancel to zero.

The diffraction amplitudes make sense. Since the reflection melds smoothly

into the forward branch, and since the amplitude of the reflection at the faultis one-half that for a continuous reflection, the amplitude of the forwardbranch near the fault is one-half, and therefore so is that of the backwardbranch. In this sense, the half-amplitude of the reflection at the fault may beviewed as the whole amplitude minus the backward branch. Therefore,canceling the backward branch with the forward branch (by summing alongthe diffraction hyperbola, as in migration) not only removes the entirediffraction but restores the reflection amplitude at the fault.

The above model of a terminating reflector is geologically unreal; a reflectorcannot just terminate. There must be an additional contribution from the faultface. However, for a vertical fault, it is still true that there must be forward

and backward branches of equal amplitude and opposite polarity, to providecancellation when two blocks are brought together in the manner ofFigure 4 ,(The physical necessity for backward and forward branches, as the gapbetween reflector terminations decreases) .

One situation that is close to the classical model is that of a vertically faultedthin horizontal layer ( Figure 5, The differentiated form of the classicalresponse, obtained from a thin layer (here assumed to be hard) ), where thesignal from the fault face is very small.


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Figure 5

We remember that if the layer is thin, relative to the seismic wavelength, the

reflection response approximates to a differentiated version of the incidentpulse; then the forward and backward diffractions also approximate todifferentiated versions of the classical response.

The response changes if the fault plane is inclined. Although the diffractionsremain grossly hyperbolic, the pulse shape and the amplitude of the forwardand backward branches may no longer be the same. However, as suggestedin Figure 6, (If the fault is inclined the details of the diffractions change,

Figure 6

but bringing the fault blocks together (a) still requires cancellation of theforward branch of one by the backward of the other (b) ), it is still necessarythat the forward branch from the left block should cancel the backwardbranch from the right block (and the reverse) when they are broughttogether; therefore, these must have the same amplitude and form, butopposite polarity. The reasonableness of this is confirmed when we rememberthat the tails of the diffractions contain the fault plane reflections, which mustalso cancel.


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If we complete the fault, as in Figure 7 (The reflection from the inclined faultplane in (a) is found among the tails of the forward branch of the upblock and

the backward branch of the downblock (b).

Figure 7

The impression is therefore of stronger diffractions on the downthrown side),there is ordinarily a strong forward branch from the upthrown block and astrong backward branch from the downthrown block, and the lateral offsetmeans that these must approach each other. Again, these tails contain thefault-plane reflection.

The backward branch from the upthrown block and forward branch from thedownthrown block ordinarily remain well separated, as in the figure. However,if there is a rapid increase of velocity with depth, the lower diffraction isflatter, and the two could again approach each other. In this case the zone ofconfluence would contain the reflection from the underside of the faultplaneobtainable because the rapid increase of velocity with depth turns thedowngoing rays upward again.


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The above observations apply to sharp rupture of the reflector at the fault...as occurs with brittle rocks. With more plastic rocks, there is likely to be draginto the fault This may be normal or reverse, and may be modified byreactivation or by isostatic rebound (footwall uplift). For any type and degreeof drag, we can synthesize the response using Huygens hyperbolas.

For normal drag into an inclined fault, as suggested in Figure 8 (Likelyappearance of a normal fault with normal drag),

Figure 8

a common effect is that only the forward branch from the upthrown block and

the backward branch from the downthrown block remain visible ( Figure 8 ).

Similarly, the dominant diffraction from the layers pierced by a salt plug isoften the backward branch of reversed polarity ( Figure 9, The termination ofa stratal reflector (a) at a salt plug often yeilds (b) an amplitude increase

caused by the concave focusing,


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Figure 9

and a diffraction dominated by the backward (whose tail contains thereflection from the salt-positive or negative, as the case may be).

The preceding figures illustrate the diffractions obtained when the faultedlayers are horizontal. If the faulted layers dip, the diffraction hyperbolas aretangent to the reflection at the point of apparent faulting, and have theirmaximum amplitude at this point; then the true position of the fault isthrough the apices of the diffraction hyperbolas ( Figure 10, The situationwhere the faulted layers dip, and the distinction between apparent,

unmigrated and migrated positions of the fault plane ).


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Figure 10

The figure also shows that the apparent position of the fault plane, joiningthe terminations of the stratal reflections, does notcorrespond to theunmigrated position of the fault-plane reflection, which is contained within thediffraction tails. The fault pick that we make on the unmigrated section-fromdiffraction apex to diffraction apex-is the migratedposition of the fault.

On an unmigrated section, an inclined fault-plane reflection must alwaysappear to cut across the reflections it truncates; Figure 11 (The fault planecorresponds (approximately) to the apices of the diffractions. The fault-planereflection is contained within the tails of the diffractions, and therefore must

cut across stratal reflections) is a case in point.


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Figure 11

A full hyperbola without any change of polarity at the apex ( Figure 12, Themessage of a full hyperbola with no change of polarity is a tight syncline-notin itself a fault) generally indicates a buried focus (often, a tight syncline)-nota fault (though it may also signal a shallow source off the line).


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Figure 12

In principle, diffractions can help us to distinguish between normal andreverse faults, in that the diffraction seen most clearly has the same polarityas the faulted reflection for a normal fault, and the opposite polarity for areverse fault ( Figure 13 , Signature of an inclined normal fault-planereflection (for a positive stratal reflection).


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Figure 13

(b) The corresponding reverse fault, with its different-polarity diffract

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