Did you ever take a cross-country drive? Hour after hour of tedious driving, as the highway climbed hills and dropped into valleys? The monotonous gray rocks exposed in road cuts largely went unnoticed, right? You passed pretty scenery, but it was static and seemed to tell no story simply because you did not have a basis in your mind with which to interpret your natural surroundings. It was much the same for scholars of generations past, before the establishment of modern science. The Earth was a closed book, hiding its secrets in a language that no one could translate. Certainly, ancient observers marveled at the enormity of mountains and oceans, but with the knowledge they had at hand they could do little more than dream of supernatural processes to explain the origin of these features. Gods and monsters contorted the Earth and spit flaming rock; and giant turtles and catfish shook the ground. Then, in fifteenth-century Europe, an intellectual renaissance spawned an age of discovery, during which the Earth was systematically charted, and the pioneers of science cast aside dogmatic views of our universe that had closed peoples’ minds for the previous millennia and began to systematically observe their surroundings and carry out experiments to create new knowledge. The scientific method was born.

FIGURE 1.1. Sketch by Leonardo da Vinci showing details of folded strata in the mountains of Italy (ca. 1500 AD). In recent years it was discovered that in addition to his careful observations of the natural world, da Vinci also completed insightful friction experiments.

In geology, the stirrings of discovery are evident in the ink sketches of the great artist and inventor Leonardo da Vinci (1452–1519), who carefully drew the true shapes of rock bodies in sketches to understand the natural shape of the Earth (Figure 1.1). In the seventeenth century came the first description of rock deformation. Nicholas Steno (1631–1686) examined outcrops where the bedding of rock was not horizontal, and speculated that strata that do not presently lie in horizontal layers must have in some way been dislocated (the term he used for deformed). Perhaps Steno’s establishment of the principle of original horizontality can be viewed as the birth of structural geology. By the beginning of the eighteenth century, the structural complexity of rocks in mountain ranges like the Alps was widely recognized (Figure 1.2), and it became clear that such features demanded explanation.

Bernese Alps, Switzerland

FIGURE 1.2. Aerial view of the central European Alps (Switzerland) looking approximately north. Some of the higher peaks include Jungfrau (4,158m) and Eiger (3,970m). The flowing ice of Aletsch glacier is marked by dark moraines. [NASA]

The pace of discovery quickened during the latter half of the eighteenth century and through the nineteenth century. In his “Theory of the Earth with Proofs and Illustrations,” James Hutton (1726–1797) proposed the concept of uniformitarianism and provided an explanation for the nature of unconformities. Since the publication of this book in 1785 there has been a group of scientists who recognize themselves as geologists. These new geologists defined the geometry of structures in mountain ranges, learned how to make geologic maps, discovered the processes involved in the formation of rocks, and speculated on the origins of specific structures and on mountain ranges in general.

FIGURE 1.3. Model of mountain building and associated deformation as represented by G. P. Scrope (1825). The uplift is caused by intrusion of an igneous core, and the folds are generated by down-slope movement.

Ideas about the origin of mountains have evolved gradually. At first, mountain ranges were thought to be a consequence of a vertical push from below, perhaps associated with intrusion of molten rock along preexisting zones of weakness, and folds and faults in strata were attributed to gravity sliding down the flanks of these uplifts (Figure 1.3). Subsequently, the significance of horizontal forces was emphasized, and geologists speculated that mountain ranges and their component structures reflected the contraction of the Earth that resulted from progressive cooling. In this model, the shrinking of the Earth led to wrinkling of the surface. One of the more notable discoveries (about 1850) was the recognition by James Hall (1811–1898) that Paleozoic strata in the Appalachian Mountains of North America were much thicker than correlative strata in the interior of the continent. This discovery led to the development of the geosyncline theory, a model in which deep sedimentary basins, called geosynclines, evolved into mountain ranges. Contraction theory and geosynclinal theory, or various combinations of the two, were widely accepted until the 1960s, when the views of Alfred Wegener (1880–1930), Arthur Holmes (1898–1965), and Harry Hess (1906–1969) led to the formulation of a very different model. Building on the work of Alfred Wegener’s continental drift theory and Arthur Holmes’s mantle convection model, Harry Hess proposed the revolutionary idea of a mobile seafloor (seafloor spreading hypothesis) that lead to the formulation of plate tectonic theory. In this theory, the Earth consists of several, rigid plates that change in space and time. The interaction between these plates offers a unifying explanation for the occurrence of mountain ranges, ocean basins, earthquakes, volcanoes, and other previously disparate geologic phenomena.

As the foundations of geology grew, diverse features of rocks and mountains gained names, and the once amorphous, nondescript masses of rock exposed on our planet became history books that preserve the Earth’s biography. Perhaps your concept of the planet has evolved rapidly as well, because of the courses in geology and other sciences that you have taken thus far. Now, as you drive across the countryside, you scare the daylights out of your passengers as you twist to see and discuss roadside outcrops. The rocks are no longer gray masses to you, but they contain recognizable patterns and shapes and fabrics. The purpose of this book is to increase your ability to interpret these features, and particularly to use them as clues to understanding the processes that have shaped and continue to change the outer layers of the Earth.

1.2 CLASSIFICATION OF GEOLOGIC STRUCTURES

When you finished your introductory geology course, you probably had a general concept of what a geologic structure is. The term probably brings to mind images of folds and faults. Perhaps you had the opportunity to take a field trip where you saw some of these structures in the wild. These features are formed in response to pushes and pulls associated with the forces that arise from the movement of tectonic plates or as a consequence of differential buoyancy between parts of the lithosphere. But what about bedding in a sedimentary rock and flow banding in a rhyolite flow; are these structures? And what about slump folds in a debris flow; are they structures? Well . . . yes, but the link between their formation and deformation is less obvious. So, maybe we need to have a more general concept of a geologic structure.

The most fundamental definition of a geologic structure is a geometric feature in rock whose shape, form and distribution can be described. From this definition it is obvious that there are several ways in which geologic structures can be subdivided into groups. In other words, by necessity there are several different, yet equally valid classification schemes that can be used in organizing the description of geologic structures. Different schemes are relevant for different purposes, so we will briefly look at various classification schemes for geologic structures that will return in subsequent chapters. At first, these various classification schemes may seem very confusing. Thus, we recommend that you start by recognizing the basic geometric classes as the foundation of your understanding.

As you learn about these classes, refer back to the lists below, and see how a particular geometric class fits into several classification schemes.

I. Classification based on geometry, that is, on the shape and form of a particular structure:

• Planar (or subplanar) surface.

• Curviplanar surface.

• Linear feature.

This subdivision represents perhaps the most basic classification scheme. In this scheme we include the following classes of structures: joint, vein, fault, fold, shear zone, foliation, and lineation.

II. Classification based on timing of formation:

• Syn-formational - formed at the same time as the material that will ultimately form the rock.

• Penecontemporaneous - formed before full lithification, but after initial deposition.

• Post-formational - formed after the rock has fully formed, as a consequence of phenomena not related to the immediate environment of rock formation.

III. Classification based on geologic significance:

• Primary - formed as a consequence of the formation process of the rock itself.

• Local gravity-driven - formed due to slip down an inclined surface; slumping at any scale driven by local excess gravitational potential.

• Local density-driven - formed due to local lateral variations in rock density, causing a local buoyancy force.

• Fluid-pressure driven - formed by injection of unconsolidated, fluidized material due to sudden release of pressure.

• Igneous/Volcanic – formed by the intrusion and extrusion or moving magma and volcanic flows.

• Impact – formed by the application of large forces from the impact of falling objects on Earth's surface.

• Tectonic - formed by lithospheric plate interactions, due to regional interactions between the asthenosphere and the lithosphere, due to crustal- or lithosphere-scale gravitational potential energy and the tendency of lithosphere to achieve isostatic compensation.

All but the last can also be grouped as non-tectonic structures, meaning that they are not directly related to the forces associated with tectonics. We purposely say “can” because in many circumstances these categories of structures do form in association with tectonic activity. For example, gravity sliding may be triggered by tectonically-generated seismicity, and salt domes may be localized by movement of tectonic normal faults. We discuss these categories here, while the last category, deformation structures, forms the focus of this book.

1.3 SEDIMENTARY STRUCTURES

When you look at an outcrop of sedimentary rock, the most obvious fabric to catch your eye is the primary layering or stratification (Figure 1.4), which is called bedding.

FIGURE 1.4. Differential erosion of bedding surfaces in the Claron (formerly Wasatch) Formation, Bryce Canyon (Utah). [Wikimedia]

What defines bedding in an outcrop? In the Painted Desert of Arizona, beds are defined by spectacular variations in colors, and outcrops display garish stripes of maroon, red, green, and white. In the Grand Canyon and the Rocky Mountains, beds are emphasized by contrasts in resistance to erosion; sandstone and limestone beds form vertical cliff faces while shale layers form shallow slopes. On the cliffs that form the east edge of the Catskill Mountains in New York State, there are abrupt contrasts in grain size between adjacent beds, with a coarse conglomerate juxtaposed against siltstone or shale. Strictly speaking, a bed is the smallest subdivision of a sedimentary unit. Stratigraphers divide sequences of strata into the following units: supergroup, group, formation, member, bed. Several beds make a member, several members make a formation, and so on. Criteria for defining units are somewhat subjective. Basically, a unit is defined as a sequence of strata that can be identified and mapped at the surface or in the subsurface over substantial region. The basis for recognizing a unit can be its age, its component sequence of lithologies, and the character of its bedding. A bed has a definable top or bottom and can be distinguished from adjacent beds by differences in grain size, composition, color, sorting, and/or by a physical parting surface. All of the features defining bedding, with the exception

of parting, are a consequence of changes in the source of the sediment (or provenance) or the depositional environment.

In some outcrops, bedding is enhanced by the occurrence of bedding-parallel parting. Parting forms when beds are unroofed (i.e., overlying strata are eroded away) and uplifted to shallower depths in the crust. As a consequence, the load pushing down on the strata decreases and the strata expand slightly. During this expansion, fractures form along weak bedding planes and define the parting. This fracturing reflects the weaker bonds between contrasting lithologies of adjacent beds, or the occurrence of a preferred orientation of sedimentary grains (e.g., mica). If a sedimentary rock has a tendency to have closely spaced partings, it is said to display fissility. Shale, which typically has a weak bedding-parallel fabric due to the alignment of constituent clay or mica flakes, is commonly fissile.

There are three reasons why platy grains like mica have a preferred orientation in a sedimentary rock. First, the alignment of grains can reflect settling of asymmetric bodies in Earth’s gravity field. Platy grains tend to lie down flat. To understand why, throw a handful of coins or cards into the air: after the coins fall they lie flat against the floor. You will rarely see a coin stand on its edge. Second, the alignment of grains can reflect flow of the fluid in which the grains were deposited. In a moving fluid, grains are reoriented so that they are hydrodynamically stable, meaning that the traction caused by the moving fluid is minimized (as is the case if the broad face of the grain parallels the flow direction). Typically, grains end up being imbricated, meaning that they overlap one another like roof shingles. Imbrication is a useful primary sedimentary feature that can be used to define paleocurrent direction, which is the direction of the current when the sediment was deposited. Third, the alignment of grains can form as a consequence of compaction subsequent to deposition. As younger sediment is piled on top, water is progressively squeezed out of the older sediment below and the grains mechanically rotate into an orientation with their flat surfaces roughly perpendicular to the applied load.

Bedding in outcrops is often highlighted by differential weathering and erosion (Figure 2.1). For example, chemical weathering (like dissolution of carbonate) of a sequence of strata containing alternating limestones and quartz sandstones will result in an outcrop face on which the quartz-rich layers stand out in relief. Fresh limestone and dolomite are almost identical in color, but weathering of a sequence of alternating limestones and dolostones will result in a color-banded outcrop, because the dolostones tend to weather to a buff-tan color and the limestones become grayish. Erosion of a sequence of alternating sandstones and shales may result in a stair-step outcrop face, because the relatively resistant sandstone beds become vertical cliffs and the relatively weak shale beds create slopes.

1.3.1 Use of Bedding in Structural Analysis

Recognition of bedding is critical in structural analysis. Bedding provides a reference frame for describing deformation of sedimentary rocks, because when sediments are initially deposited, they form horizontal or nearly horizontal layers, a concept referred to as the Law of Original Horizontality. Thus, if we look at an outcrop and see tilting or folding, what we are noticing are deviations in bedding attitude from original horizontality.

In complexly deformed and metamorphosed sedimentary rocks, geologists have to search long and hard to find subtle preserved manifestations of original bedding (e.g., variations in grain size, color, composition). Only by finding bedding can a geologist unravel the folding history of a region. When found, bedding is labeled S0 (pronounced S-zero), where the S is an abbreviation for surface. Later we will discuss surfaces in rocks that are formed by deformation, like rock cleavage, which are labeled S1, S2, and so on.

The study of certain depositional structures within beds and on bedding surfaces is useful in tectonic analysis because they may provide important information on depositional environment (the setting in which the sediment was originally deposited), on stratigraphic facing or younging direction (the direction in which strata in a sequence are progressively younger), and on current direction (the direction in which fluid was flowing during deposition). Facing indicators allow you to determine whether a bed is right-side-up (facing up) or overturned (facing down), with respect to the Earth’s surface. Recognition of facing is powerful both for stratigraphic studies and for structural studies. For example, the structural interpretation of a series of parallel beds in two adjacent outcrops depends on the facing—if the facing is the same in both outcrops, then the strata are probably homoclinal, meaning that they have a uniform dip. However, if the facing is opposite, then the two outcrops are likely on different limbs of a fold whose hinge area is not exposed.

1.3.1.1 Graded Beds and Cross Beds

Patterns within beds may contain information about stratigraphic facing and current directions that are often critical for tectonic interpretations. Graded beds display progressive fining of clast/grain size from the base to the top (Figure 2.2), and are a consequence of deposition from turbidity flows. A turbidity flow is a cloud of sediment that moves down a slope under water because the density of the sediment-water mixture is greater than that of clean water, and denser liquids sink through less dense liquids. Turbidite flows are triggered by major storms or earthquakes (because of their association with seismicity, the occurrence of turbidites may indicate that the sediment source region was tectonically active), and move down gentle slopes at considerable speed. Typically, a flow is confined to a submarine channel or canyon; when a broadening of the channel or a decrease in slope slows the speed of a turbidity current, the sediment cloud settles. During settling, the largest grains fall first, and the finest grains last. Each turbidity flow produces a separate graded sequence or a turbidite, which is often capped by pelagic sediment, meaning deep-marine sediments like clay and plankton shells. Turbidites display an internal order, called a Bouma sequence [1] (Figure 1.5), which reflects changing hydrodynamic conditions as the turbidity current slows down.

FIGURE 1.5. Graded bedding in a turbidite (or Bouma) sequence.

In pre–plate tectonics geological literature (i.e., pre- Beatles), thick sequences of turbidites were referred to as flysch, a term originating from Alpine geology. Flysch was considered to be an orogenic deposit, meaning a sequence of strata that was deposited just prior to and during the formation of a mountain range. Exactly why such strata were deposited, however, was not understood. Modern geologists now realize that the classical flysch sequences are actually turbidites laid down in a deep trench marking an active plate boundary (like a subduction zone). Turbidite flows are common in ocean trenches, which represent seismically active, convergent plate margins. During and after deposition, the trench turbidites are scraped up and deformed by continued convergence between plates, and may eventually be caught in a continental collision zone.

FIGURE 1.6. (a) Terminology of cross bedding, and (b) cross beds in the Coconino Sandstone, Oak Creek Canyon (Arizona). [2.3]

Cross beds are surfaces within a bed that are oblique to the overall bounding surfaces of the bed (Figure 1.6). Cross beds, which are defined by subtle partings or concentrations of grains, form when sediment moves from the windward or upstream side of a dune, ripple, or delta to a face on the leeward or downstream side, where the current velocity is lower and the sediment settles out. Thin beds parallel to the upper bounding surface are called topset beds, the inclined layers deposited parallel to the slip face are called foreset beds, and the thin beds parallel to the lower bounding surface are called bottomset beds (Figure 1.6a). The foreset beds, which typically are curved (concave up) and merge with the topset and bottomset beds, are the cross beds. If the topset beds and the upper part of the foreset beds are removed by local erosion, the bottomset beds of the next higher layer of sediment are juxtaposed against the foreset beds of the layer below. Thus, cross beds tend to be truncated at the upper bedding surface, whereas they are asymptotic to the lower bedding surface (Figure 1.6b). This geometry provides a clear stratigraphic facing indicator. The current direction in a cross-bedded layer is taken to be approximately perpendicular to the intersection between the truncated foresets and overlying bed.

1.3.1.2 Surface Markings

Local environmental phenomena, such as rain, desiccation (i.e., drying), current traction, and the movement of organisms, affect the surface of a bed of sediment. If the sediment is unlithified, these phenomena leave an imprint that is known as a surface marking.

FIGURE 1.7. Asymmetric ripple marks. The arrow indicates approximate current direction during deposition.

For example, ripple marks are ridges and valleys on the surface of a bed formed as a consequence of fluid flow. If the current flows back and forth, as along a beach, the ripples are symmetric, but if they form in a uniformly flowing current, they are asymmetric (Figure 1.7). The crests of symmetric ripples tend to be pointed, whereas the troughs tend to be smooth curves. Thus, symmetric ripples are good facing indicators. Asymmetric ripples are not good facing indicators, but do provide current directions. Flute casts are asymmetric troughs formed by vortices (mini tornadoes) within the fluid that dig into the unconsolidated substrate. The troughs are deeper at the upstream end, where the vortex was stronger. They get shallower and wider at the downstream end, because the vortex dies out. Flute casts can be used as stratigraphic facing indicators.

1.3.1.3 Disrupted Bedding

Load casts, also called ball-and-pillow structures, are bulbous protrusions extending downward from a sand layer into an underlying mud or very fine sand layer (Figure 1.8). They form prior to lithification where a denser sand lies on top of less dense mud and a disturbance by a storm or an earthquake causes blobs of sand to sink into the underlying mud. Load casts are useful stratigraphic facing indicators when they retain some connection to the host layer.

FIGURE 1.8. Load cast or ball-and-pillow structure; wine stakes for scale (Eifel, Germany).

Clastic dike in Proterozoic sandstones. Note that the dike sharply cuts across bedding and that very coarse clasts are preserved in its center (Sudbury, Ontario).

Disrupted bedding in turbidite; hammer for scale (Cantabria, Spain).

Where sand and mud layers are progressively buried, it is typical for the mud layers to compact and consolidate before the sand layers do. As a consequence, the water in the sand layer is under pressure. If an earthquake, storm, or slump suddenly cracks the permeability barrier surrounding the sand, water and sand are released and forced into the mud along cracks. When this happens near the Earth’s surface, little mounds of sand, called sand volcanoes, erupt at the ground surface. The resulting wall-like intrusions of sand (or in some localities, even conglomerate) are called clastic dikes (Figure 1.8). At depth, partially consolidated beds of sand and mud break into pieces, resulting in a chaotic layering that is known, simply, as disrupted bedding (Figure 1.8).

Studies of disrupted bedding, sedimentary dikes, and sand volcanoes in lake and marsh deposits provide an important basis for determining the recurrence interval of large earthquakes. In these studies, investigators dig a trench across the deposit and then look for disrupted intervals within the sequence. Radiocarbon dating of organic matter in the disrupted layers defines the absolute age of disruption events and allows us to estimate the recurrence of earthquakes.

1.3.2 Conformable and Unconformable Contacts

Earlier we defined a contact as any surface between two geologic units. There are three basic types of contacts:

1) depositional contacts, where a sediment layer is deposited over preexisting rock;

2) fault contacts, where two units are juxtaposed by a fracture on which sliding has occurred; and

3) intrusive contacts, where one rock body cuts across another rock body. In this section, we consider in more detail the nature and interpretation of depositional contacts.

Relatively continuous sedimentation in a region leads to the deposition of a sequence of roughly parallel sedimentary units in which the contacts between adjacent beds do not represent substantial gaps in time. Gaps in this context can be identified from gaps in the fossil succession. The boundary between adjacent beds or units in such a sequence is called a conformable contact. For example, we say, “In eastern New York, the Becraft Limestone was deposited conformably over the New Scotland Formation.” The New Scotland Formation is an argillaceous limestone representing marine deposition below wave base, whereas the Becraft Limestone is a pure, coarse-grained limestone representing deposition in a shallow-marine beach environment. Bedding in the two units is parallel, and the contact between these two units is gradational.

FIGURE 1.9. The principal types of unconformities:
(a) a disconformity,
(b) an angular unconformity,
(c) a nonconformity, and
(d) a buttress unconformity.

If there is an interruption in sedimentation, such that there is a measurable gap in time between the base of the sedimentary unit and what lies beneath it, then we say that the contact is unconformable. For example, we say, “In eastern New York, the Upper Silurian Rondout Formation is deposited unconformably on the Middle Ordovician Austin Glen Formation,” because Upper Ordovician and Lower Silurian strata are absent. Unconformable contacts are generally referred to as unconformities, and the gap in time represented by the unconformity (that is, the difference in age between the base of the strata above the unconformity and the top of the unit below the unconformity) is called a hiatus. In order to convey a meaningful description of a specific unconformity, geologists distinguish among four types of unconformities that are schematically shown in Figure 1.9:

• Disconformity - beds of the rock sequence above and below the unconformity are parallel to one another, but there is a measurable age difference between the two sequences. The disconformity surface represents a period of non-deposition and/or erosion (Figure 1.9a).

• Angular unconformity - strata below the unconformity have a different attitude than strata above the unconformity. Beds below the unconformity are truncated at the unconformity, while beds above the unconformity roughly parallel the unconformity surface. Therefore, if the unconformity is tilted, the overlying strata are tilted by the same amount. Because of the angular discordance at angular unconformities, they are quite easy to recognize in the field. Their occurrence means that the sub-unconformity strata were deformed (tilted or folded) and then were truncated by erosion prior to deposition of the rocks above the unconformity. Therefore, angular unconformities are indicative of a period of active tectonism. If the beds below the unconformity are folded, then the angle of discordance between the super- and sub-unconformity strata will change with location, and there may be outcrops at which the two sequences are coincidentally parallel (Figure 1.9b).

• Nonconformity – This term is used for unconformities at which strata were deposited on a basement of older crystalline rocks. The crystalline rocks may be either plutonic or metamorphic. For example, the unconformity between Cambrian strata and Precambrian basement in the Grand Canyon is a nonconformity (Figure 1.9c).

• Buttress unconformity - This type of unconformity (also called onlap unconformity) occurs where beds of the younger sequence were deposited in a region of significant pre-depositional topography. Imagine a shallow sea in which there are islands composed of older bedrock. When sedimentation occurs in this sea, the new horizontal layers of strata terminate at the margins of the island. Eventually, as the sea rises, the islands are buried by sediment. But along the margins of the island, the sedimentary layers appear to be truncated by the unconformity. Rocks below the unconformity may or may not parallel the unconformity, depending on the pre-unconformity structure. Note that a buttress unconformity differs from an angular unconformity in that the younger layers are truncated at the unconformity surface (Figure 1.9d).

grenvilleunconf

FIGURE 1.10. Unconformable contact between mid-Proterozoic Grenville gneiss (dark gray) and Cambrian sandstone and Pleistocene soils (southern Ontario, Canada).

Unconformities represent gaps in the rock record that can range in duration from thousands of years to billions of years, so they are time markers as well. Examples of great unconformities, representing millions or billions of years, occur in the Canadian Shield, where Pleistocene till buries Proterozoic and Archean gneisses. In Figure 1.10 the classic unconformity between Paleozoic sedimentary rocks and Precambrian gneisses is shown and many introductory geology books show this contact in the Grand Canyon.

FIGURE 1.11. Classic outcrop of an angular unconformity in the Caledonides at Siccar Point (Scotland). The hammerhead rests on the unconformity, which is tilted due to later deformation.

It is a special experience to put your finger on a major unconformity and to think about how much of Earth’s history is missing at the contact. Imagine how James Hutton felt when, in the late eighteenth century, he stood at Siccar Point along the coast of Scotland (Figure 1.11), and stared at the Caledonian unconformity between shallowly dipping Devonian Red Sandstone and vertically dipping Silurian strata and, as the present-day waves lapped on and off the outcrop and deposited new sand, suddenly realized what the contact meant. His discovery is one of the most fundamental in field geology.

FIGURE 1.12. Some features used to identify unconformities: (a) scour channels in sediments, (b) basal conglomerate, (c) age discordance from fossil evidence, and (d) soil horizon or paleosol.

How do you recognize an unconformity (Figure 1.12) in the field today? Well, if it is an angular unconformity or a buttress unconformity, there is an angular discordance between bedding above and below the unconformity. A nonconformity is obvious, because crystalline rocks occur below the contact. Disconformities, however, can be more of a challenge to recognize. If strata in the sequence are fossiliferous, and you can recognize the fossil species and know their age, then you can recognize a gap in the fossil succession. Commonly, an unconformity may be marked by a surface of erosion, as indicated by scour features, or by a paleosol, which is a soil horizon that formed from weathering prior to deposition of the overlying sequence. Some unconformities are marked by the occurrence of a basal conglomerate, which contains clasts of the rocks under the unconformity. Recognition of a basal conglomerate is also helpful in determining whether the contact between strata and a plutonic rock is intrusive or whether it represents a nonconformity.

1.3.3 Compaction and Diagenetic Structures

When a clastic sediment initially settles, it is a mixture primarily of grains and water. The proportion of solid to fluid varies depending on the type of sediment. Gooey mud, which consists of clay and water, contains more water than well-packed sand. Progressive burial of sediment squeezes the water out, and the sediment compacts. Compaction results in a decrease in porosity (>50% in shale and >20% in sand) that results in an increase in the density of the sediment.

Lateral variation in the amount of compaction within a given layer, or contrasts in the amount of compaction in a vertical section, is a phenomenon called differential compaction. Differential compaction within a layer can lead to lateral variation in thickness that is called pinch-and-swell structure. Pinch-and-swell structure can also form as a consequence of tectonic stretching, so again, you must be careful when you see the structure to determine whether it is a depositional structure or a tectonic structure.

The compaction of mud leads to development of a preferred orientation of clay in the resulting mudstone. Clay occurs in tiny flakes shaped like playing cards. In a wet sediment the flakes are not all parallel to one another, as in a standing house of cards, but after compaction the flakes are essentially parallel to one another, as in a collapsed house of cards. The preferred orientation of clay flakes, as we have seen, leads to bedding plane fissility that produces a shale. For example, in the Gulf Coast sequence of the southern United States this progression is preserved in drill cores that were obtained to study the relationship between oil maturation and clay mineralogy. In contrast, the compaction of sand composed of equant grains causes the grains to pack together more tightly, but produces little rock fabric.

FIGURE 1.13. Pitted pebble; coin for scale.

FIGURE 1.14. Suture-like stylolites in limestone. Pocket knife for scale.

Deeper compaction can cause pressure solution, a process by which soluble grains preferentially dissolve along the faces at which stress is the greatest. In pure limestones or sandstones, this process causes grains to suture together, meaning that grain surfaces interlock with one another like jigsaw puzzle pieces. In conglomerates, the squeezing together of pebbles results in the formation of indentations on the pebble surfaces creating pitted pebbles (Figure 1.13). In limestones and sandstones that contain some clay, the clay enhances the pressure solution process. Specifically, pressure solution occurs faster where the initial clay concentration is higher. As a result, distinct seams of clay residue develop in the rock. These seams are called stylolites (Figure 1.14). In rocks with little clay (<10%), stylolites are jagged and tooth-like in cross section, like the sutures in your skull. The teeth are caused by the distribution of grains of different solubility along the stylolite. In rocks with more clay, the stylolites are wavy and the teeth are less pronounced, because the clay seams become thicker than tooth amplitude. Some of the dissolved ions that are removed at pressure-solved surfaces precipitate locally in the rock in veins or as cement in pore spaces, whereas some get transported out of the rock by moving groundwater. The proportion of reprecipitated to transported ions is highly variable, but as much as 40% of the rock can be dissolved and removed during formation of stylolites.

Some sedimentary rocks exhibit color banding that cuts across bedding. This color banding, which is called Liesegang banding, is the result of diffusion of impurities, or of reactions leading to alternating bands of oxidized and reduced iron. Because it can be mistaken for bedding or cross bedding, it is mentioned in the context of primary sedimentary structures. To avoid mistaken identity, search the outcrop to determine whether sets of bands cross each other (possible for Liesegang bands, but impossible for bedding), and whether the bands are disrupted at fractures or true bedding planes, because these are places where the diffusion rate changes.

1.3.4 Penecontemporaneous Structures

If sediment layers have an initial dip, meaning a gentle slope caused by deposition on a preexisting slope or tilting prior to full lithification in a tectonically active region, gravity can pull the layers down the slope. The ease with which sediments move down a slope is increased by fluid pressure in the layers, which effectively keeps the layers apart. Movement is resisted by weak electrostatic adhesion between grains, but this resistance can be overcome by the energy of an earthquake or a storm, and the sediment will move down the slope. If the sediment completely mixes with water and becomes a turbid suspension flowing into deeper water, then all of the preexisting primary structure is lost and the grains are resedimented as a new graded bed (turbidite; see above) farther down the slope. If the flowing mixture of sediment and water is dominantly sediment, it churns into a slurry containing chunks and clasts that are suspended in a matrix. Such slurries are called debris flows, and where preserved in a stratigraphic sequence, they become matrix-supported, poorly sorted conglomerates containing a range of clast sizes and shapes.

If the beds were lithified sufficiently prior to movement, so that they maintain cohesion, then the movement is called slumping. During slumping, the sedimentary layers tend to be folded and pulled apart and are thrust over one another. The folds and faults formed during this slumping are called penecontemporaneous structures, because they formed almost (Greek prefix pene) at the same time as the original deposition of the layers. Penecontemporaneous folds and faults are characteristically chaotic. The folds display little symmetry, and folds in one layer are of a different size and orientation than the structures in adjacent layers. Penecontemporaneous faults are not associated with pronounced zones of brittle fracturing (we turn to the characteristics of this behavior in Chapter X). One key to the recognition of slump structures in a sedimentary sequence is that the deformed interval is intraformational, meaning that it is bounded both above and below by relatively undeformed strata. Commonly, intervals of penecomtemporaneous structures occur in a sequence that also includes debris flows and turbidites, all indicative of an unstable depositional environment. While slump structures can be mistaken for folding adjacent to a tectonic detachment fault, the opposite, tectonic folds mistakenly interpreted as slump structures, may also occur. Not a simple matter to distinguish between the two!

We tend to think of debris flows and landslides as being relatively small structures, capable of disrupting a hillslope and perhaps moving a cottage or two, but generally not much more. However, the geologic record shows that catastrophic landslides of enormous dimension have occurred on occasion. In northern Wyoming, for example, a giant Eocene slide in association with volcanic eruption displaced dozens of mountain-sized blocks and hundreds of smaller blocks. One such large block, Heart Mountain, moved intact for several tens of kilometers, apparently riding on a cushion of compressed air above a nearly planar subhorizontal (detachment) fault (Figure 1.15).

FIGURE 1.15. The Heart Mountain detachment, emplacing Paleozoic carbonates on Eocene deposits (Wyoming).

1.4 SALT STRUCTURES

Salt is a sedimentary rock that forms by the precipitation of evaporite minerals (typically halite [NaCl] and gypsum or anhydrite calcium sulfates) from saline water. Salt deposits accumulate in any sedimentary basin, meaning a low region that is the site of deposition, where saline water, such as seawater, evaporates sufficiently for salt to precipitate. Particularly thick salt deposits lie at the base of passive-margin basins, so named because they occur along tectonically inactive edges of continents. To see how these basins form, imagine a supercontinent that is being pulled apart, like Pangea in the early Mesozoic. This process, called rifting, involves brittle and ductile faulting, the net result of which is to thin the continental lithosphere until it breaks and an oceanic ridge is formed. During the early stages of rifting, the rift basin is dry or contains freshwater lakes. Eventually, the floor of the rift drops below sea level and a shallow sea forms. If evaporation rates are high, various salts (typically, halite and gypsum/anhydrite) precipitate out of the seawater and are deposited on the floor of the rift. When the rift evolves into an open ocean, the continental margins become passive margins that gradually subside. With continued subsidence, the layer of evaporite (salt) is buried by clastic sediments and carbonates typical of continental-shelf environments. We’ll discuss this tectonic environment in more detail later (Extension Tectonics, Chapter X), but for now we leave you a picture of a thick pile of sediment with a layer of salt near its base. This is the starting condition for the formation of salt intrusions.

Salt differs from other sedimentary rocks in that it is much weaker and, as a consequence, is able to flow plastically, like a viscous fluid, under conditions in which other sedimentary rocks behave in a brittle fashion. In some cases, deformation of salt is due to tectonic faulting or folding, but because salt is so weak, it may deform solely in response to gravity, and thereby cause deformation of surrounding sedimentary rock. As gravity is the only reason for salt movement, the deformation resulting from its movement is called halokinesis (combining the Greek words for salt and movement, respectively), and the resulting body of salt creates a salt structure.

1.4.1 Why Halokinesis?

salt movement, or halokinesis, begins in response to three factors:

(1) the development of a density inversion;

(2) differential loading; and

(3) the existence of a slope at the base of a salt layer.

All three of these factors occur in a passive-margin basin setting. Salt is a nonporous and essentially incompressible material. So when it gets buried deeply in a sedimentary pile, it does not become denser. In fact, salt actually gets less dense with depth, because at greater depths it becomes warmer and expands. Other sedimentary rocks (like sandstone and shale), in contrast, form from sediments that originally had high porosity and thus become denser with depth because the pressure caused by overburden makes them compact. This contrast in behavior, in which the density of other sedimentary rocks exceeds the density of salt at depths greater than about 6 km, results in a density inversion, meaning a situation where denser rock lies over less dense rock. Salt density is about 2200 kg/m3, whereas the density of the sedimentary rocks is about 2500 kg/m3. A density inversion is an unstable condition because the salt has positive buoyancy. Positive buoyancy means that forces in a gravity field cause lower density materials to try to rise above higher density materials, thereby decreasing the overall gravitational potential energy of the system. Negative buoyancy, the reverse, is a force that causes a denser material to sink through a less dense material (see Section 2.2.4). A familiar example of positive buoyancy forces is the push that your hand feels when you try to hold an air-filled balloon under water. Holding a brick under water illustrates negative buoyancy. When the positive buoyancy force exceeds the strength of the salt and is sufficient to upwarp strata that lie over the salt structure, then it will contribute to the formation of the salt structure.

Differential loading of a salt layer takes place when the downward force on the salt layer caused by the weight of overlying strata varies laterally. This may occur where there are variations in the thickness or composition of the overlying strata, variations in the original surface topography of the salt layer, or changes in the thickness of the overlying strata due to faulting. Regardless of its cause, differential loading creates a situation in which some parts of the salt layer are subjected to a greater vertical load than other parts, and the salt is squeezed from areas of higher pressure to areas of lower pressure. For example, imagine a layer of salt whose upper surface initially bulges upward to form a small “dome.” The weight of a column of rock and water from sea level down to a horizontal surface in the salt layer on either side of the dome is greater than the weight of the column on the top of the dome, because salt is less dense than other sedimentary rocks. As a consequence, salt is squeezed into the dome, making it grow upwards. A salt layer that has provided salt for the production of a salt structure, and thus has itself been changed by halokinesis, is called the source layer.

FIGURE 1.16. Satellite image of salt domes (hills) and salt glaciers (dark tongues) in the Zagros mountains (western Iran). [NASA]

The combination of differential loading and buoyancy force drives salt upward through the overlying strata until it reaches a level of neutral buoyancy, meaning the depth at which it is no longer buoyant. At this level, salt has the same density as surrounding strata. The density of mildly compacted clastic strata equals that of salt at depths of around 500–1500 m below the surface of the basin, depending on composition. At the level of neutral buoyancy, salt may begin to flow laterally, much like a thick pile of maple syrup flows laterally over your pancakes. This process, which is also driven by gravity (above the level of neutral buoyancy, the salt is subjected to a negative buoyancy force), is known as gravity spreading. Where salt is extruded at the land surface, it becomes a salt glacier (Figure 1.16). At the seafloor, salt also spreads like a salt glacier, except that during movement it continues to be buried by new sediment.

1.4.2 Geometry of Salt Structures and Associated Processes

In response to positive buoyancy force and to differential loading, salt will flow upward from the source bed, which thins as a consequence. If the source bed thins to the point of disappearing and the strata above the source bed and below the source bed become juxtaposed, we say that the contact between these two beds is a primary weld. In general, a weld is any contact between strata that were once separated by salt. At any given time, a region may contain salt structures at many stages of this development. Geologists working with salt structures have assigned a rich vocabulary to these structures based on their geometry; some are shown in Figure 1.17. The name assigned to a specific structure depends on its shape today, but in the context of geologic time, this shape may be only temporary.

FIGURE 1.17. Schematic diagram showing the stages in the formation of salt structures and associated terminology. Structural maturity and size increase toward the structures in the rear. Sequence in (a) shows structures rising from line sources, whereas structures in (b) originate from point sources.

Because salt both rises up into preexisting strata and rises during the time of deposition of overlying strata, geologists distinguish between two types of salt structure growth. If the salt rises after the overlying strata have already been deposited, then the rising salt will warp and eventually break through the overlying strata. This process is called upbuilding. If, however, the rise of the salt relative to the source layer occurs coevally with further deposition, the distance between the source layer and the surface of the basin also increases. This process is called downbuilding. As salt moves, it deforms adjacent strata and creates complex folds and local faults. When salt diapirs approach the surface, the overlying strata are arched up and therefore are locally stretched, resulting in the development of normal faults in a complex array over the crest of the salt structure.

The structural geometry of passive-margin basins is complicated because sedimentation continues during salt movement. Sedimentary layers thicken and thin as a consequence of highs and lows in elevations caused by the salt, and the resulting differential compaction causes further salt movement. The sedimentation pattern may change in a locality when a salt structure drains out and flows into a structure at another locality. Thus, in regions where halokinesis occurs, it is common to find places where an arch evolves into a basin, or vice versa, a process we call inversion.

1.4.3 Other Gravity-Driven Structures

The formation of salt structures is a dynamic process that is intimately linked to faulting in the overlying strata. Salt is so weak that it makes a good glide horizon on which faulting and movement of overlying strata occurs. In fact, on many passive margins, a thick package of sedimentary rocks tends to detach and slump seaward, gliding on a detachment fault in the layer of salt at its base. This movement resembles the slumping of sediment of a hillslope, though the scale of displacement it quite different. As slumping occurs, the landward portion of the basin is stretched and is therefore broken by a series of normal faults whose dip tends to decrease with depth. This change in dip with depth makes the faults concave up, which are called listric faults. As movement occurs on a listric fault, the strata above the fault arch into rollover folds (Figure 1.18).

FIGURE 1.18. Cross sections showing “down-to-Gulf” type movement of a passive-margin salt wedge; the sections also show listric normal faults and salt rollers.

The sequence (a) through (d) shows successive stages in the evolution of the margin, and accompanying extension.

Many listric normal faults intersect the ground surface in southern Texas, because this region is part of the passive-margin basin along the Gulf Coast. Because the faults dip south and they transport rock toward the Gulf of Mexico, they are sometimes called down-to-Gulf faults. Slip on these faults thins the stratigraphic section above the salt layer and thus results in differential loading. As a result, the salt can rise beneath the fault, evolving from a salt dome to a salt roller, to a diapir that eventually cuts the overlying fault. At the toe of the passive margin wedge, a series of thrust faults develop to accommodate displacement of the seaward-moving section, just as thrusts develop at the toe of a hillside slump.

1.4.4 Importance of Salt Structures

Why spend so much time dealing with salt structures and passive margins? Simply because these regions are of great economic and societal importance. Passivemargin basins are major oil reservoirs, and much of the oil in these reservoirs is trapped adjacent to salt bodies. Oil rises in the upturned layers along the margins of the salt body and is trapped against the margin of the impermeable salt. In recent years, salt bodies in passive-margin basins are being used as giant storage tanks for gas or oil, and are being considered as potential sites for the storage of nuclear waste. This is one of many examples where structural geology is central to reaching important societal decisions.

1.5 IGNEOUS STRUCTURES

You may recall from your introductory geology course that there are two principal classes of igneous rocks, and that these classes are distinguished from one another based on the environment in which the melt cools. Extrusive rocks are formed either from lava that flowed over the surface of the Earth and cooled under air or water, or from ash that exploded out of a volcanic vent. Intrusive rocks cooled beneath the surface of the Earth. During the process of intruding, flowing, settling and/or cooling, igneous rocks can develop primary structures. In this context, we use the term “primary structure” to refer to a fabric that is a consequence of igneous processes.

Where do magmas, the melt phase of igneous rocks, come from? If you have not had a course in igneous petrology, we’ll quickly outline the nature of magmatic activity. Magma forms where conditions of heat and pressure cause existing rock (either in the crust or in the mantle) to melt. Commonly, only certain minerals within the solid rock melt (the ones that melt at a lower temperature), in which case we say that the rock has undergone partial melting. A magma formed by partial melting has a composition that differs from that of the rock from which it was extracted. For example, a 1–6% partial melt of ultramafic rock (peridotite) in the mantle yields a mafic magma which, when solidified, forms the gabbro and basalt that characterizes oceanic crust. Melting of an intermediate-composition crustal rock (diorite) yields a silicic magma which, when solidified, forms granite or rhyolite. Once formed, magma is less dense than the surrounding rock, and buoyancy forces cause it to rise. The density decrease is a consequence of the expansion that accompanies heating and melting, the formation of gas bubbles within the magma, and the difference in composition between magma and surrounding rock.

Magma moves by oozing up through a network of cracks and creeping along grain surfaces. The difference between the pressure within the magma and the pressure in the surrounding rock is so substantial that, as magma enters the brittle crust, it can force open new cracks. Magma continues to rise until it reaches a level of neutral buoyancy, defined as the depth where pressure in the magma equals lithostatic pressure in the surrounding rock, meaning that the buoyancy force is zero. At the level of neutral buoyancy, the magma may form a sheet intrusion, or may pool in a large magma chamber that solidifies into a bloblike intrusion called a pluton. If the magma pressure is sufficiently high, the magma rises all the way to the surface of the Earth, like water in an artesian well, and is extruded at a volcano.

1.5.1 Structures Associated with Sheet Intrusions

One aspect of sheet intrusions that is of particular interest to structural geologists is their relationship to stress. In Chapter 3, we will introduce the concept of stress in detail, but for now, we point out that stress acting on a plane is defined as the force per unit area of the plane. Intuitively, therefore, you can picture that the Earth’s crust is held together least tightly in the direction of the smallest stress, which we call the direction of least principal stress. Sheet intrusions, in general, form perpendicular to the direction of the least principal stress, assuming no preexisting planes of weakness (such as faults). For example, in regions where the greatest stress is caused by the weight of the overlying rocks, and is therefore vertical, the least principal stress is horizontal, so vertical dikes form. Dike swarms, which are arrays of subparallel dikes occurring over broad regions of the crust, probably represent intrusion at depth in association with horizontal extension, which causes the horizontal stress to be tensile.

FIGURE 1.19. Types of sheet intrusions around a volcano.

Not all dikes occur in parallel arrays. In the immediate vicinity of volcanoes, the ballooning and/or collapse of a magma chamber locally modifies the stress field and causes a complex pattern of fractures. As a result, the pattern of dikes around a volcano (Figure 1.19) includes ring dikes, which have a circular trace in map view, and radial dikes, which run outward from the center of the volcano like spokes of a wheel. At a distance from the volcano, where the local effects of the volcano on the stress field are less, radial dikes may change trend to become perpendicular to the regional least principal stress.

Sills, or subhorizontal sheetlike intrusions, form where local stress conditions cause the least principal stress to be vertical, and/or where there are particularly weak horizontal partings in a stratified sequence. Sill intrusion can result in the development of faults. If the thickness of the intrusion changes along strike, there is differential movement of strata above the intrusion. Laccoliths resemble sills in that they are concordant with strata at their base, but unlike sills, they bow up the overlying strata to create a dome. For example, the laccoliths of the Henry Mountains in Utah are several kilometers in diameter.

Sheet intrusions occur in all sizes, from thin seams measured in centimeters, to the Great Dike of Zimbabwe, which is nearly 500 km long and several kilometers wide. Considering the dimensions of large intrusions (tens to hundreds of kilometers long), it is important to keep in mind that a large volume of magma can flow past a given point. The occurrence of flow may be recorded as primary igneous structures in the rock. For example, examination of dike-related structures may show the presence of drag folds, scour marks, imbricated phenocrysts, and flow foliation, particularly along the walls of an intrusion.

1.5.2 Structures Associated with Plutons

The nature of primary structures found in plutonic rocks depends on the depth in the Earth at which the intrusion solidified, because these structures reflect the temperature contrast between the intrusion and the country rock. Remember that the Earth gets warmer with depth: at the surface, the average temperature is ~10°C, whereas at the center it may be as much as 4000°C. The change in temperature with depth is called the geothermal gradient. In the shallow crust, the geothermal gradient is in the range of 20°C/km to 40°C/km. At greater depths, however, the gradient must be less, because temperatures at the continental Moho (at about 40 km depth) are in the range of about 700°C (~15°C/km), and temperatures at the base of the lithosphere (at about 150 km) are in the range of about 1280°C (<10°C/km). The origin of the increased geothermal gradient near the surface is the concentration of radioactive elements in the minerals of more silicic rocks. Granitic magma begins to solidify at temperatures between 550°C to 800°C. Therefore, the temperature contrast between magma and country rock decreases with depth. In the case of shallow-level plutons, which intrude at depths of less than about 5 km, the contrast between magma and country rock is several hundred degrees. Contacts at the margins of such shallow intrusions are sharp, so that you can easily place your finger on the contact. Angular blocks of country rock float in the magma near the contact, and the country rock adjacent to the contact may be altered by fluids expelled by the magma or may be baked (a rock type called hornfels). At greater depth, the temperature contrast between magma and country rock decreases, and contacts are more gradational, until the country rock itself is likely to be undergoing partial melting. Minerals that melt at lower temperatures (like quartz and feldspar) turn to liquid, while refractory minerals (that is, minerals that melt at higher temperatures, such as amphibole and pyroxene) are still solid, though quite soft. Movement of the melt causes the soft solid layers to be contorted into irregular folds. When this mass eventually cools, the resulting rock, which is composed of a marble-cake-like mixture of light and dark contorted bands, is called a migmatite (Figure 1.20).

FIGURE 1.20. Migmatite from the North Cascades (Washington State, USA) showing complex folding and disruption.

Many plutons exhibit an intrusion foliation that is particularly well developed near the margin of the pluton and is subparallel to pluton–host rock contact. This foliation is defined by alignment of inequant crystals and by elongation of chunks of country rock or early phases of pluton that were incorporated in the magma (called xenoliths, which means “foreign rocks”). Such fabric is a consequence of shear of the magma against the walls of the magma chamber and of the flattening of partially solidified magma along the chamber walls in response to pressure exerted as new magma pushes into the interior of the chamber. Similarly, intrusion foliation can be developed along the margins of dikes.

Because intrusion foliation forms during the formation of the rock and is not a consequence of tectonic movements, it is a nontectonic structure. However, it may be difficult to distinguish from schistosity resulting from tectonic forces. Plutons tend to act as mechanically strong blocks, so that regional deformation is deflected and concentrated along the margins of the pluton. Interpretation of a particular foliation therefore depends on regional analysis and the study of deformation microstructures. For example, if the fabric remains parallel to the boundary of the intrusion, even when the boundary changes and individual grains show no evidence for solid-state deformation, the foliation is likely a primary igneous structure. The distinction between tectonic and primary structures in plutons has proven to be quite difficult and more often than not is ambiguous.

1.5.3 Structures Associated with Volcanic Extrusion

As basaltic lava flows along the surface of the Earth, the surface of the flow may wrinkle into primary folds that resemble coils of rope, or may break into a jumble of jagged blocks that resembles a breccia. Lava flows with the rope-coil surface are called Pahoehoe flows, and lavas with the broken-block surface are called Aa flows. The wrinkles in a Pahoehoe lava should not be mistaken for tectonic folds, and the jumble of blocks caused by autobrecciation (that is, breaking up during flow) of Aa lavas should not be mistaken for a tectonic breccia related to faulting.

If basaltic lava is extruded beneath seawater, the surface of the flow cools quickly, and a glassy skin coats the surface of the flow. Eventually, the pressure in the glass-encased flow becomes so great that the skin punctures, and a squirt of lava pushes through the hole and then quickly freezes. The process repeats frequently, resulting in a flow composed of blobs (centimeters to meters in diameter) of lava. Each blob, which is called a pillow, is coated by a rind of finegrained to glassy material. As the pillows build out into a large pile, creating a pillow basalt, successive pillows flow over earlier pillows and, while still soft, conform to the shape of the earlier flow surface (Figure 1.21). As a result, pillows commonly have a rounded top and a pointed bottom (the “apex”) in cross section, and this shape can be used as a stratigraphic facing criterion.

FIGURE 1.21. Pillow basalt from the Point Sal Ophiolite, California (USA). The asymmetric shape of the pillows and location of the “points” (apex) indicate that the stratigraphic top of the flow is up.

In 1902, Mt. Pelee on the Caribbean island of Martinique erupted. It was a special kind of eruption, for instead of lava flows, a spine of rhyolite rose day by day from the peak of the volcano. This spine, as it turned out, was like the cork of a champagne bottle slowly being worked out. When the cork finally pops out of the champagne bottle, a froth of gas and liquid flows down the side of the bottle. Likewise, when, on the morning of May 2, the plug exploded off the top of Mt. Pelee, a froth of hot (>800°C) volcanic gas and ash floated on a cushion of air and rushed down the side of the mountain at speeds of up to 100 km/h. This ash flow engulfed the town of St. Pierre, and in an instant, almost 30,000 people were dead. When the ash stopped moving, it settled into a hot layer that welded together. Such a layer of welded tuff is called an ignimbrite, often displaying a foliation. Volcanic ash is composed of tiny glass shards with jagged spinelike forms that are a consequence of very rapid cooling. When the ash settles, the glass shards are still hot and soft, so the compaction pressure exerted by the weight of overlying ash causes the shards to flatten, thereby creating a primary foliation in ignimbrite that is comparable to bedding.

Rhyolitic lavas commonly display subtle color banding, called flow foliation, that has been attributed to flow of the lava before complete solidification. The banding forms because lavas are not perfectly homogeneous materials. Since the temperature is not perfectly uniform, there may be zones in which crystals have formed, while adjacent regions are still molten. Shear resulting from movement of the lava smears out these initial inhomogeneities into subparallel bands. To visualize this, think of a bowl of pancake batter into which you have dripped spoonfuls of chocolate batter. If you slowly stir the mixture, the blobs of chocolate smear out into sheets. Chocolate blobs that were initially nearby would smear into parallel sheets with an intervening band of pancake batter. In the flow, movement of the lava smears out blobs of contrasting texture into layers, which, when the rock finally freezes, have a slightly different texture than adjacent bands and thus are visible markers in outcrop. Commonly, continued movement causes previously formed layers to fold, so flow-banded outcrops typically display complex primary folds.

1.5.4 Cooling Fractures

FIGURE 1.22. Columnar jointing in the Massif Central, France; (a) side view, (b) top view.

As shallow intrusions and extrusive flows cool, they contract. Because of their fine grain size, these bodies are susceptible to forming natural cracks, or joints, in response to the thermal stress associated with cooling. When such joints are typically arranged in roughly hexagonal arrays that isolate columns of rock, the pattern is called columnar jointing (Figure 1.22). Popular tourist and movie director destinations like Devil’s Tower in Wyoming, Giant’s Causeway in Ireland, or the Massif Central in southern France offer spectacular examples. The long axes of columns are perpendicular to isotherms (surfaces of constant temperature) and thus they are typically perpendicular to the boundaries of the shallow intrusion or flow. If you look closely at unweathered columnar joints, the surfaces of individual joints are ribbed. We will learn later that this feature is a consequence of the way in which fractures propagate through rock.

1.6 IMPACT STRUCTURES

Glancing at the Moon through a telescope, the most obvious landforms that you see are craters. Like early Earth, the Moon has been struck countless times by meteors, and each impact has left a scar which, because the Moon is tectonically inactive and has no atmosphere or water, has remained largely unchanged through succeeding eons. The Earth has been pummeled at least as frequently as the Moon, but many objects disintegrate and burn in the atmosphere before reaching the surface, and the scars of many that did strike the surface have been erased by erosion and particularly by tectonics. The vast majority of impacts on the Earth-Moon system occurred prior to about 3.9 Ga, when the solar system contained a multitude of fragments that were not yet incorporated into planets. Considering that 70% of today’s Earth’s surface is underlain by oceanic lithosphere, most of which is less than 200 million years old, and that all but a relatively small portion of the continental crust has either been covered by younger strata or has been involved in plate tectonics, it is not surprising that impact structures are so rare on Earth.

Whereas impact structures on Earth are rare, and often difficult to recognize, they do exist. For our discussion we distinguish three categories, based on the most obvious characteristic of the impact:

(1) relatively recent surficial impacts that are defined by a visible crater,

(2) impacts whose record at the present Earth surface is the disruption of sedimentary strata, and

(3) impacts whose record is a distinctive map-view circular structure in basement.

FIGURE. 1.23. (a) Barringer Meteor Crater of Arizona (USA) and (b) geologic cross section showing distribution of impact structures.

Today there are about 150 impacts recognized on our planet but only a few dozen obvious impact craters can be seen. One of the largest and perhaps most famous is the approximately 50,000-year-old Barringer Meteor Crater in Arizona, which is 1.2 km in diameter and 180 m deep, and is surrounded by a raised rim that is about 50 m high (Figure 1.23a). The size of the impacting object is estimated only to be 30–50 m! As shown in the cross section (Figure 1.23b), the impact created a breccia that is about 200 m thick beneath the floor of the crater. The raised rims of the crater are not only composed of shattered rock ejected from the crater, but are also sites where bedrock has been upturned. Ancient impact sites that are no longer associated with a surficial crater dot the Midcontinent region of the United States. These sites are defined by relatively small (less than a few kilometers across) semicircular disruption zones, in which the generally flat-lying Paleozoic strata of the region are fractured, faulted, and tilted. They were originally called cryptovolcanic structures (from the Greek crypto, meaning “hidden”), because it was assumed that they were the result of underlying explosive volcanism. Typically, steeply dipping normal faults, whose map traces are roughly circular, define the outer limit of these structures. These faults are cross cut by other steep faults that radiate from the center of the structure like spokes of a wagon wheel. This fracture geometry is similar to that around volcanoes, which appeared to support the volcanic interpretation. Near the center of the structure, bedding is steeply dipping, and faulting juxtaposes units of many different ages.

FIGURE 1.24. Shatter cones that were formed by the Sudbury impact that occurred around 1.85 Ga (Ontario, Canada). The apex of a shatter cone points in the direction from which the impacting object came. Pocket knife for scale.

Locally, the strata are broken into huge blocks that jumbled together to create an impact breccia. Throughout this region, rocks are broken into distinctive shatter cones, which are conelike arrays of fractures similar to those found next to a blast hole in rock (Figure 1.24). The apex of the cone points in the direction from which the impacting object came. In impact structures, shatter cones point up, confirming that they were caused by impact from above, as would be the case if the structure was due to an incoming meteor.

Why do impact structures have the geometry that they do? To see why, think of what happens when you drop a pebble into water. Initially, the pebble pushes down the surface of the water and creates a depression, but an instant later, the water rushes in to fill the depression, and the place that had been the center of the depression rises into a dome. In the case of meteor impact against rock, the same process takes place. The initial impact gouges out a huge crater and elastically compresses the rock around the crater. But an instant later, the rock rebounds. At the margins of the affected zone it pulls away from the walls, creating normal faults, and in the center of the zone, it flows upward, creating the steeply tilted beds. Because the rock in the near surface behaves in a brittle manner when this occurs, this movement is accompanied by faulting and brecciation.

The incredibly high pressures that develop during an impact create distinctive changes in the rocks of the impact site. The shock wave that passes through the rock momentarily subjects rocks to very high pressures, a condition that causes shock metamorphism. Shock metamorphism of quartz yields unusual high-density polymorphs like stishovite, and characteristic deformation microstructures. In addition, the kinetic energy of impact is suddenly transformed into heat, with the result that rocks of the impact site are momentarily heated to temperatures as high as 1700°C. At such temperatures, the whole rock melts, only to freeze quickly into glass of the same composition as the original rock. In some cases, melt mixes with impact breccia, and injects into cracks between larger breccia fragments, forming a glasslike rock called pseudotachylyte.

Impact structures affecting now exposed basement crystalline rocks characteristically create distinctive circular patterns of erosion in the basement that stand out in satellite imagery or through geophysical methods. One of the best known basement impact structures is the Sudbury complex in southern Ontario, Canada. Not only are the characteristic features of impact, like shatter cones and pseudotachylyte, readily visible in the field, but the Sudbury impact, occurring about 1.85 Ga, was large enough to affect the whole crust and cause an impact melt that produced valuable economic deposits that have been mined for many years.

1.7 DEFORMATION STRUCTURES

FIGURE 1.25. Collage of deformation structures in the field, showing joints, fold (nappe), (normal) faults, mylonite and cleavage.

The primary focus of the material in this text is deformation structures, which are ubiquitous in the field and produce some of the most inspiring images in geology (Figure 1.25). As we’ll examine in the upcoming chapters, we can use multiple classifications for deformation structures, with the classification based on dominant processes and conditions (called Deformation Regimes) the underlying organization of the material in this book.

I. Classification based on the process of formation, that is, the deformation mechanism:

• Fracturing: related to development or coalescence of cracks in rock.

• Frictional sliding: related to the slip of one body of rock past another, or of grains past one another, both of which are resisted by friction.

• Plasticity: resulting from deformation by the internal flow of crystals without loss of cohesion, or by non-frictional sliding of crystals past one another.

• Diffusion: resulting from material transport either solid-state or assisted by a fluid (dissolution).

II. Classification based on the mesoscopic cohesiveness during deformation:

• Brittle: formed by loss of cohesion across a mesoscopic discrete surface.

• Ductile: formed without loss of cohesion across a mesoscopic discrete surface.

• Brittle/ductile: involving both brittle and ductile aspects.

Note that the scale of observation (in this case, mesoscopic) is critical in the distinction between brittle and ductile deformation, because ductile deformation can involve microscopic-scale fracturing and frictional sliding. These terms are not synomymous with frictional and plastic, as we will describe below.

III. Classification based on the strain significance, in which a reference frame, usually the Earth’s surface, is defined:

• Contractional: resulting in shortening of a region.

• Extensional: resulting in extension of a region.

• Lateral-slip (or Wrench): resulting from movement without either shortening or extension.

Note that shortening in one direction can be, but does not have to be, accompanied by extension in a different direction, and vice versa. Also, regional deformation usually results in the vertical displacement of the Earth’s surface, a component of deformation that is commonly overlooked.

IV. Classification based on the distribution of deformation in a volume of rock:

• Continuous: deformation occurs through the rock body at all scales.

• Penetrative: occurs throughout the rock body, at the scale of observation; up close, there may be spaces between the structures.

• Localized: continuous or penetrative structure occurs only within a definable region.

• Discrete: structure occurs as an isolated feature.

•

1.7.1 Structural Analysis and Scales of Observation

From introductory material, you know what a structure is and you know what a geologist means by deformation. We also know that there is a group of people who call themselves structural geologists. But what do structural geologists do? One way to gain insight into the subject of structural geology is to think about the type of work that structural geologists carry out. Not surprisingly, structural geologists do structural analysis, which involves a range of activities:

Descriptive analysis - The characterization of the shape and appearance of geologic structures. It includes development of a precise vocabulary (jargon) that permits one geologist to create an image of a structure that any other geologist can understand, and development of methods for uniquely describing the orientation of a structure in three-dimensional space. For example, the creation and interpretation of geologic maps and cross-sections (Figure 1.26).
Kinematic analysis - The determination of the movement paths that rocks or parts of rocks have taken during transformation from the undeformed to the deformed state. This subject includes, for example, use of features in rocks to define the direction of movement on a fault.
Strain analysis - The development of mathematical tools for quantifying the displacement and distortion of a rock. This activity includes the search for features in rock that can be measured to define strain.
Dynamic analysis - The development of an understanding of stress and its relation to deformation. This activity includes the use of tools for measuring the present-day state of stress in the Earth, and the application of techniques for interpreting the state of stress responsible for microstructures in rocks.
Mechanism analysis - The study of processes on the atomic scale to grain scale that allow structures to develop. This activity includes study of both fracture and flow of rock.
Tectonic analysis - The study of the relationship between structures and global tectonic processes. This activity includes the study and interpretation of regional-scale or megascopic structural features, and the study of relationships among structural geology, stratigraphy, and petrology.

$C:\Users\vdpluijm\Google Drive\Earth Structure\ES3\ES3 Figures\USGS_GQ-18_1 Map.jpg$

$C:\Users\vdpluijm\Google Drive\Earth Structure\ES3\ES3 Figures\USGS_GQ-18_1 Sections.jpg$

Figure 1.26. Vintage example of a geologic map and cross sections of folded rocks in the US Appalachians (Tennessee). Click on map for larger version.

GoogleEarth: https://psgt.earth.lsa.umich.edu/chapter/1/USGS_GQ-18_1.kmz

[USGS]

From the list above you will notice that many of the definitions refer to the scale of observation. For the results of a structural analysis to be interpretable, the scale of our analysis must be taken into account. For example, a bed of sandstone in a single outcrop in a mountain may appear to be undeformed. But the outcrop may display only a small part of a huge fold that cannot be seen unless you map at the scale of the whole mountain. Structural geologists commonly refer to these relative scales of observation by a series of subjective prefixes. Micro refers to features that are visible optically at the scale of thin sections, or that may only be evident with the electron microscope; the latter is sometimes referred to as submicroscopic. Meso refers to features that are visible in a rock outcrop, but cannot necessarily be traced from outcrop to outcrop. Macro refers to features that can be traced over a region encompassing several outcrops to whole mountain ranges. In some circumstances, geologists use the prefix mega to refer to continental-scale deformational, such as the movements of tectonic plates over time. Of course there are no sharp boundaries between these scales, and their usage will vary with context, but a complete structural analysis tries to integrate results from several scales of observation.

Each scale of observation has its own set of tools. For example, optical and electron microscopes are used for observations on the microscale, while satellites may be used for observations on the macroscale. The mesoscopic recognition and description of rocks and their structures are of fundamental importance to field analysis, which requires a set of eyes (± corrective lenses), a hammer, a compass, and a hand lens. Field work is, in fact, pretty much a low-tech, low-budget affair, except for transportation (e.g., the High Himalayas, Antarctica), or some similarly remote setting that requires extensive logistics (like expeditions, planes, and helicopters). For structural field work we record observations on lithologies and rock structures in notebooks or on portable devices and we measure the orientation of geometric elements with a compass. The compass to a structural geologist (Figure 1.27) is like the handlens to a petrologist, the smartphone to a college student, or the stethoscope to a family doctor: it is the professional’s tool (and should be clearly visible at all times).

FIGURE 1.27. Measuring the orientation of structural elements in outcrop. In this case, a lineated surface is measured using a Breithaupt® structural geology compass that measures dip and dip direction (the latter is perpendicular to strike).

Basic geometric principles, the ways of describing geometric features, and concepts related to constructions such as structure contours and spherical projections are explained in structural geology laboratory manuals, with only short descriptions of terms and associated concepts below, some of which are illustrated in Figure 1.28:

• Apparent dip - Dip of a plane in an imaginary vertical plane that is not perpendicular to the strike. The apparent dip is less than or equal to the true dip.

• Attitude - Orientation of a geometric element in space.

• Cross section - Plane perpendicular to the Earth’s surface.

• (True) dip - The slope of a surface; formally, the angle of a plane with the horizontal measured in an imaginary vertical plane that is perpendicular to the strike.

• Dip direction - Azimuth of the horizontal line that is perpendicular to the strike.

• Foliation - General term for a surface that occurs repeatedly in a body of rock (e.g., bedding, cleavage).

• Lineation - General term for a penetrative linear element, such as the intersection between bedding and cleavage or alignment of elongate grains

• Pitch or Rake - Angle between a linear element that lies in a given plane and the strike of that plane.

• Plunge - Angle of linear element with earth’s surface in imaginary vertical plane.

• Plunge direction - Azimuth of the plunge direction.

• Position - The geographic location of a geometric element (e.g., an outcrop).

• Profile plane - Plane perpendicular to a given geometric element; for example, the plane perpendicular to the hinge line of a fold.

• Strike - Azimuth of the horizontal line in a dipping plane or the intersection between a given plane and the horizontal surface (also trend).

• Trace - The line of intersection between two nonparallel surfaces.

• Trend - Azimuth of any feature in map view; sometimes used as synonym for strike

FIGURE 1.28. (a) Attitude of a plane: dip, dip direction, and strike. The strike of dipping plane ABCD is the intersection with an imaginary horizontal plane (line EF). The dip of plane ABCD is given by the angle α, while line GH represents the dip direction. Note that there are two possible directions of dip for a given strike. Thus, when using dip and strike, the general direction of dip must also be included (e.g., 090°/45° N). Alternatively, the attitude of plane ABCD, using dip direction and dip, is uniquely described by 000°/45°. (b) Attitude of a line: plunge, plunge direction, and pitch. Line EF lies in dipping plane ABCD. The plunge of line EF is the angle β, which is measured from the horizontal (line EG) in an imaginary vertical plane that contains both line EF and line EG; line EG is called the plunge direction. The attitude of line EF, using plunge and direction of plunge, is given by 30°/315°. The pitch (angle γ) is the angle that line EF makes with the horizontal (strike) of a plane (here: plane ABCD) containing the line. Note that when the pitch of a line is recorded, the attitude of the reference plane must be given, as well as the side from which the pitch angle is measured (here: 000°/45°, 40° W).

1.7.2 Guidelines for Structural Analysis

As we close this chapter, we want to offer a few general comments on structural analysis. Good scientific work requires that one separates observations from interpretations, which equally holds for geologic mapping. Yet, a geologic map, a cross section, or a block diagram without interpretation misses the unique insights of the investigator. If you spend the time collecting and digesting data, you are best suited to make the interpretations (or, educated guesses). The following suggestions may help with map interpretation, but note that they hardly do justice to the intricate process of structural analysis and interpretation; our aim with these guidelines is mostly to point you in the right direction:

• Strata are deposited horizontally. This is the Law of Original Horizontality, which makes bedding an internal reference frame.

• Strata follow one another in chronological, but not necessarily continuous, order. This is known as the Law of Superposition.

• Separated but aligned outcrops of the same lithologic sequence imply stratigraphic continuity.

• Strata occur in laterally continuous and parallel layers in a region.

• Sharp discontinuities in lithologic patterns are faults, unconformities, or intrusive contacts.

• Deformed areas can be subdivided into a number of regions that contain consistent structural attitudes (structural domains). For example, an area with folded strata can be subdivided into regions with relatively constant dip direction (or even dip), such as the limbs and hinge areas of large-scale folds.

• The simplest but internally consistent interpretation is most correct. This is also known as the “least-astonishment principle”.

The assumptions on which interpretations are based do not hold universally; in fact, after some field experience you may disagree with one or more of the points listed. In our experience, however, these guidelines enable a reasonable, first-order interpretation of the geometry of an area. Each individual guideline is valid under a given set of circumstances, but remember that, except for the laws, they remain mere assumptions; no more, no less. Whenever possible your assumptions should be tested by adding more observations, and when the assumptions continue to hold, only then may your interpretation be valid. This approach follows a proven scientific method, called the testable working hypothesis, which eventually leads to a model. If the model is very successful it may become a law. Until then, there is always room for alternative interpretations and models.

Increasingly, subsurface data from drilling and geophysical methods are available to a structural geologist that are used to test and constrain one’s interpretation. Drilling is restricted to the upper 10 km of the crust, but provides samples of deeply buried layers that can be compared with exposed rock units. This is a powerful test for the cross sections and block diagrams you construct. Using two-dimensional and three-dimensional deep seismic reflection imaging we get an indirect view of the deeper parts of the Earth. Seismic reflection profiles are obtained by recording the travel times of sound waves that bounce off layers in the Earth. Correlation of these reflectors with features that are exposed at the surface or obtained from drilling gives important information on the nature of the deep structure.

1.8 Deformation Regimes

We consider the basic geometric classes of deformation structures to be a manifestation of the mesoscopic cohesiveness of rock. Joints, veins, and certain types of faults are manifestations of primarily frictional processes, whereas cleavage, foliation, and folding are largely manifestations of plastic processes. Thus, in this book we subdivide our discussion of deformation structures into two parts that reflect the characteristic conditions: “The Frictional Regime”, representative of the upper crust, and “The Plastic Regime”, representative of the deeper crust and mantle.

Processes governed by friction are characteristic for deformation in the upper part of the crust, where temperatures and pressures are relatively low, exemplified by joints and faults. Plastic processes define deformation in the middle and lower part of the crust and the mantle, as they are favored under conditions of greater pressure and temperature, exemplified by folds, foliations and shear zones. We schematically illustrate this in a synoptic diagram of a vertical fault that cuts the upper and middle crust in Figure 1.29.

FIGURE 1.29. Synoptic diagram a displacement zone that cuts deep into the crust, showing the frictional and plastic regimes, the frictional–plastic transition, metamorphic facies (P-T) and crustal strength (represented by differential stress, σd). Characteristic rock types associated with displacement are also indicated. This is sometimes called the Sibson-Scholz fault model.

While generally correct, as with all models, this classification has oversimplifications. For example, a sudden increase in rate of deformation may cause rock that was is deforming by plastic processes, such as folding, to suddenly behave in a frictional manner (by fracturing). You can see this remarkable effect by, respectively, slow and quick pulling of a piece of SillyPutty®. However, you’ll hopefully find the deformation regime classification used here a robust and informative approach to examine deformation structures on all scales, and the associated processes.

1.9 Plate Tectonics

Ultimately, most crust and mantle structures are a consequence of plate tectonic activity and the actions of associated forces, which describes the slow (on the order of centimeters per year) but steady motion of segments of an outer, strong layer of the Earth, called the lithosphere, over a weaker part of the mantle, called the asthenosphere. These key terms and their characteristics will be defined in subsequent chapters. The forces of plate motion, especially the interactions at plate boundaries, produce the deformation structures we study in the field and in the laboratory that are the focus of this text.

FIGURE 1.30. The principal deformation settings of plate tectonics. Three types of plate boundaries arise from the relative movement (arrows) between lithospheric plates: C—Convergent (or sudbuction) boundary characterized by contractional structures (Chapter 13), D—Divergent (or rift) boundary characterized by extensional structures (Chapter 12), and L—Lateral slip (or transfer) boundary characterized by wrench structures (Chapter 14).

The three end-member types of plate motions are convergence, divergence and lateral slip (Figure 1.30), each of which are explored in the final block of chapters. Without the activity that arises from plate motion, including deformation, volcanism and earthquakes, the Earth would be as dead as the Moon. In other words, plate tectonics provides the dynamic framework to examine the occurrence and significance of deformation structures that are found on local, regional and global scales.

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v.3.0; last update: 14-Sep-2023

[1] Named after the Dutch sedimentologist Arnold Bouma (1932-2011).