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INTRODUCTION TO GEOLOGY LECTURE 1/2 TITLE 3/4 Plate Tectonics and Earth Structure Minerals: 5/6 The Building Blocks of Rocks Igneous Rocks: 7/8 Plutonism and Volcanism Earthquakes 9/10 Metamorphic Rocks 11/12 Sedimentation and Sedimentary Rocks 13/14 Sedimentary Environments 1 : Glaciers and Glaciations 15/16 Sedimentary Environments 2 : Deserts and Wind 17/18 Sedimentary Environments 3 : Fluvial and Groundwater 19/20 Sedimentary Environments 4 : Shoreline and Pelagic 21/22 Stratigraphy and Structure 23/24 Geological Mapping and Construction of Cross-Sections 25/26 The Formation and Occurrence of Petroleum 27/28 Petroleum Migration, Reservoirs and Traps 29/30 Petroleum Exploration Methods STUDENT NOTES The Introduction to Geology Lecture Notes are provided as a guide to the work covered in the lecture programme outlined above. They constitute a summary of the more comprehensive power point presentation already given to the trainees. HOW THE EARTH WORKS - PLATE TECTONICS The earth is a "geologically active" planet; its major surface features, continents, oceans and the great mountain ranges, owe their existence and ever changing forms to "tectonic" activity. The following notes provide an outline description of the Plate Tectonic model, in which a relatively simple system of moving lithospheric plates provides an explanation of the major surface features of our planet and the tectonic processes that produce them. STRUCTURE OF THE EARTH Figure 1.9 pg. 13 illustrates the structure of the earth, with a central solid iron core, surrounded by a liquid iron core, the lower mantle, and then the upper mantle, consisting of a partially molten, weak asthenosphere and a strong lithosphere with a surficial crust of light rock. About 90% of the Earth is made up of the four elements iron, oxygen, silicon and magnesium, which are the fundamental building blocks of most minerals. A large proportion of the iron, being heavier, sank to the core of the Earth, and lighter elements such as silicon, aluminium, calcium, potassium and sodium have become more concentrated in the crust. THE EARTH'S MAJOR TOPOGRAPHICAL FEATURES When viewed from space the major features of the earth's surface appear to be the continents and ocean basins. Mountain ranges occur as linear features, extending for thousands of kilometres in some continental areas. If, however, we imagine the waters of the oceans to be removed……. we should be aware of the great ocean ridge system and, on a rather smaller scale, the deep ocean trenches. The plot illustrating the distribution of topographic levels on the earth's surface shows two dominant levels: 1. continents (average height c.1 km) forming c. 40% of the earth's surface 2. oceans (average depth c. 4km) forming c. 60% of the earth's surface. The proportion of the earth's surface occupied by the extremes of height (>3km.) and depth (>5km) are very small, forming only about 1 % of the total. The two dominant surface levels reflect the two different types of crust (Figure 1.12 pg. 17): 1. continental crust which has a composition close to that of granite has an average thickness of 33km. 2. oceanic crust which has a composition close to that of gabbro or basalt has an average thickness of 7 km. Compared with oceanic crust, continental crust is thick and of lower density; this results in the continental areas being topographically higher than oceanic areas, according to isostatic principles, and hence the two dominant levels. Mountain ranges, ocean ridges and trenches form the major topographical anomalies: 1. mountain ranges; this measure c.300-800 km. across; the cordilleran belt of the western Americas and the Alpine-Himalayan belt are the major mountain belts of the continental areas. Another type of mountain belt forms the series of island arcs found mainly around the northern and western margins of the Pacific and along the north eastern margin of the Indian Ocean. The topographical relief of these island arc systems is as great, if not greater, than those of the continental ranges, but because they are partially submerged, they appear less significant. 2. ocean ridges; these are more significant volumetrically than the continental mountain ranges; typically measure 500-1000 km across. They occupy about 30% of the surface area of the oceans. Their huge mass and excess topographic relief is isostatically compensated by the underlying mantle being hotter and less dense than its equivalent beneath the deep sea plains. 3. ocean trenches; these form discontinuous features situated near continental margins, (e.g. western South America), or bordering the convex side of island arcs, (e.g. those of the North West Pacific). They are generally about 100-150 km. across and extend to depths 2-3 km. deeper than the average ocean depth. The deepest trenches are over 11 km. below sea level. PRESENT DAY TECTONIC ACTIVITY The obvious manifestations of present day tectonic activity are seismicity (earthquakes) and vulcanicity (volcanic activity). Study of the distribution of current tectonic activity reveals a close correlation with the topographical anomalies noted. 1. seismicity; the vast majority of earthquakes are concentrated in narrow, well defined belts, corresponding with the topographical anomalies. More than 80%of the total earthquake energy is released in the circum-Pacific belt. Thus it seems that the young mountain ranges, ocean ridges and the deep ocean trenches which represent extreme disturbances of the earth's relief are also the sites of current tectonic activity. Study of the distribution of earthquakes in terms of depth of focus reveals the following pattern: ocean ridges are associated with shallow focus earthquakes (generally<65km). These are concentrated along a central rift or along faults that offset the axial rift. deep earthquakes (>300km.) are associated with ocean trenches, especially those situated on the margins of the Pacific Ocean. A typical cross section of the earthquake activity associated with ocean trenches shows that the earthquake foci lie on an inclined plane, which outcrops at the site of the trench and dips below the adjacent island arc or continental margin. This inclined zone of earthquake activity is known as a Benioff zone. 2. vulcanicity; the distribution of present day volcanic activity is very similar to the pattern of earthquake distribution. This therefore implies a close relationship between vulcanicity and tectonic instability. 75% of currently and historically active volcanoes lie in the circumPacific belt, especially along the volcanic island arcs. In addition, there are many volcanoes associated with the ocean ridge system, and along certain major faults and lineaments in the ocean basins. The greatest concentration of vulcanicity on the continents is found in the African rift system. 3. stable and unstable tectonic zones; present day tectonic activity reveals a pattern of relatively narrow zones of tectonic instability corresponding with the major topographical anomalies; these are separated by broad areas in which there appears to be tectonic stability. When we look closely at continental areas, we find evidence of the existence of stable and unstable zones extending back 2,500 million years into geological time. The stable zones are known as cratons and the unstable zones as mobile belts. Continental masses it seems are made up of former mobile belts that have subsequently stabilized, and most continents contain the nowstabilized products of mobile belts that were active 3,800 million years ago. In contrast, the stable parts of the oceanic areas contain no evidence of instability older than 200 million years; the crust of the present day oceanic areas all appears to have formed within the last 200 million years and contrasts markedly with the great antiquity of some continental cratonic areas. THE PLATE TECTONIC MODEL The Plate Tectonic model offers very satisfactory account of how the major features of the lithosphere can be explained in terms of a relatively simple system of moving lithospheric plates. The model provides a unifying theme throughout the science of geology, and whilst regarded as "revolutionary" by some, the Plate Tectonic model has developed from earlier models of "Continental Drift" and "Seafloor Spreading". Plate Tectonics The central idea behind plate tectonics is that the lithosphere is divided into twelve large rigid plates, each moving as a distinct unit (Figure 20.4 pg. 204 plus back page). The plates consist of rigid lithosphere (with either a thick continental of thin oceanic crust), which 'float' on the partially molten asthenosphere. Convection currents within the asthenosphere are thought to be the driving force behind plate movement. For example, where hotter matter rises under the ocean, it flows apart and carries the plates along with it. When this hot matter has cooled and starts to sink, the plates sink also (Fig 1.19 pg. 20). The plates are therefore constantly moving, which explains why the Atlantic Ocean apparently did not exist 150Ma ago. At this time, it has been established that Eurasia, Africa, and the Americas were all one continent called Pangea. Continental Drift Ever since the development of accurate maps, several authors have commented upon the apparent "jig-saw" fit of continental coastlines that lie on opposite sides of major oceans. Some of these authors have argued that the jig-saw fit is not accidental, but reflects the break up of former super-continents and the subsequent separation and drifting apart of continental fragments. Until the 1950's, most physicists poured scorn upon the concept of continental drift, arguing that there was no satisfactory mechanism for producing the postulated drift. However, since the 1850's onwards, geologists have been collecting large amounts of data (some lithological, some palaeontological) apparently supportive of the continental drift hypothesis. Moreover, in 1928, a very famous British geologist (Arthur Holmes) suggested that slow moving convection currents within the earth's mantle could provide a viable mechanism by which continental drift could be accomplished. Arthur Holmes' mechanism did not appeal to sceptical scientists, and continental drift proponents only attained scientific respectability following certain developments in geophysics during the 1950's. In particular, the measurement of rock magnetism provided very strong evidence in support of continental drift. When an igneous rock (i.e. a rock derived from molten earth materials) crystallizes, any magnetic minerals developed during the crystallization process acquire a weak magnetic field that is aligned parallel to the earth's magnetic field at the time of crystallization. Measurement of the rock's magnetic field allows the magnetic pole position and the rock's magnetic latitude to be determined at the time of crystallization. If continental drift takes place, relative to a fixed magnetic pole, the magnetic fields of igneous rocks of different ages should record different magnetic pole positions for a given continent at different points in time. In fact, relative to a specified continent the apparently changing magnetic pole positions with time can be represented by a polar wandering curve. Polar wandering curves for Europe and North America Note that with the Atlantic closed, the curves are remarkably similar for both continents for Cambrian to Triassic Periods. The development of the Atlantic Ocean and the increasing separation of Europe and North America is responsible for the differences in the Jurassic to Recent polar wandering curves. These data imply continental break-up during the Trias, and Jurassic to Recent enlargement of the Atlantic Ocean as Europe and North America "drift" ever further apart. Sea Floor Spreading This model was developed by Hess during the early 1960's. Hess suggested that continental drift occurred by means of a sea floor spreading mechanism, driven by convection currents operating within the earth's mantle. He suggested that the ocean ridge system formed at the sites of upwelling convection currents; the ocean floors themselves took part in the convective circulation and moved symmetrically away from the ocean ridges like giant conveyor belts. New ocean floor was created at the ridge axes by consolidation of molten earth materials (magma and lavas) that "filled the gap" left as older parts of the ocean floor moved away from the ridge axis. As they moved away from the ridge axis, these older parts of the ocean floors cooled and subsided to the lower level of the deep sea plains. In this model the continents were passively carried along by the conveyor system; and since individual continents were known to have moved over 1,000 km in 100200 million years, the average rate of "drift" is of the order of 1-2 cm. per year. Old ocean floor, which had become cool and dense, sank back into the mantle convective circulation at the sites of the great ocean trenches. Thus in Hess's model, ocean floors are continually being created and destroyed (a cycle of activity taking about 200 million years, i.e. equivalent to the age of the oldest ocean floor). Continental materials however do not become dense enough to sink back into the mantle, they are effectively the "scum" on the upper surface of the convective circulation; they are not recycled within the mantle, hence the great age of certain parts of the continents. Magnetism and Magnetic Reversals Motions in the fluid iron core of the Earth set up a dynamo action thus generating the Earth's magnetic field (Figure 19.3 pg. 478). Rocks are magnetized in the direction of the magnetic field at the time of their formation. The rocks can be dated radiometrically and thus the history of the magnetic field recorded. Such studies have shown that the field reverses direction (the reason for which is unexplained) and such reversals are evident on the sea-floor. Figure 19.15, 19.16 and 19.17 (pgs. 485 – 486) illustrates the symmetrical pattern of magnetized rocks either side of a MOR. Supporting evidence for Hess's model was presented in 1963 by Vine and Matthews. They recognized the significance of linear patterns of magnetic anomalies (with an amplitude of c.1 % of the earth's field strength) observed on the ocean floors. These magnetic anomalies are caused by reversals in the polarity of the earth's magnetic field. Roughly every 500,000 years or so, the earth's magnetic field reverses its polarity, and thus rocks that develop magnetic minerals as they crystallize from a molten state within the earth's magnetic field will record either normal or reversed polarity. For example, rocks which crystallize at times when the earth's field is reversed, acquire a rock magnetic field that is opposed to the present magnetic field of the earth. Thus the total magnetic field strength recorded at the sites of such reversely magnetized rocks will be less than that recorded for normally magnetized rocks. If ocean ridges represent the sites where new ocean floor is continuously being created (as in Hess's model) by consolidation of molten earth material, then the ocean floors should be paved with strips of normally and reversely magnetized rocks as noted by Vine and Matthews. Since 1963, the symmetry of the magnetic anomaly strips with respect to the ocean ridge system has been demonstrated. Magnetic anomalies have now been mapped and correlated for the entire world's oceans; this allows patterns and rates of spreading to be identified and compared on a world-wide basis. Note particularly the offsets that commonly appear to displace segments of the anomaly pattern. These offsets are caused by structures known as transform faults. Transform faults were a predicted, and now verified consequence of the sea floor spreading mechanism. The geometry of such a fault is illustrated; note that the true sense of displacement is opposite to that of the ridge axis offset, and that seismic activity only occurs along the section between the ridge offsets. The Emergence of the Plate Tectonic Model The Plate tectonic model has evolved and developed from the fusion of ideas and evidence assembled with respect to the models of continental drift sea floor spreading. Consider the following observations: a. There is detailed and accurate fit of continents, e.g. between South America and Africa after 4,000 km of drift that has taken nearly 200 million years to accomplish. This testifies to a lack of distortion of the continents involved in this process. b. The ocean floors are paved with magnetic anomaly stripes. These have maintained their shape and continuity, in some cases for tens of millions of years throughout the spreading process. c. The patterns of tectonic activity indicate large stable, tectonically "quiet" areas, in both continental and oceanic domains, whilst between these stable areas, there are relatively narrow zones that are tectonically very active. Moreover these tectonically active zones occur in continuous belts around the earth's surface. It thus follows that since seismic activity is caused by the relative displacement of rock masses on either side of a fracture system, the continuous zones of seismic/tectonic activity mark the limits of large, relatively rigid and stable areas that are in a state of relative motion with respect to each other. The stable zones are now known as lithospheric plates (lithospheric because each plate is limited in depth to the base of the lithosphere). The tectonically active zones represent the sites of plate boundaries where adjacent plates inter-react in response to the convective circulation within the earth's mantle. PLATES AND PLATE BOUNDARIES The Major Plates The nature of the relative plate motions indicates three main types of plate interaction which generate three main types of plate boundary. Adjacent plates may: a. Diverge, giving what are known as constructive boundaries b. Converge, resulting in destructive boundaries c. Slide past each other, in which case the plates have transform boundaries. Note that plate boundaries are not necessarily coincident with continent/ocean margins. Continental margins that are also plate boundaries are said to be active (e.g. the west coast of South America), whereas continental margins that lie within the interior of a plate are said to be passive (e.g. the east coast of South America). 1. Constructive Boundaries: these form where lithospheric plates move apart (Figure 20.10 pg. 509). If a constructive boundary is initiated within a continental area, there will be stretching, thinning and rifting of the continental lithosphere leading ultimately to the break-up of the continental mass. These processes will be accompanied by igneous activity and thermal doming, with molten materials poured out on the earth's surface and intruded at depth. Divergent Plate Boundaries: Continental Rifting: Once the continent is ruptured, continuation of the plate divergence process results in the formation of a new ocean basin, with new oceanic lithosphere generated at the sites of the new ocean ridge (spreading ridge). Divergent Plate Boundaries: The Ocean Ridge: The new oceanic lithosphere is generated from a combination of processes that include the ductile inflow of asthenospheric material at depth, the intrusion and consolidation of magma within the lithosphere and the surface outpourings of lavas along the ridge axis. East African Rift, Red Sea, Gulf of Aden Map, Atlantic Ocean Map: Illustration of the process of continental rifting (East African Rift), break-up (Red Sea Rift) and the development of a young ocean (Gulf of Aden). Illustration of a mature ocean with a long well developed spreading ridge system and the large separation of continental masses that were formerly contiguous. 2. Destructive Boundaries: these are formed wherever lithospheric plates are converging and moving towards each other (Figure 20.12 pg. 513). Essentially there are two types of convergent boundary: a. Subduction zone; these occur at sites where oceanic lithosphere is returned to the mantle; the downgoing or subducting plate slides beneath the upper plate as convergence between the plates continues. Benioff zones (the inclined zones of seismicity, noted earlier) mark the sites where the two converging plates are in contact. Convergent Plate Boundaries; Oceanic-Oceanic Convergent Plate Boundaries; Continental-Oceanic Illustrations of subduction zones: In both cases an oceanic trench marks the sites where the subducting plate begins its descent towards the mantle. At greater depths (100-200km.) portions of the downgoing slab start to melt; the melts produced are relatively light and buoyant; they rise to higher levels and are responsible for the intrusive and volcanic activity associated with island arcs and cordilleran mobile belts. Island arcs are formed in those situations where the overriding plate consists of oceanic lithosphere (e.g. Japanese Islands), whereas cordilleran mobile belts form where the upper plate is composed of continental lithosphere (e.g. South American Andes). The main differences between these two types of subduction zone lies in the proportions and compositions of the molten products. At depths of 650-700 km Benioff zones cannot be recognized, from which observation it can be inferred that, seismically, the downgoing plate ceases to be distinct from surrounding mantle material. Pacific Ocean Map: The present day Pacific Ocean, largest of the world's oceans is ringed by subduction zones, and is thus an ocean that has entered its declining phases of activity, to the extent that destruction of the oceanic portions of its constituent lithospheric plates is proceeding space. Ultimately most oceans reach a state where the rate of plate destruction via subduction zones, exceeds the rate of production of new oceanic lithosphere at their spreading ridges. This results in their gradual contraction and ultimate closure. Continents on opposing sides of the ocean converge and eventually collide with each other to produce a: b. Collision Zone; Continental lithosphere is generally too light to sink back into the mantle i.e. it is not subducted. Plate convergence that results in continental collision and the closure of the intervening ocean usually continues for some time after the initial impact of the continental masses. This results in the leading edges of the colliding continents becoming strongly deformed (earthquakes etc.) and thickened. causing the formation of continental mountain beltsThe Himalayas, Tectonic Elements and Sections: The present day Himalayas represent the results of continental collision between the Asian continent and continental portions of the Indian Plate. The initial impact between these continents took place c. 50 million years ago as the former Tethys Ocean closed. Since then it has been calculated that there has been at least a further 1500 km. of continued convergence that has caused intense deformation of the impacted continental crust, such that Tibet now has continental crust that is approximately twice the normal thickness. Further west along the Alpine-Himalayan belt the present day Mediterranean Sea represents a mere remnant of the formerly extensive Tethys Ocean; this ocean will finally close when the African and European continents eventually impact each other. Following continental collision, all that usually remains to mark the presence of the former ocean is a narrow zone (known as a suture) in which fragments of oceanic lithosphere are preserved, e.g. the Indus Suture marks the boundary between Asian and Indian continents and the site of the former Tethys Ocean. 3. Transform or Conservative Boundaries: these are plate boundaries defined by transform faults/fault systems, where plates neither converge nor diverge. Plate movement vectors are thus parallel to transform fault traces, and the rates of motion on transform faults are determined ultimately by rates of sea floor spreading. Both convergent and divergent plate boundaries are segmented by transform faults. Plate map of the world: Transform faults effectively segment convergent and divergent boundaries so as to allow different rates of convergence or divergence in different parts of the system (a necessary requirement when moving "rigid" plates on the surface of a sphere, where tangential velocities vary between the poles and the equator of the relative motion). One of the best known and most studied examples of a transform boundary is the San Andreas Transform between the Pacific and the American Plates. This is in effect a 1000 km offset of the East Pacific Rise spreading ridge and its northward continuation, the Juan de Fuca ridge. THE WILSON CYCLE Perhaps one of the most surprising features of planet earth is the relatively great age of much of the continental areas, compared with the relative youth of the oceanic areas. All the oceanic lithosphere that we see today has been produced within the last 200 million years. Plate tectonics provides a very satisfactory explanation for this, and indeed many other geological phenomena. Ocean basins, it seems, undergo a cycle of renewal and destruction over a period lasting 200 million years or so, whereas continental lithosphere, once created, is essentially permanent; it may undergo reworking, but it always essentially remains as continental lithosphere. The cycle of renewal and destruction of ocean basins has become known as the Wilson Cycle in honor of a famous Canadian geologist, Tuzo Wilson. In general terms, the cycle begins with the birth of a new ocean basin from a continental rift; this gradually develops into a broad mature ocean which eventually closes as subduction on its margins outpaces spreading at its ridges. Ultimate closure leads to the collision of continents on opposing margins of the ocean, leaving a suture between continental masses as the only relict of that former ocean's lithosphere. Supplementary Reading/Information: Chapters 1, 19 and 20 - Earth MINERALS: THE BUILDING BLOCKS OF ROCKS The Earth is host to three main rock types:(see Figure 3.1, pg 53) Igneous – formed from hot molten liquid (magma) sourced from within the Earth (Intrusives and Extrusives) Metamorphic – igneous or sedimentary rocks altered by action of heat and pressure e.g. during burial Sedimentary – igneous or metamorphic rocks eroded and transported away from their source e.g. by water). Sediments are re-deposited in layers in lowland areas. Minerals are the building blocks of these rocks. About 90% of the Earth is made up of the elements iron, oxygen, silicon and magnesium, which are the fundamental building blocks of most minerals. Rocks are described as aggregates (a mixture) of minerals, although certain minerals can occur by themselves in large, impure quantities e.g. Calcite is the dominant constituent of the rock limestone. A mineral is defined as a naturally occurring inorganic solid that possesses a definitive chemical structure which gives it a unique set of physical properties. A mineral must exhibit the following characteristics: Naturally occurring Inorganic Solid Definite chemical structure BONDING ATOMS There are almost 4,000 recognized minerals on Earth. The elements which combine to form an individual mineral are held together by electrons in the outer shell of the atoms. The electrons involved in bonding are called valence electrons, and the number of these available determines the number of bonds it will form e.g. silicon has 4 valence electrons and thus forms 4 bonds, oxygen forms only 2 bonds and hydrogen only 1. There are a number of types of bonds, defined by the behavior of the valence electrons: Ionic Bonds: One or more valence electrons are transferred from one atom to another. One atom becomes stable by giving up its valence electrons, and the other makes itself stable by using them to complete its outer shell. The loss or gain of an electron results in a net positive or negative charge respectively. Atoms with a charge imbalance are called ions, and as unlike charges attract, ions attract one another to form a neutral chemical compound. Ionic compounds consist of an orderly arrangement of oppositely charged ions assembled in a definite ratio that provides overall electrical neutrality e.g. sodium chloride. Thus, the sodium atom becomes a positively-charged ion which attracts the negatively charged chlorine ion. Covalent Bonds: Some atoms combine together by sharing electrons in order to acquire a full outer shell. For example, in chlorine gas (Cl2), each chlorine atom shares an electron in its outer shell, which has 7 electrons, in order to achieve a more stable arrangement. This is known as a covalent bond. Both ionic and covalent bonds may occur within the same compound. For example, in silicate minerals, silicon is bonded to oxygen covalently, to form the basic building block common to all silicates. These are then ionically bonded to metallic ions, producing various electrically neutral chemical compounds. MINERAL GROUPS Despite there being nearly 4,000 minerals, composed of various combinations of the 100 elements, 98% (by weight) of the Earth's crust is made up on only a few dozen abundant minerals, with only eight elements composing the bulk of these minerals. The two most abundant elements are silicon and oxygen, which combine to form the framework of the most common mineral group. Silicates: All silicates have the same fundamental building block, the silicon-oxygen tetrahedron. The tetrahedral consist of one silicon atom bound covalently to four oxygen atoms. This however leaves an imbalance, as the four oxygen have a total charge of –8 and the silicon +4. Thus the tetrahedron is a negatively charged ion (SiO44-) which achieves stability by bonding to positively charged ions. The tetrahedral are bound together by ions such as A13+ , Fe3+, Mg2+, Fe2+ , Ca2+, K+, Na+. The tetrahedral can also bind together by sharing of oxygen atoms between adjacent silicon atoms. Hence, the tetrahedral may form single chains, double chains or sheet structures. Due to the sharing of oxygen atoms, the ratio of oxygen: silicon differs in each of the silicate structures, i.e. Oxygen: Silicon Ratio Isolated Tetrahedron 4:1 Single Chain 3:1 Sheet silicates (3-D) 2:1 As more oxygen is shared, the percentage of silicon increases, and silicate minerals are therefore described as having high or low silicon content. Most silicate structures carry an ionic charge and are neutralized by the inclusion of charged metallic ions that bond them together into a variety of crystal configurations. Metallic ions of a similar atomic size can substitute in for one another e.g. Fe2+ and Mg2+ are a similar size, as are Ca2+ and Na+. Due to this substitution, an individual mineral may contain varying amounts of certain elements. For example, olivine (Mg,Fe)2SiO4 may contain varying proportions of iron and magnesium. Thus olivine is actually a family of minerals with a range of compositions (known as a solid solution) between two end member compositions. When Ca2+ substitutes for Na+, the structure gains a positive charge. To maintain neutrality, A13+ can substitute for Si4+. This double substitution is common in the mineral plagioclase feldspar, the end members of this family being anorthite (CaAl2Si2O8) and albite (NaAlSi3O8). There are a number of major groups of silicate minerals, most of which form by crystallization from a cooling magma. Each group has a particular silicate structure, and displays a characteristic cleavage. Silicate minerals tend to cleave between the silicon-oxygen tetrahedral rather than across them, due to their strong covalent bonds. Table 1 shows the main silicate mineral groups, formulae, cleavage and structure. Supplementary Reading/Information Chapter 3 - Earth IGNEOUS ROCKS: PLUTONISM AND VOLCANISM This lecture will address the origin of igneous rocks, and the reasons for their variation in composition and texture. The two basic categories of igneous rock are: Plutonic rocks: Intrusive igneous bodies where the magma has crystallized at depth (i.e. high pressure and temperature). Cooling of the magma is slow and hence plutonic rocks tend to be coarse-grained e.g. granite, gabbro. The pluton may be of any shape or size, depending on the volume of magma and the manner of emplacement. Volcanic rocks: Extrusive igneous rocks, where the magma crystallizes at the Earth's surface. It therefore cools very quickly and produces fine-grained rocks e.g. basalt, rhyolite. Texture: Igneous rocks typically show a mosaic of interlocking crystals of minerals such as quartz, feldspar, mica etc. The crystals usually show good crystal form with sharp faces, unlike sedimentary or metamorphic rocks where the crystals have been altered or weathered into more rounded shapes. Igneous rock classification: Plagioclase is a major mineral component of igneous rocks, and igneous rocks are in fact classified according to the relative proportions of sodium plagioclase (albite, NaAlSi3O8) and calcium plagioclase (anorthite, CaAl2Si2O8). The change in composition of plagioclase is basic to classifying both intrusive and extrusive igneous rocks. Figure 15.3 pg. 382 illustrates the names assigned to intrusive (plutonic) and extrusive (volcanic) igneous rocks according to the plagioclase composition. Figure 3.17 pg. 73 illustrates igneous rock classification in relation to the proportions of the various igneous minerals present. Rhyolite is the fine-grained equivalent of granite, and being rich in silica, these are classed as felsic igneous rocks. Basalt is the fine-grained equivalent of gabbro, and having a lower proportion of silica, these are classed as mafic igneous rocks. Peridotite and ultramafic rocks are rich in mafic minerals such as olivine, pyroxene, amphibole and black mica (biotite). Colour: Felsic igneous rocks tent to be lighter in colour than basic igneous rocks due to their higher proportion of silica. Felsic rocks may be white to light grey in colour, becoming progressively darker grey to black the more mafic the composition becomes. PLUTONISM MAGMATIC DIFFERENTIATION (FRACTIONATION) Magma in the Earth's mantle is taken as the starting material from which igneous rocks crystallize. This magma is typically mafic (basaltic) in composition. By some mechanism, this magma must differentiate to allow the production of magmas of different compositions, which then crystallize to give the range of igneous rock compositions which exist. This mechanism is called FRACTIONAL CRYSTALLISATION i.e. crystallization in which crystals does not react continuously with the melt. See Figure 15.6 pg. 385 and Figure 15.8 pg, 387. This may occur in a number of ways: Partial Melting: Some minerals melt earlier than others, for example albite and micas. Thus the composition of the liquid in a partially melted rock will differ significantly from that if the rock were completely melted. If this melt is removed then it can recrystallized to form a rock of very different composition from the starting material. Crystal Zoning: For example, plagioclase feldspar is zoned, as anorthite crystallizes first and albite later (Figure 15.5 pg. 384). The feldspar crystals therefore have Ca-rich cores and Narich rims. If, during crystallization, the crystals already formed, were removed, then the remaining fluid would be rich in albite. The next crystals to form from this fluid would therefore be more albitic, thus producing a rock of more felsic composition. Crystal Settling: There is evidence that within a plutonic body e.g. a magma chamber, early-formed crystals begin to settle in the bottom of the chamber, with later-forming crystals settling later, thus forming a layered structure. A classic example of this is illustrated by the Palisades cliff, on the Hudson River west band, New York (Figure 15.17 pg. 386). A basaltic intrusion was emplaced into sandstone and the upper and lower contacts cooled to form finegrained basalt. The mineral olivine crystallized first and sank to the bottom of the intrusion, followed by pyroxene and then plagioclase feldspar. However, the extent to which crystal settling is an important magma fractionation process is questionable as it requires an infinite amount of time for small crystals to settle in a viscous magma. Other mechanisms must be in operation to enable huge bodies of granite to exist. Other Fractionation Mechanisms: Within a magma chamber, convective motion occurs, allowing layers of crystals to be deposited on the walls and ceiling. Thermal variation within a chamber causes diffusion of ions, and thus concentration of elements and the creation of chemical zones in the chamber. The oxygen concentration of the magma in different parts of the chamber will also affect the course of crystallization. A magma chamber may contain two melts which are immiscible, and will thus each produce their own crystallization products. In contrast, melts from two different magma chambers may, on rising to the surface, meet and mix, thus crystallizing to produce a rock of different composition. Although magma in the mantle is essentially basaltic in composition, variation in the source can occur, particularly in subduction zones, when combinations of igneous, metamorphic and sedimentary rocks are assimilated into the mantle. This can lead to the formation of large granitic bodies in subduction zones. Where magma rises through continental crust, for example above subduction zones, there is more potential for fractionation of the magma to occur. As the magma rises through continental crust, rocks can be stopped (i.e. broken off) and melted into the magma, thus changing its composition. The magma will become progressively more felsic as fractionation processes occur. In contrast, magmas produced directly from the mantle with no fractionation are basaltic. These rocks are dominant in mid-ocean ridge environments, and for example in Iceland, which is situated directly on the Mid-Atlantic ridge. Igneous Rock Families: Fractionation processes allow the formation of igneous rocks which fall into three basic families, each associated with a certain geological setting (Figure 15.4, pg. 392): Calc-alkaline: Plagioclase; K-feldspar; Quartz; Mica; Amphibole; Pyroxene. Characteristic of plate convergence and subduction zones. Mafic/Ultramafic: Calcic plagioclase; Pyroxene; Olivine. Characteristic of mid-ocean ridge are, as an oceanic lithosphere. Alkaline: Na & K Feldspar; Feldspathoids; Biotite; No Quartz. These rocks are less abundant, but form along continental rifts and intraplate regions. FORMS OF MAGMATIC INTRUSIONS Magmatic intrusions can be of varied size and shape. Figure 15.14, pg. 393 shows the different forms of intrusion, and the list below describes each term. Pluton: a large igneous body congealed from magma underground. Country Rock: the invaded rock surrounding igneous intrusions. Sill: Tabular pluton where the magma is injected between beds of layered rock concordantly (i.e. parallel to the rock layers) Laccolith: Similar to a sill but mushroom shaped, not tabular. The overlying rock layers are domed upwards. Dike: Tabular pluton that cuts across the layering of the country rock i.e. discordant. Ring Dike: The erosional remnant of an intrusion that filled a cylindrical feature. Lopolith: A large usually concordant intrusive whose centre has sagged downwards to form a bowl-shaped body. Batholith: These are the largest of plutons, discordant intrusives at least 100km2. Stock: Similar to a batholith but smaller. VOLCANISM The distribution of volcanoes on Earth correlates strongly with plate boundaries. Magma erupts from volcanoes as lava, differing from the parent magma in that it has lost some of the volatile constituents. Also characteristic of volcanism are pyroclastic deposits i.e. volcanic rock fragments ejected into the air. The different types of lava and pyroclastics are described below: LAVA Pahoehoe: this is highly fluid lava which spreads in sheets. The elastic skin is dragged to produce a ropy texture (Figure 16.5 pg. 403). Aa: this is very viscous, slow moving lava whose thick skin is broken into a very rough jagged surface. Pillow lava: piles of ellipsoidal sack-like blocks which form during underwater eruptions. Tongues of lava cool quickly on contact with water and crack radially (Figure 16.6 pg. 403). PYROCLASTICS Particle size Rock formed on cementation Dust (Very fine) Ash (>2mm) Bombs (>6mm) Volcanic tuffs LITHIFICATION Volcanic breccias When pyroclastic fragments cement together (lithification), fine material such as dust and ash forms volcanic tuff, and coarse material forms volcanic breccia. Volcanic eruptions can also produce glowing clouds of hot ash, dust and gases which flow down the side of the volcano. This cloud is known as a nuѐe ardente (Figure 16.12 pg 406). It leaves poorly sorted, non-bedded deposits, which on compaction are called welded tuffs or ignimbrites. Eruptions involving pyroclastic deposits are known as phreatic i.e. large volumes of gas and steam. TYPES OF VOLCANO AND ERUPTIONS Hawaiian Volcano: These eruptions are slow and steady, non-violent, with very fluid basaltic lava. Gases escape readily and therefore pressure does not build up and the volcano is prevented from blowing its top. Strombolian Volcano: The lava is also basaltic, but more viscous, allowing pressure to build up and small explosions to occur every few minutes. The lava does not flow very far from the volcano centre. The Strombolian eruption is usually loud but not dangerous. It is named from a volcano on the island of Stromboli between Italy and Sicily. Vulcanian Volcano: This is named after Vulcano, a peak in northern Sicily. It is active only intermittently, each eruption lasting up to months, the volcano blowing its top with great force. A large volume of material is blown out of the crater, producing clouds of ash and gas, often followed by a lava flow. The dark clouds of ash rise into the stratosphere where the particles remain for years, altering weather around the entire earth. Vesuvian Eruption: Named after Mt Vesuvius near Naples. Ash and pyroclastic deposits are ejected vertically. Plinian Eruption: This is the most violent eruption, the force and volume of materiel ejected being extremely large. Peleean Eruption: This is a variation of the Plinian type but includes a pyroclastic flow which destroys everything in its path. Named after Mt Pelee which in 1902 destroyed a city on the island of Martinique. Mt St Helens is a combination of Plinian and Peleean, therefore being the deadliest volcano. ANATOMY OF A VOLCANO The extrusive material issue from a central vent or pipe (Figure 16.21 pg 410), and gives rise to a volcanic cone. Basaltic, fluid lavas produce volcanoes with gentle slopes known as shield volcanoes (Figure 16.22 pg. 411). More felsic, viscous lavas barely flow and produce volcanic cones. Pyroclastic material ejected from a vent produces cinder cones (Figure 16.25 pg. 412). When a volcano emits lava and pyroclasts, a composite cone or stratovolcano (Figure 16.27 pg. 414), built of alternating layers of lava and pyroclastic beds is formed. At the summit of most volcanoes, above the vent, is a crater. Calderas are large basin-shaped depressions which form due to collapse after a violent eruption of large volumes of magma (Figure 16.29 and 16.30 pgs. 415 and 416). Supplementary Reading/Information Chapters 15 and 16 - Earth METAMORPHIC ROCKS The term metamorphism is derived from a Greek word meaning change. It is a solid state process whereby the mineralogical and/or structural state of a rock is adjusted to changed conditions, usually of pressure and temperature, within the earth's crust. Metamorphic processes occur between the fields of igneous and sedimentary processes, usually in the temperature range 300-700°C, and in the pressure range 0-15kb (ca. 0-45km thickness of crust). Two types of pressure exist - hydrostatic (or confining) pressure due to the weight of the overlying rocks, and directed pressure resulting from earth's movements. Usually the confining pressure is greater, but the directed pressure affects the texture/appearance of the metamorphic rock. TYPES OF METAMORPHISM Thermal or Contact Metamorphism This occurs where temperature is high, for example adjacent to an igneous intrusion. The 'country rocks' surrounding the magma become heated and their texture and mineralogy may be changed i.e. metamorphosed. The margin of altered rock surrounding the intrusion often appears bleached and baked, and is called an aureole. The width and nature of the aureole depend on the nature of the country rocks, the size and temperature of the intrusion and its depth in the crust (i.e. the temperature of the wall rock), and also on the availability of fluids to transfer heat. Due to conduction, the temperature drops rapidly away from the pluton and therefore sequential zones of different grade metamorphism occur (i.e. high grade close to intrusion and low grade further away). This may be on a scale of cm to km depending on the temperature, and is reflected by variations in mineralogy and texture. The most spectacular aureoles may be expected where hot intrusions (i.e. 1000°C) are intruded at shallow crustal levels. Figure 17.18 pg. 444 shows an example of a skarn, which are banded rocks produced by contact metamorphism of limestone and dolomite. Figure 17.19 pg. 445 illustrates the sequence of minerals formed on contact metamorphism of sandstones I shales. Minerals containing volatiles are present in the outer zones, and dry, gas-free minerals in inner zones. Rocks of the aureole will often appear spotted. These are new metamorphic minerals that have grown. The rocks also become baked, hardened and indurated, often all traces of cleavage being lost. These hard, tough rocks produced are called hornfelses. Dynamic or Dislocation Metamorphism In this type of metamorphism the dominant variable is pressure not temperature. Dynamic metamorphism is found in major fault zones within the earth's crust. Along fault planes, the rock may be mechanically ground and broken up by deformational pressures, and the broken rocks produced are called cataclastics. They are the product of dynamic or syntectonic metamorphism. When this type of metamorphism occurs at depth, the deformation may be more ductile, leading to the production of fine grained rocks with a well developed layered structure, known as mylonites. As an indicator of the possible intensity of the mylonitisation process, there is evidence that some pebbles and fossils are stretched to 50-300 times their original size. Regional Metamorphism Contact and dislocation metamorphism tend to occur on a localized scale i.e. near to intrusions and fault zones. In the large P/T region between the fields of thermal and dislocation metamorphism is the domain of regional metamorphism (also known as dynamothermal metamorphism). Regional metamorphism tends to be characteristic of mobile or orogenic (mountain) belts, or portions of the earth's crust which become mobilized due to plate interactions. Geologists use minerals as gauges of P and T, as different minerals form under different P and T conditions. This is very important when mapping regional metamorphism. By mapping index minerals in the field, geologists can define broad zones or belts of metamorphism, ranging from least to most intense. Lines on a map connecting points on a map where an index mineral first appears are called isograds, signifying a change in metamorphic grade. Mineral assemblages (the presence of 2 or 3 minerals) are a more precise guide to the conditions of P and T experienced by the rock. Assemblages formed below about 300°C are referred to as very low grade, 300-500°C as low grade, 500-600°C as medium grade and over 600°C as high grade, merging into rocks that have been partially melted. Figure 17.12 pg. 440 gives an impression of how mineral composition of a metamorphic rock changes with increasing metamorphic grade. The sequence of minerals is not the same in all metamorphic environments, as P and T may not vary at the same rate as metamorphism becomes more intense. Note that the sequence of minerals formed is strongly dependent on the original rock type e.g. shale or basalt. Metamorphic rocks formed at high T and P may later be subjected to another set of metamorphic conditions. If the second event is of lower T and P to the first, then the rocks will be lowered in grade. This process is called retrograde metamorphism. TEXTURAL CHANGES IN METAMORPHISM Contact Metamorphism Pure sandstones and limestones will recrystallized to give rocks with an interlocking mosaic of equant crystals called quartzite and marble respectively. These rocks tend to be massive i.e. with no foliation or texture. As a rock is heated, it aims to minimize the ratio of surface to volume to become more stable. This leads to growth of larger crystals of equant size and thus a homogeneous rock (typically 1-2mm diameter). New minerals may grow into very large crystals surrounded by a finer matrix. These are porphyroblasts or metacrysts. Dislocation Metamorphism The main characteristic of this metamorphism is broken or strained crystals (cataclasis). At higher pressure, as already mentioned, mylonite forms. This is a rock flour, smeared out and flattened into laminae. Mylonites often contain porphyroclasts, which are larger fragments of the original rock which are surrounded by a finer grained matrix. These porphyroclasts show much evidence of strain such as bent cleavage, bent twins, strain twinning, broken grain boundaries etc. Regional Metamorphism This usually produces foliated rocks. Foliation is a set of parallel planes cutting the rock at an angle to the bedding of the original sediment. For example, shale metamorphoses to a slate (Figure 17.4 pg. 435). A slate shows a good fracture cleavage (or foliation) which is not shown by shales which instead part easily along their bedding planes. This fracture cleavage is at an angle to the bedding plane and allows cleavage I breakage into thin sheets at regular intervals. This perfect cleavage makes slate good for roofing tiles or flagstones. Good foliation is common in mica-bearing slates and schists formed from shales. Crystals or minerals tend to show preferred orientation along the foliation plane, as these planes often act as small shear planes. When crystals become aligned in a parallel manner it is known as lineation (Figure 17.7 pg. 436). Good lineation is common in mafic rocks which metamorphose to form large numbers of elongate amphibole crystals. Foliation and lineation are the product of preferred crystal orientation i.e. the tendency of crystals towards parallel alignment, with their shortest dimension perpendicular to the major compressional force (Figure 17.8 pg. 437). Schist: If metamorphic minerals grow to a larger size and become visible, they often become coarsely foliated, with segregation of minerals into lighter and darker layers. Gneiss: This is the extreme of schistosity. Light and dark minerals become segregated into coarse bands, with no tendency to split or part along these bands. NOMENCLATURE The names assigned to different metamorphic rocks are illustrated in Figure 3.19 pg. 76. Zeolite: The lowest grade rocks contain a variety of zeolite minerals. These are complex hydrous aluminosilicates formed by the alteration of mafic volcanic rocks. Greenschist: Slightly higher grade, greenschists contain chlorite (a sheet silicate), epidote (an aluminosilicate), actinolite (an amphibole) and albite (sodium plagioclase feldspar). These minerals all contain much iron, magnesium and calcium. Amphibolite: Higher grade again, these are characterized by hornblende (an amphibole), sodium calcium plagioclase feldspar, and garnet. Pyroxene Granulites: These are the highest metamorphic grade of mafic volcanics, similar in composition to a gabbro or basalt. Blueschists: These form in metamorphic belts where P is high and T is low. The characteristic mineral is glaucophane (a blue amphibole). Eclogites: These form at high P and variable moderate to high T. They are rich in pyroxene and garnet. Paired Metamorphic Belts: In subduction zones, cold subductifl9 slab with sediments on top sinks rapidly, reaching high P though the slab is still relatively cold. Blueschist is produced on the oceanic side of the subduction zone. However, on the landward side, igneous rocks generated be melting of the subducting plate rise to the surface, and the increased T transforms shallow buried volcanics and sediments. A paired metamorphic belt is produced with high Plow T rocks on the oceanic side and low P, high T rocks on the landward side. Metasomatism: This is when metamorphism results in changes in bulk chemical composition of the rock. This means that some chemical components have been transported by fluid e.g. hydrothermal fluids associated with a magmatic body may convect around permeable rock. Elements such as silica, sodium and potassium are highly soluble in hot aqueous fluids, and may be removed. However, much regional and contact metamorphism is isochemical, which means there is little change except for loss of water and carbon dioxide. Supplementary Reading/Information: Chapter 17 - Earth SEDIMENTATION AND SEDIMENTARY ROCKS Sedimentation is the final stage of the process beginning with erosion and transportation of eroded materials to sites of deposition. Particles settle out of suspension and are deposited in a layer. Physical sedimentation is where air and water currents transport solid materials to lowland areas. Chemical sedimentation is dominantly the process where sea water or other bodies of saline water precipitate dissolved substances in order to keep a constant composition. Diagenesis is the name given to the chemical and physical changes that occur after deposition, to alter composition and texture, and thus convert soft sediment to rock i.e. to lithify it. Diagenesis involves a range of processes. On compaction, water is driven out, e.g. a mud with 60% water can be compacted to mudstone with 10% water. Unstable minerals may recrystallize, and the growth of clay minerals is favored. Oil, gas and coal form as the result of diagenesis of the original sedimentary organic matter. Physical and chemical sedimentation follow a general downhill trend in response to gravity i.e. erosion begins in mountains/slopes and material proceeds to rivers and eventually the sea. The depositional patterns of sediments are strongly influenced by tectonics and the resulting geomorphic environment. Geologists analyze the sediments in order to decipher the paleogeography of the environment at the time of deposition. CLASTIC SEDIMENTS These are made up of weathered particles or detritus, e.g. shales, sandstones and conglomerates. Clastic (or detrital) sediments account for 3/4 of the Earth's sediments due to the dominance of mechanical erosion. Shale is three times more common than any of the coarser clastics. Sandstone: Sandstones are classified on the basis of their grain size. If the grains of sandstone are all of a similar size, it is well sorted. If there is a large range in sizes it is poorly sorted. Sorting is related to the type of depositing current e.g. beach sand is well sorted whereas debris-flow sand is poorly sorted. Classification Grain Size (mm) Very coarse sand 1.0 – 2.0 Coarse sand 0.5 – 1.0 Medium sand 0.25 – 0.5 Fine sand 0.125 – 0.25 Very fine sand 0.0625 – 0.125 Grains of sand are eroded during transportation and therefore become more rounded as they become more distal to their source. Sedimentary structures are the internal structures of sedimentary rocks and are very useful in reconstructing the sedimentary environment. For example, bedding is the planar surface, originally horizontal, on which sediment was deposited. Cross-bedding can occur due to ripples or current on the sediment surface and the direction of cross bedding indicates the direction of the current. Mud cracks indicate periodic drying out, and if a sediment becomes rapidly compressed e.g. by seismic shaking, dewatering structures can occur. From these sedimentary structures, it is possible to construct a paleocurrent map, showing the directions of sediment transport. The mineralogy of sandstone allows it to be traced back to its source. Quartz arenites contain almost entirely quartz grains. Arkoses contain abundant feldspar. Lithic arenites contain lots of fine-grained rock fragments from shales, slates, schists or volcanics. Graywackes consist of quartz and feldspar grains surrounded by a finer clay matrix. Gravel and Conglomerate: These contain large pebbles, and must have been deposited by stronger currents e.g. mountain rivers. A size of 25 cm diameter is approximately the limit that any river can carry. Pebbles become more abraded; rounder and smaller the further they have been transported. The pebbles may also become aligned such that they point in the direction of current flow. Conglomerates form in higher energy environments, for example during storms or on talus slopes. These are the slopes at the foot of continental margins, coral reefs or mountains, where boulders and debris accumulate. Mud and Shale: These are the most abundant sediments on Earth, but due to the fine grain size, they reveal least about their formation. The material is usually studied by electron microscopes and X-ray diffraction. They are defined as sediments with a large component of clay-size material (<1/256mm). Muds and shales are the result of slow settling from a very gentle transporting current. Below the depth of wave transport, muds and shales are constantly being formed on the ocean floor, blanketing ridges, continental shelves, trenches etc. Muds contain the remains of the decay of organisms and are therefore attractive to other organisms as a food source e.g. worms, burrowing clams, crustaceans etc. eat sediment, digest the organic matter and excrete the unused inorganic bulk. This leaves tracks, burrows and trails. This reworking and modification of sediments by organisms is known as bioturbation. Black shales contain abundant organic matter, having formed in a poorly oxygenated environment in which organic matter has not had chance to decay. On burial, this organic matter may alter to form oil and gas. CLASTIC SEDIMENTARY ENVIRONMENTS Figure 12.1 pg. 301 illustrates some of the environments in which sands can be deposited. Alluvial: This environment includes river channels, meander belts on flood-plains, alluvial fans and alluvial plains. As the channel migrates, it leaves behind a distinctive sedimentary sequence, with coarse sand and gravel on the channel floor, grading into fine sand, silts and muds on the flood plain at the top. This is known as the fining-upward alluvial cycle (Figure 12.10 pg. 310). Desert: The desert environment is dry enough to allow sand to be blown by the wind (eolian sedimentation). The dunes consist of fine, well sorted sand grains, with characteristic patterns of cross-bedding, indicating wind direction. Dune deposits grade into alluvial deposits of desert rivers. Glacial: Glacial environments include alluvial environments in front of the ice, an eolian environment where glacial rock flour is transported by strong off-glacier winds and deposited as loess, and the glaciomarine environment where glaciers calve icebergs in the sea. Under the ice, the deposits are known as tills, and are heterogeneous and poorly sorted. Glacial environments can be recognized by the presence of striated bedrock (scraped by rocks in overlying ice), and eskers (under-ice streams). Deltaic: The delta environment is complex, but acts as a major dropping point for river sediments. This environment is usually characterized by the stratigraphic pattern of alluvial freshwater deposits and fossiliferous marine deposits. Coarsening upwards of sediments may be evident, developed as the river mouth advances, depositing coarser sands of the channel over finer silts and muds offshore (Figure 12.12 pg. 311). Beach and Bar: Beach sands are well sorted and rounded, with bedding gently inclined towards the sea, and oscillation ripples in the surf zone. Typical of this environment would be a fine-grained subtidal sediment overlain by medium- to coarse-grained tidal zone sand deposits, then beach sands and topped by dune sands or salt-marsh organic-rich muds. Shallow Marine: Sedimentation on continental shelves is determined by the action of wave bottoms and tidal currents. Muds are deposited in depressions sheltered from currents, sands and silts in areas of weaker currents and medium to fine-grained sands in ribbons on shallower parts of the shelves. Turbidite: Turbidite currents formed by sub-marine slumps (Figure 11.36 pg. 283) deposit a characteristic sequence of sediments on the oceans abyssal plains. Turbidite sequences grade up from coarse structureless sand, to medium-grained, bedded sands, then finer sands and finally silts and muds (Figure 12.13 pg. 312). If the deposit formed close to the slump which caused the current, then the muddy top will be missing, and further away, the coarse base will be absent. Pelagic: Pelagic clays are fine-grained red clays which are the nonturbidite clastic deposits of the deep sea. The rate of sedimentation is so slow that iron in the clay becomes oxidized by sea water, giving the red colour. The clays are finely laminated, and manganese crusts and nodules are common. A variety of sedimentary environments exist at the same time in a region, and to define the sets of simultaneously deposited sediments, the word facies is used (Figure 12.14 pg. 313). For example, facies 1 may consist of marine offshore muds, whereas facies 2, deposited at the same time, may consist of shoreline sands. The extent of a given facies changes with time, and this may be due to marine transgression (advance of marine sediments over non-marine) or regression (advance of non-marine deposits). CHEMICAL SEDIMENTS Carbonates are the most common chemical sediments, formed due to the abundance of calcium and bicarbonate ions in sea water. Ca2+ + HCO3- = CaCO3 + H+ Limestone (CaCO3) is the most common carbonate rock, and also the related rock dolomite, CaMg(CO3)2. Many marine organisms, from one-celled animals to oysters, clams and other invertebrates, secrete some calcium carbonate. In this process of biological precipitation, the organisms extract calcium carbonate from the water and precipitate it to make their shells. Carbonate sedimentation is favored in warm tropical seas, especially in the coral reef habitat. Coral Reefs: Reefs are thought to originate from corals and algae colonizing the shores of volcanic islands and forming a fringing reef. As the island slowly sinks due to subsidence associated with sea-floor spreading, the deposition of coral (calcium carbonate which cements to the dead coral below) may keep pace with the sinking, and gradually builds up the reef. Eventually the volcanic centre disappears and is replaced by an atoll (coral island) with a central lagoon (Figure 12.17 pg. 316). Carbonate is also deposited in other environments, not always marine. The shallow platform in the area of the Bahamas Island has lead to deposition of carbonate over a large area, forming a carbonate platform. Abundant here are carbonate sands or oolites. These are spherical grains of aragonite (the unstable form of calcium carbonate), which begin from a shell nucleus and are rolled around by currents, depositing layer upon layer of calcium carbonate. Carbonate oozes formed from the remains of these organisms in the deep oceans are buried and lithified to form chalk. Calcium carbonate is also deposited around hot springs by algae and nonbiological precipitation, to form tufa (porous) and travertine (denser). Stalactites and stalagmites are formed by precipitation of calcium carbonate from saturated waters dripping from limestone. On entering the cave atmosphere, carbon-dioxide is lost, causing supersaturation of carbonate and hence precipitation. OTHER CHEMICAL SEDIMENTS Silica: Most of the chemically deposited silica is secrete biologically by small algae known as diatoms, and single-celled organisms known as radiolaria. They populate much of the surface of the ocean and freshwater lakes, and extract silica from the water to form their opaline, amorphous silica shells. Where these organisms are abundant due to a high supply of nutrients in the water, the shells of dead organisms sing to form silica-rich diatom ooze and radiolaria ooze, which cement and harden into diatomite and radiolarite respectively (Figure 12.27 pg. 323). These may lithify and recrystallized to form cherts. Sulfide: Organisms can control chemical sedimentation indirectly by changing the chemical conditions in the environment. Bacteria can change sulfur from its oxidized state, sulfate to its reduced state, sulfide. This process produces the smelly gas, hydrogen sulfide, which is a powerful reducing agent and changes ferric iron to ferrous iron, thus precipitating the mineral pyrite, FeS2. The activities of these bacteria keep the environment free of oxygen. The ocean may also become anoxic (deoxygenated) where basins are cut off from aerated waters by a ridge or barrier. In these basins, as organic matter decays, oxygen is used up, and is not replenished fast enough. Bottom waters may therefore become reducing, as in fjords. This deoxygenation can also be caused by pollution. Phosphate input will encourage algae and other plants to grow, to the point that the surface waters lack oxygen, a process called eutrophication (Figure 12.31 pg. 326). Coal: Swamps are areas of rich plant growth, which, on dying, falls to the waterlogged soil. The water and its rapid burial prevent it from oxidizing, and thus the vegetation does not decay completely. It forms pear, which after burial and chemical transformations becomes lignite (soft, brown, coal-like material). Burial to greater depths and thus higher temperatures metamorphoses the lignite to bituminous (soft) coal and eventually to anthracite (hard) coal. Evaporites: These are salts formed by the evaporation of sea water, such as halite (NaCl), gypsum (CaSO4.2H2O) and anhydrite (CaSO4). As sea water evaporates, a sequence of minerals is precipitated, starting with calcium carbonate and proceeding to sodium chloride and finally magnesium and potassium minerals. The concentrated solution formed at the surface from which evaporites precipitate is known as a brine (Figure 12.33 pg. 329). Brine is denser than sea water and sinks, removing the evaporites and the surface waters are replenished with sea water from the open ocean. Evaporites are paleoclimatic indicators as the extensive evaporation is only found in tropical or subtropical seas, or in lakes or salars of arid or semi-arid regions. SEDIMENTARY STRUCTURES The movement of sand grains in current creates ripples and dunes on the streambed as well as familiar horizontal bedding planes. These structures are known as bedforms. See Figure 8.9 and 8.10 (pg. 183 and 184). Information on the connection between current and sedimentary structures comes from laboratory experiments where streams are simulated using flumes. Ripples formed in an alluvial environment have a gentle slope upstream and a steep slope downstream, and are thus asymmetrical ripples. Ripples formed by wave action are symmetrical, with much sharper crests than current ripples. Formed by the back and -forth movement of waves, these are known as oscillation ripples. The inclined bedding associated with the formation of ripples is known as cross-bedding (Figure 8.11 pg. 184). Different cross-bedding forms are diagnostic of different environments. The angle of the cross-bedding is the angle of the downstream, or lee, slope of the ripple. As water velocity increases, the ripples move faster and grow larger, and are called dunes. As the dunes grow larger, small ripples form and climb up their backs and disappear over the lee slope. As the velocity of the water increases even further, ripples and dunes disappear and the bed becomes flat again. Supplementary Reading/Information: Chapter 12 - Earth SEDIMENTARY ENVIRONMENTS: GLACIERS & GLACIATION Glaciers cover 10% of the Earth's surface; though during ice ages have been 3 times more extensive. A glacier is a thick ice mass. Valley or alpine, glaciers (Figure 10.11 pg. 240) are a stream of ice that flow down valley, whereas ice sheets flow out in all directions and are continental-scale features e.g. Antarctica. Ice caps are areas of ice covering upland/plateau areas e.g. Iceland. Piedmont glaciers form where valley glaciers merge at the base of steep mountains to form a sheet. GLACIER MOVEMENT The upper part of a glacier (upper 50m) is brittle and referred to as the zone of fracture, where tension creates cracks (crevasses). Below 50m depth, the pressure is sufficient for the ice to behave plastically and to flow. The glacier may also move by the whole mass of ice slipping along the ground. Due to drag created by valley walls, ice flow is greatest in the centre of the glacier. The overall rate of flow of glaciers varies greatly from cm/day up to several metres/day. Glaciers can move in surges where the rate of flow periodically quickens. GLACIER BUDGET The zone of accumulation is where snow accumulates and ice forms, its outer (lower) limit being defined by the snow line. Below the snow line is the zone of wastage where melting occurs. Glaciers can also waste by the process of calving – large blocks of ice breaking off the front of the glacier which become icebergs where the glacier has reached a lake or the sea (Figure 10.5 pg. 236). The glacial budget is the balance between accumulation and ablation (loss). If accumulation>ablation = glacier advances If ablation>accumulation = glacier retreats GLACIAL EROSION Due to the competency of ice, glaciers can carry huge blocks of rock that no other erosional agent can transport, and cause great erosion. Glaciers erode the land in two ways plucking and abrasion. Plucking occurs when meltwater penetrates cracks and joints in the bed rock and freezes. As the water expands it breaks the rock into dust and blocks which become part of the glacier load. Abrasion is where the ice load scratches and polishes the bedrock. The pulverized rock formed by abrasion is called rock flour. Long scratches and grooves created by larger material are called striations (Figure 10.18 pg. 244). The rate of glacial erosion is dependent on: 1. Rate of glacial movement 2. Thickness of the ice 3. Shape, abundance and hardness of the rock fragments in the base of the glacier 4. The erodibility of the surface beneath the glacier GLACIAL EROSIONAL LANDFORMS Valley or alpine glaciers tend to create a sharp, angular accentuated topography whereas ice sheets override terrain and instead subdue the topography. Glaciated valleys display a Ushaped glacial trough and are wider and deeper than river valleys. The glacier straightens the valley creating truncated spurs. Tributary glaciers on either side of the main trunk glacier leave hanging valleys. Depressions formed by abrasive plucking and scouring, once the glacier has retreated, become filled with water and are called pater noster lakes. At the head of a glacial valley is a cirque, where accumulation occurs. Fjords form where steep sided glacial valleys have been flooded by the sea. As a group of cirques around a mountain become enlarged they leave between them horns and arêtes - spires and sharp-edged ridges of rock (figure 10.20 pg. 246). As the ice moves it can carve small hills in the bedrock known as roche mouton née (Figure 10.27 pg. 251), with a gentle abraded slope facing the oncoming ice and a steep, plucked face on the 'shadow' slope. GLACIAL DEPOSITS Glacial drift is the term used to describe sediments of glacial origin and can be divided into two distinct types: 1. Till – materials deposited directly by the glacier 2. Stratified drift – sediments laid down by glacial meltwater. Till is unsorted sediment as the ice carries material of all sizes, whereas stratified drift is sorted by the water according to the size and weight of the fragments. Boulders lying free at the surface which are of different composition to the bedrock on which they sit are called glacial erratics i.e. derived from a source outside the area (Figure 10.17 pg. 244). OTHER GLACIAL FEATURES The sides of a valley glacier accumulate large quantities of debris which, when the glacier wastes, leaves ridges called lateral moraines (Figure 10.19 pg. 245). Medial moraines form where two valley glaciers meet to form a single ice stream. End moraines form at the terminus of a glacier. The terminal moraine is that which marks the furthest extent of the glacier, and recessional moraines mark stationary positions during glacial retreat. During retreat, till is laid down forming a gently undulating surface of ground moraine. On the downstream edge of most end moraines is a ramplike surface of stratified drift called an outwash plain for an ice sheet and a valley train for a valley glacier. Basins or depressions in the end moraines and outwash are kettles, formed where blocks of stagnant ice are buried in drift and melt. Drumlins are streamlined asymmetrical hills composed of till, with a steep slope facing the ice and gentle slope in the direction of movement. They form when glaciers advance over pre-deposited drift and reshape it. Sinuous ridges of sand and gravel called eskers form where streams flow in tunnels beneath the ice near a glacier terminus. Kames are steep-sided hills of sand and gravel, formed where meltwater washes sediment into depressions in the wasting terminus of a glacier. THE ICE AGE By studying and dating ancient glacial deposits, it is known that the Ice Age was characterized by a series of advances and withdrawals of glaciers. Ocean floor sediments provide a record of climatic cycles and studies of core indicate that glacial/interglacial cycles have occurred about every 100,000 years and that the Ice Age consisted of about 20 such cooling/warming cycles. The Ice Age began 2 – 3 Ma ago and occurred during the Pleistocene epoch. There is also evidence for three earlier periods of glacial activity at 2 billion, 600 million and 250 million years ago. In addition to massive erosional and depositional work, Ice Age glaciation has other effects, including the forced migration of animals, changes in stream/river courses, isostasy (adjustment by rising of the crust after the weight of the ice is removed), climate changes caused by the existence of the glacier (i.e. cooler temperatures, less evaporation in arid areas, pluvial lakes etc.) and, most notably, changes in sea-level (advanced shorelines during glacial periods). CAUSES OF GLACIATION Any explanation of the causes of glacial ages must account for: 1. What causes the onset of glaciation, i.e. what causes a drop in temperature? 2. What causes the alteration of glacial and interglacial stages, i.e. short term changes? Two of the main hypotheses for the occurrence of glacial ages are: Plate Tectonics: Plate movements mean that land masses have shifted in relation to one another and moved to different latitudinal positions. This is accompanied by changes in oceanic circulation, altering the transport of heat, moisture and thus the climate. Climatic changes due to shifting plates are extremely gradual and cannot explain glacial/interglacial cycles. Variations in Earth's Orbit: Climatic oscillations which are responsible for glacial/interglacial cycles can be brought about by variations in the Earth's orbit. These Milankovitch cycles (after the scientist who developed this hypothesis) relate to variations in the receipt of solar energy, the corresponding surface temperature of the Earth and the degree of contrast between seasons. This is due to: Variations in the shape (eccentricity) of Earth's orbit around the sun Changes in obliquity (the angle the axis makes with the plane of Earth's orbit The precession of Earth's axis Supplementary Reading/Information: Chapter 10 - Earth SEDIMENTARY ENVIRONMENTS: DESERTS A dry climate such as exists in a desert is defined as one in which annual precipitation is less than the annual loss of water by evaporation. Dryness is thus related to temperature, since it effects precipitation. 30% of the Earth's land surface is classified as dry, and in these dry regions, two climatic types are recognized: 1. desert which is arid 2. steppe which is semi-arid Steppe is the marginal transition zone separating deserts from more humid regions. Dry lands are concentrated in the sub tropics and in the middle latitudes. LOW LATITUDE DESERTS A virtually unbroken desert environment stretches almost 10,000 km from the Atlantic Coast of North Africa to North West India. There is a smaller area of tropical desert and steppe in northern Mexico and south-western United States. 40% of the Australian continent is desert. These dry regions coincide with zones of high air pressure called subtropical highs. In these places air currents are subsiding, which causes compression and warming, leading to blue skies and ongoing drought (Figure 9.1 pg. 212 and Figure 9.10 pg. 217). MIDDLE LATITUDE DESERTS Middle latitude deserts and steppes exist principally because of their position in the deep interiors of large land masses. The presence of mountain ranges act as a barrier to prevailing winds carrying maritime moisture. The dry region often present on the leeward side of such mountain ranges is referred to as a rainshadow desert. For example, in Asia, the Himalayas prevent moist monsoon air from the Indian Ocean from reaching the interior. WEATHERING In dry lands, the rate of weathering is very slow due to the lack of moisture and scarcity of organic acids from decaying plants. Oxidation is the main weathering process (chemical) in deserts. Water: Desert streams are ephemeral, which means they carry water only in response to specific episodes of rainfall i.e. they flow intermittently. When rain falls, it cannot soak in so high run-off and flash floods are common. As the surface material is not anchored by vegetation, a large volume of material is eroded by these intermittent desert water flows. Rivers that cross dry areas are rare and originate outside the desert. For example, the Nile originates in the Central African mountains and traverses 3,000 km of the Sahara without a single tributary. Wind: The main role of wind in dry areas is not erosion but transportation and deposition of sediment to create dunes. Similar to water, a wind increases in velocity as height increases from the ground and heavier particles are carried as 'bed load' closer to the ground. The bed load of a wind consists of sand which moves by saltation (bouncing along the surface) (Figure 9.3 pg. 213). The suspended load consists of finer dust particles (i.e. silt) which are swept high into the atmosphere by the wind. The wind velocity in a thin layer close to the ground is almost zero and in order to become a suspended load, dust must be disturbed and lifted from the ground before the wind can transport it. In dry regions, wind can be erosive. Deflation is the lifting and removal of loose material, which can create shallow depressions called blowouts, the size of which is limited by the level of the water table. When wind has removed large amounts of dust and sand, desert pavement is left – that is a coarse cover of pebbles and gravel (Figure 9.8 pg. 216). Wind can also erode by abrasion, leaving polished and shaped stones called ventifacts. WIND DEPOSITS There are two types of wind deposit: 1. Dunes - mounds and ridges of sand from the wind's bed load 2. Loess - extensive blankets of silt once carried in suspension When moving air encounters an obstacle, a shadow of slower moving air exists behind it, and hence sand is deposited. The mound of sand created grows into a dune. Sand grains saltate up the windward side, and sand accumulates in the shadow just beyond the crest of the dune. This is known as the leeward side of the dune, or slip face, with an angle of 34° maintained. Sand deposited on the slip face is cross bedded. Slumps and slides occur on the slip face to maintain the angle of repose and the dune slowly migrates in a windward direction. TYPES OF DUNES Barchan Dunes: crescent shaped solitary sand dunes where sand supplies are limited and the surface relatively flat and hard (Figure 9.26 pg. 228). Transverse Dunes: Dunes form in a series of long ridges separated by troughs. They form where the prevailing winds are steady and sand plentiful (Figure 9.27 pg.228). Barchanoid Dunes: These are intermediate between barchan and transverse. Longitudinal Dunes: Long ridges of sand which form parallel to the prevailing wind, and where sand supplies are limited (Figure 9.28 pg. 229). Parabolic Dunes: These form where vegetation partially covers the sand. They are similar to barchans except their tips point into the wind rather than downwind. They often form along coasts where there is a strong on-shore wind. Star Dunes: These are isolated hills of sand resembling star shapes. They form where wind directions are variable. LOESS These deposits of windblown silt lack any visible layers or structure. Deserts and glacial deposits are the two principal sources of loess. The thickest, most extensive loess deposits are in Western and Northern China, blown from the desert basins of Central Asia. Loess in the United States and Europe is, in contrast, from glaciation (Figure 9.29 pg. 229). BASIN AND RANGE LANDSCAPE Basin and range regions form where the climate is arid with interior drainage. Mountains are uplifted and carved by streams which carry debris into the basins. The more erosion occurs, the less is the contrast in relief. Sporadic rains produce cones of debris at the base of a slope or mouth of a canyon i.e. alluvial fan. With time, adjacent fans coalesce to form an apron of sediment called a bajada along the mountain front. When rainfall is particularly abundant, streams may flow across the alluvial fan to the basin floor to form a playa lake. This soon disappears due to evaporation and infiltration. In the late stages of erosion, the mountain ranges are reduced to a few large lumps of bedrock called inselbergs. Supplementary Reading/Information Chapter 9 - Earth SEDIMENTARY ENVIRONMENTS: FLUVIAL AND GROUNDWATER The hydrologic cycle is the continuous interchange of water between the oceans, atmosphere and continents, including precipitation, evaporation, infiltration, runoff and transpiration. FLUVIAL Initially, water flows as runoff (thin sheets) but after a short distance, threads of current develop and tiny channels called rills form. The amount of runoff depends on the infiltration capacity of the land. Rills become streams and the factors determining a stream velocity are: gradient cross sectional shape roughness of the channel discharge Gradient and roughness usually decrease downstream whilst the other properties increase. The base level is the lowest point to which a stream may erode its channel, which may be: 1. Ultimate base level (sea level) 2. Temporary or local base level Lowering base level causes a stream to erode and raising base level causes deposition. The work of a stream includes erosion, transportation, (as dissolved load, suspended load and bed load), and when its velocity decreases, deposition. The capacity of a stream is the maximum load of solid particles it can transport and competence is the maximum particle size a stream can transport. A streams (or rivers) depositional feature includes deltas (where the stream is slowed on entering a large body of water such as a lake or ocean) and natural levees (banks of sediment on either side of the stream deposited during flood. Two general types of stream valley exist: Narrow V-shaped valleys form because the stream is downcutting towards base level often contain waterfalls and rapids. Wide, flat-floored valleys form when a stream has cut its channel closer to base level, and its energy is directed from side to side so that erosion produces a flat valley floor or floodplain. Streams flowing upon floodplains often meander (Figure 8.22 and 8.25 pg. 194195) and widespread meandering may result in shorter channel segments called cut offs and abandoned bends called oxbow lakes. The land area that contributes water to a stream is called a drainage basin. These are separated from each other by imaginary lines called divides. The network of streams in a drainage basin forms various patterns including dendritic, radial, rectangular and trellis. GROUNDWATER Groundwater is one of the most important and widely available resources - the largest reservoir of freshwater that is readily available to humans. Geologically, groundwater is important as an erosional agent. Its dissolving action creates subterranean caverns and also surface depressions called sinkholes. THE DISTRIBUTION OF UNDERGROUND WATER Groundwater is that water which completely fills the pore spaces in sediment and rock in the subsurface zone of saturation. The upper limit of the zone is the water table (Figure 7.13 pg. 163) and the zone of aeration is above the water table where the soil, sediment and rock are not saturated. Groundwater generally moves within the zone of saturation. The quantity of water that can be stored in this zone depends upon its porosity – the volume of open space. However, the primary factor controlling the movement of groundwater is the permeability – the ability to transmit a fluid through interconnected pore spaces. The water table has a very irregular surface, mainly because groundwater moves very slowly, and due to variations in rainfall and permeability from place to place. The water table usually follows the surface topography, with its highest elevations beneath hills, descending towards the valleys. The movement of groundwater into channels maintains stream flow even during dry periods. These streams are said to effluent. In arid regions, permanent streams usually originate in wet regions before flowing through the arid area. Under these conditions, the zone of saturation is supplied by downward seepage from the stream channel. These streams that provide water to the water table are called influent streams. Generally, groundwater moves under the force of gravity but sometimes it may move upwards from zones of higher pressure to lower – for example, the deeper you go into the zone of saturation the higher the water pressure. The pressure is greater beneath a hill but low beneath a stream channel. Hence water will migrate towards the channel, which may involve some upward movement. Impermeable layers of rock or sediment that hinder or prevent water movement are termed aquitards e.g. clay. In contrast, permeable rock strata that can transmit groundwater freely are called aquifers (Figure 7.14 pg. 163). SPRINGS Springs occur wherever the water table intersects the land surface and a natural flow of groundwater results. Wells are bored into the zone of saturation to withdraw groundwater, creating roughly conical depressions in the water table known as cones of depression. Artesian wells are when groundwater rises above the level where it was initially encountered. For this to occur, the water must be stored under pressure i.e. stored in an aquifer which is confined by aquitards above and below. One end of the aquifer must be open to the surface so that it can receive water. ENVIRONMENTAL PROBLEMS OF GROUNDWATER Groundwater is being exploited at an increasing rate, causing various environmental problems. This includes: Overuse of groundwater by intense irrigation means that in some areas the water table has dropped by 1 m annually. It takes thousands of years for the groundwater to be fully replenished. Land subsidence – withdrawal of groundwater means that the weight of overburden packs the sediment more tightly together and the ground subsides. This is particularly prominent in areas underlain by thick layers of loose sediment. Contamination – sewage, farm wastes and fertilizers are common sources of pollution. This is a major problem when aquifers that supply a large part of the water supply to a population become contaminated. HOT SPRINGS AND GEYSERS Groundwater circulating at depth becomes heated and if it rises to the surface, emerges as a hot spring (Figure 7.21 pg. 173). When groundwater is heated in an underground chamber, it expands and converts to steam which escapes as a geyser. Groundwater from hot springs and geysers usually contains more material in solution than groundwater from other sources due to its temperature. When the water contains a lot of dissolved silica, geyserite is deposited around the spring, whereas in limestone areas the calcite deposit formed is travertine. Geothermal energy is harnessed by tapping underground reservoirs of steam and hot water. A good geothermal reservoir will have: 1. A potent source of heat such as a large magma chamber e.g. in volcanic regions 2. Large and porous reservoir with channels connecting it to the heat source 3. A cap of low permeability rock to inhibit the flow of water and heat to the surface KARST A landscape that has to a large extent been shaped by the dissolving power of groundwater - exhibits what is known as karst topography - an irregular terrain punctuated with many depressions called sinkholes or sinks. A karst topography will have dripstone features in caverns, collectively known as speleothems. Supplementary Reading/Information: Chapters 7 and 8 - Earth SEDIMENTARY ENVIRONMENTS: SHORELINE AND PELAGIC SHORELINE Winds create surface currents, gravity produces tides, and density differences create deep ocean circulation, all of which have an impact on the shoreline and how it is shaped. WAVES Wind drags the surface of oceans into waves. The tops of waves are crests separated by troughs (Figure 11.13 pg. 267). The vertical distance between the two being wave height. Horizontal distance between crests is wave length and wave period is the time interval between the passings of two crests. These physical properties of a wave depend on: 1. Wind speed 2. Length of time the wind has blown 3. Fetch – the distance the wind has travelled across open water With increasing distance from a stormy area, waves lose energy and gradually change to swells (lower height and longer length). Two types of wind generated waves are: 1. Waves of oscillation - generated in the open sea in which the wave form advances as the water particles move in circular orbits. 2. Waves of translation - turbulent advance of water formed near the shore as waves of oscillation break into surf. Erosion of shorelines by waves is due to wave impact, pressure and abrasion (grinding action of water carrying rock fragments). As waves approach a shoreline, the near-shore end of a wave is slowed first as it reaches shallower ground, and therefore the waves bend. This wave refraction (Figure 11.12 pg. 267) means that wave impact is concentrated against the sides and ends of headlands. Although wave refraction tends to bend waves around towards a parallel trend with the shore, most waves still reach the shore at an angle, but the backwash of water from each breaking wave moves straight down the slope of the beach. This causes a zigzag motion called beach drift, where sand and pebbles are moved along the coastline. Oblique waves also produce longshore currents that flow parallel to the shore and are capable of carrying large quantities of sediments e.g. in parts of California, 1-5 million tons of sediment are moved along the shore each year (Figure 11.21 pg. 272). SHORELINE FEATURES Shoreline erosion produces various features such as: wave-cut cliffs: from the cutting action of surf against the base of coastal land (Figure 11.5 pg. 262) wave-cut platforms: relatively flat, benchlike surfaces left behind by receding cliffs (Figure 11.7 pg. 263) sea arches: when a headland is eroded and two caves from opposite sides coalesce sea stacks: formed when the roof of a sea arch collapses DEPOSITIONAL FEATURES Deposition of sediment transported by beach drift and longshore currents forms various features such as: spits: elongated ridges of sand that project from the land into the mouth of an adjacent bay baymouth bars: sand bars that completely cross a bay tombolos: ridges of sand that connect an island to the mainland or another island The Atlantic and Gulf Coastal Plains have a shore line characterized by barrier islands – low ridges of sand parallel to the coast and 3 – 30 km offshore. Barrier Island may originate as spits or as sand dune ridges from a period when sea-level was lower. SHORELINE EVOLUTION A shoreline is continually modified and initial erosion may increase its irregularity e.g. weaker rocks eroded more easily than stronger ones. However, if a shoreline remains stable, marine erosion and deposition will even out to produce a straighter more regular coast. Local factors that influence shoreline erosion are: 1. the proximity of a coast to sediment - laden rivers 2. the degree of tectonic activity 3. the topography and composition of the land 4. prevailing winds and weather patterns 5. the configuration of the coastline and nearshore areas HUMAN RESPONSES TO SHORELINE EROSION In order to preserve buildings and development in coastal locations, natural migration of sand is controlled by various means. This includes: 1. Building structures such as: groins (short walls built at a right angle to the shore to trap moving sand) (Figure 11.22 pg. 273); breakwaters (structures built parallel to the shoreline to protect it from the force of large breaking waves); seawalls (barriers to prevent waves from reaching the area behind the wall). 2. Beach nourishment - sand is added to replenish eroding beaches 3. Buildings are relocated away from the beach The nature of shoreline erosion problems along the American Pacific and Atlantic coasts is very different. Atlantic and Gulf coast development has occurred on barrier islands which receive the full force of major storms. The Pacific coast however has narrow beaches backed by steep cliffs and mountain ranges. The problem here is narrowing beaches caused because the natural flow of sediment to the coast has been interrupted by dams built for irrigation and flood control. COASTAL CLASSIFICATION Coasts can be classified based upon changes that have occurred with respect to sea level. Emergent coasts characterized by wave-cut cliffs and platforms above sea level, develop either because an area experiences uplift or as a result of a drop in sea level. Submergent coasts are characterized by drowned river mouths (estuaries) and are created when sea level rises or the land adjacent to the sea subsides. TIDES Tides, or the daily rise and fall of sea level, are caused by the gravitational attraction of the moon and to a lesser extent the Sun. Near the times of new and full moons, the Sun and Moon are aligned and their gravitational forces add together to produce especially high and low tides known as the spring tides. Conversely, at the times of the first and third quarters of the Moon, the gravitational forces of the Moon and Sun are at right angles and the daily tidal range is less. These are neap tides (Figure 11.23 pg. 274). THE PELAGIC (OCEAN FLOOR) ENVIRONMENT 71% of the Earth's surface consists of oceans and marginal seas. In the Southern Hemisphere where there is less land mass, about 81% of the surface is water. The Pacific Ocean is the largest, containing more than half of the water in the world ocean with the greatest average depth of 3,940 metres. The volume of all land above sea level is actually only 1/18 that of the ocean. The ocean floor is characterized by mountains, deep canyons and flat plains, similar to the scenery on the continents. In the 1920s, the invention of electronic depth-sounding equipment allowed a continuous profile of the ocean floor to be imaged. The echo-sounding equipment is a device towed by a ship which sends out sound waves which are bounced off the ocean floor (Figure 11.29 pg. 278). The two-way-time (the time for the sound waves to travel from the emitter to the ocean floor and back) is directly related to the depth. Continuous two way time data is plotted to produce a profile of the ocean floor. SUBMARINE DIVISIONS The ocean floor can be divided into three major topographical units: 1. Continental margins 2. Ocean basin floor 3. Mid-ocean ridges The zones that make up the continental margin include: Continental shelf – a gently sloping submerged surfaced extending from the shoreline towards the deep-ocean basin Continental slope – a steep slope marking the true edge of the continent from shelf into deep water Continental rise – where slope merges into sediments that have moved downslope from the continental shelf to the deep-ocean floor See Figures 11.31 and 11.33 pgs. 280 – 281. SUBMARINE FEATURES Continental Margin Submarine Canyons and Turbidity Currents: Submarine canyons are deep, steepsided valleys that originate on the continental slope and may extend to depths of 3 km. Some of these canyons are seaward extensions of river valley, though most have been excavated by turbidity currents (Figure 11.36 pg. 283). The latter are downslope movements of dense, sediment-laden water. Turbidites are sediments deposited by these currents, and are characterized by a decrease in sediment grain size from bottom to top (graded bedding). Ocean Basin Floor: This lies between the continental margin and the mid-ocean ridge system. Features of the ocean basin floor include: Deep-ocean trenches - long narrow troughs that mark the boundary between two plates at a subduction zone Abyssal plains - thick accumulations of sediments deposited on the ocean floor Seamounts - isolated, steep-sided volcanic peaks on the ocean floor that originates near oceanic ridges or in association with volcanic hot spots. Mid Ocean Ridges: these are the sites of sea-floor spreading, representing more than 20% of the Earth's surface. They are the most prominent features in the oceans forming an almost continuous mountain range. Ridges are characterized by an elevated position, extensive faulting and volcanic structures developed on newly formed oceanic crust. Most of the activity associated with ridges occurs along a narrow region on the ridge crest called the rift-zone – the region where magma from the asthenosphere moves upwards to create new crust. The Pacific Ocean is older than the Atlantic. Hence the mid-ocean ridge in the Pacific – the East Pacific Rise, has largely been overridden by the subduction of the ocean below the American continents and the consequent westward migration of the American plate. In contrast, the Mid-Atlantic Ridge is still very active, standing at 2,500 – 3,000 m above the adjacent deepocean floor. In Iceland the ridge actually extends above sea level. SEAFLOOR SEDIMENTS There are three broad categories of sea floor sediment. Terrigenous sediment – consists primarily of mineral grains that have been weathered from continental rocks and transported to the ocean. Biogenous sediment – consists of shells and skeletons of marine animals and plants. Calcareous and siliceous oozes are the most common biogenous sediments. Hydrogenous sediment – consists of minerals that crystallize directly from seawater through various chemical reactions. For example, manganese nodules rounded black lumps composed of >20% manganese and other valuable metals such as Fe, Cu, Ni and Co. These manganese nodules are often a potential resource. Supplementary Reading/Information: Chapter 11 – Earth STRATIGRAPHY AND STRUCTURE Stratigraphy is the study of rock strata as a record of the geological history of an area. The geological history can be interpreted to show how that area evolved in terms of its plate tectonic setting throughout time. A sequence of sediments kilometres thick has accumulated over a length of time outwits normal comprehension. For example, if 0.1 mm of sediment accumulated in one year, this would amount to 1 km of sediment in 10 million years. This is the sort of time scale geologists work with. RELATIVE DATING The geological time scale allows the geologic events in the Earth's history to be placed in sequence according to relative age. Until the 1960's and the development of radioactive dating methods, the ages of rocks were expressed in terms of named intervals of relative time, based on the relationships between layers of sediments. The two fundamental principles behind this are: 1. A particular layer is younger than the one beneath it and older than the one on top. 2. A bed can be identified by characteristic fossils that it contains. Layers of rock can thus be mapped as formations i.e. groups of layers that have the same stratigraphic age and contain materials that have the same physical appearance and properties (lithology). In this way, an individual bed may be recognized in widely separated localities. Units of rock can be identified with distinguishable fossil assemblages, although individual species may be present in different formations. The fossils present will vary with the environment, so rocks in different places may yield quite different faunas. The study of fossils (palaeontology) provides the most useful and widely used means of relative dating and correlating sedimentary sequences. Petroleum exploration commonly uses microscopic animal fossils (micropalaeontology) or the spores and pollen from plants (palynology). RADIOACTIVE DATING Where rocks contain suitable radioactive material, absolute ages can be determined. For example, K-Ar dating measures the proportion of argon derived from the breakdown of radioactive potassium, and by knowing the rate of this breakdown and measuring the amounts present, the age of the rock can be determined. These techniques are expensive and time consuming. GEOLOGICAL TIME-SCALE Animals appeared on Earth 570 million years (Ma) ago. The time before this is referred to as the Pre-Cambrian, and the Cambrian period since then is split into three eras: the Palaeozoic (early), the Mesozoic (middle) and the Cenozoic (recent). Each of these eras is subdivided into named periods, and even smaller units still. Figure 2.25 (pg. 39) shows the eras and periods of the geological timescale. The period names are used to refer to rocks that formed during that period e.g. Cretaceous rocks formed between 135 and 65 Ma. ROCK DEFORMATION I STRUCTURE Rocks are usually deposited parallel to the Earth's surface and may be subsequently deformed by tilting or faulting. When stresses are applied to a rock it may behave in either a brittle or a ductile manner, resulting in fractured or folded rocks respectively. This depends on: a. Temperature b. Depth of rock e.g. shallow rock tends to fracture or fault whereas deeper rocks tend to deform smoothly. c. The nature of the rock e.g. crystalline basement rocks tend to be brittle (fracture), whereas sediments tend to be ductile (fold). d. The time the pressure or stress is applied for. FOLDS When a rock deforms in a ductile manner, it will fold. Folds exist on large scales (e.g. mountain belts) or on small scales (e.g. individual beds of sediments). The folding may be gentle or severe. The axis of a fold is the intersection of the axial plane (symmetrical division of a fold) with the beds (Figure 4.14 pg. 87). When the limbs of the fold dip at the same angle, the axial plane is vertical and the fold is symmetrical. When the beds in one limb dip steeper than the other, the fold is asymmetrical, and the angle of the axial plane to the vertical is known as the plunge of the fold. When the beds on both limbs of the fold dip in the same direction, the fold is overturned. When the axial plane is nearly horizontal, the fold is recumbent. (Figure 4.15 pg. 88) Upfolds, or arches of layered rocks are called anticlines, and downfolds, or troughs are called synclines (Figure 4.12 pg. 87). A steplike bend in otherwise horizontal or gently dipping beds is a monocline (Figure 4.13 pg. 87). Note that topographic expression is not necessarily a reflection of deformation i.e. hills do not necessarily correlate with the top of an anticline, nor valleys with the wells of synclines. FRACTURES When a rock deforms in a brittle manner, it will fracture. There are two categories of fracture: 1. Faults. This is where rock is displaced either side of or parallel to a fracture. Faults are common in mountain belts of where deformation is intense. Faults are assigned different names according to their sense of movement e.g. normal, reverse, thrust, dip-slip, strike-slip, oblique-slip, right-lateral, left-lateral etc. (see Figure 4.24 pg. 92) A transform fault is where two plates meet at a passive margin (i.e. not a constructive or destructive plate margin). For example, the San Andreas Fault is a classic transform fault, the two plates being offset by hundreds of kilometres. 2. Joints. This is where the rock has cracked but no appreciable movement has occurred. This tends to occur where regional stresses are applied, e.g. sediment compression leads to joint formation. If the pattern of joints is regular then the stress system must have been uniform. Intersecting joints create blocks of rock, and the joints allow the passage of water, thereby speeding up the weathering process. Joints may also provide channels for magma, leading to parallel swarms of dikes etc. Faulting leads to the formation of structures known as grabens and horsts (Figure 4.27 pg. 93). A graben forms by tensional crustal forces leading to down dropping of a faulted block, for example at mid-ocean ridges or rift valleys (East African Rift). Grabens are thus long narrow valleys bounded by one or more parallel normal faults. A horst is the opposite. It is a ridge formed by parallel reverse or normal faults. When studying structures to determine the geological history of an area, it is important to realize that a fault must be younger than the youngest rocks it cuts, and older than the oldest undisrupted formation that covers it. Fault zones in the field can be recognized by crushed, ground up rocks (cataclastics), and also by slickensides (polished striated friction marks). UNCONFORMITIES Sometimes the process of sedimentation at a particular location will cease. This may be because the sediment has built up to sea level so that no more can accumulate there. For whatever reason, the break in the sedimentary sequence is known as a hiatus. During a hiatus, it may happen that the sediments are uplifted above sea level, probably tilted in the process, and eroded. Eventually erosion will level off a surface that cuts obliquely across the bedding of the sediments. If then subsidence occurs and the process of sedimentation continues, a new sequence of strata will be deposited horizontally over the tilted and truncated older sequence. The erosional surface representing this interruption in sedimentation and differential erosion is known as an unconformity. EVOLUTION OF A SEDIMENTARY BASIN Since much oil and gas forms below the sea in sedimentary basins, it is useful to understand how such a basin may develop, and the types of structures that can be expected to form in such an environment. The sea-floor has been subsiding continuously or intermittently for at least 240 Ma. Regions where such subsidence has occurred, and sediments accumulated are referred to as sedimentary basins. When a basin first forms it may be isolated from the sea, and the first sediments probably accumulated on land. If subsidence is faster than the rate of sediment accumulation, then water depth will increase and finer-grained rocks will form such as muds and shales. As the processes of accumulation and subsidence continue, older sediments become squashed, water is squeezed out, and diagenesis or hardening of the sediment to rock (induration) occurs. Subsidence is not uniform, causing highs and lows in the basement surface and thus variation in the nature of sediment being deposited. Differential subsidence leads to the formation of gentle anticlines and synclines, and also stresses the sediments to cause faulting. A cease in subsidence may lead to a hiatus and the formation of an unconformity. At depth, the basin will become compressed by forces associated with plate motions etc. and buckled or folded. GEOLOGICAL MAPPING Geologists can record the geometry of tilts, folds and faults, and through reconstruction of maps and cross-sections, can thus determine the deformation history of rocks. For example, in oil exploration, the surface of a particular stratum such as sandstone which is favorable to oil accumulation is mapped across an area by identifying its depth in a series of boreholes, and producing a contour map of it. In the field, the geologist records information about the structure of beds by measuring their strike and dip. The dip is the angle of inclination of the bed from the horizontal in the direction of steepest descent, and the strike is at right angles to the dip direction, and is the intersection of the plane of the bed with the horizontal (Figure 4.9 pg. 86). Once the formations are mapped and the dips and strikes recorded, the sub-surface geology is reconstructed (Figure 4.11 pg. 87). For petroleum geology and petroleum exploration, stratigraphical correlation, structure and geological mapping are important for a number of reasons: 1. To understand the plate tectonic setting of an area, in order to identify areas in space and time which are likely to have had the geological conditions favorable for the generation of hydrocarbons. 2. To identify structures which can act to trap hydrocarbons in petroleum deposits. 3. To recognize areas which can be correlated with known hydrocarbon producing regions, in the hope that an analogous geological situation will lead to identification of new deposits. Supplementary Reading/Information Chapters 2 and 4 – Earth GEOLOGICAL MAPPING AND CONSTRUCTION OF CROSS SECTIONS Hills and valleys are usually carved out of layered sequences of rock, or strata, in which the individual beds differ in thickness and resistance to erosion. The surface topography and landforms are a product of erosion. Map 1 is a simple geological map, illustrating the topographic contours and also the geological boundaries between different strata. In this case, the geological boundaries are parallel to the contour lines, indicating that the strata are horizontal. This is actually quite rare in nature, as the rock strata have usually been uplifted, faulted or folded. DRAWING A CROSS-SECTION To draw a section from Map 1, along line A-B. 1. Draw a base line on graph paper, the exact length A-B. 2. Mark off on the baseline the points at which the contour lines cross the line of section, and for each point, draw a vertical from the base line mark to the appropriate height on a vertical scale i.e. 8.5mm from A is the intersection of the 700m contour. 3. A topographic surface can be constructed by joining all these intersection points together. Note: Geological details of the section are often lost if the vertical scale used is equivalent to the horizontal scale e.g. if a geological map of scale 1 : 50,000 is used to construct a cross section, an equivalent vertical scale would be 1 cm = 500m. This vertical scale would be too small to show detail, and so the scale should be vertically exaggerated, e.g. 1 cm = 200m. Care should be taken not to over exaggerate the scale because strata may then appear very distorted. DIP Strata inclined to the horizontal are dipping. The angle of dip is the maximum angle measured between the strata and the horizontal. The direction of dip is given as a compass bearing from 0° to 360°. For example, 12/270 implies a 12° angle of dip to the west. The strike is the direction at right angles to the dip. See Figure 1. Figure 1 – Southerly dipping strata in a quarry. Note the relationship between the directions of dip and strike. STRUCTURE CONTOURS Contour lines can also be drawn on a geological map for a bedding plane (geological boundary). These are called structure contours or strike lines, and join points on a bedding plane of equal height. On a map, the height of a geological boundary is known where it crosses a topographic contour line. A straight line (the strike line) can be drawn between points on a geological boundary which are at the same height (example given in lecture). Strike lines are always straight and parallel, and if the dip of the beds is constant, they will be equally spaced. CALCULATION OF ANGLE OF DIP From the spacing between structure contours, the dip of the beds can be calculated. For example, if the distance (measured with a ruler) between the 700m strike line and the 600m strike line for a particular geological boundary is 1.25cm, and the scale of the map is 2.5cm = 500m, then 100m vertical drop of the bed occurs in 1.25cm = 250m. Hence, the gradient is given by 100/250, or 1 in 2.5. The angle by trigonometry is 1/2.5 = 22°. CALCULATION OF THE THICKNESS OF A BED On map 2, the 200m structure contour line for the Q-R boundary passes through the point where the P-Q boundary is at 400m. It follows that bed Q has a vertical thickness of 200m. VERTICAL THICKNESS AND TRUE THICKNESS When beds are inclined, the vertical thickness (for example that penetrated by a borehole) is greater than the true thickness of the bed, measured perpendicular to the geological boundary. The angle a between VT (vertical thickness) and T (true thickness) is equal to the angle of dip. cosine α = T/VT and T = VT x cosine α The true thickness of a bed is thus the vertical thickness multiplied by the cosine of the angle of dip. Where the dip is low, the cosine is high (approaching 1.0), and VT and T are approximately the same. See Figure 2. INLIERS AND OUTLIERS An outcrop of a bed entirely surrounded by outcrops of younger beds is called an inlier. An outcrop of a bed entirely surrounded by older beds is called an outlier. These features are usually the product of erosion. FAULTING Refer to lecture notes on stratigraphy and structure for illustrations of the geometry of faults i.e. normal, reverse, oblique faults etc. The throw of a fault is the vertical displacement of any bedding plane. See figure 3. The angle of dip of a fault it the angle it makes with the horizontal, and the angle of hade is its angle with the vertical. Calculation of throw of a fault from a geological map Structure contours are drawn for displaced stratum on either side of a fault. If for example, the 1000m structure contour for a bed on the west side of a fault coincides with the 500m structure contour for the same bed on the east side of the fault, then the fault has a downthrow to the east of 500m. SUMMARY The maps encountered so far have been quite simple geological maps, lacking in complex structural features such as multiple faulting, folding or unconformities. Geological maps and cross-sections can be constructed for very complex geological settings, although often, the more complex the geology becomes, the more uncertain becomes the sub-surface geology in the cross section. Figure 2 – Section showing the relationship between the vertical thickness (V/T) and the true thickness (T) of a dipping bed. Figure 3 – Section through strata displaced by a normal fault (after erosion has produced a nearlevel ground surface). Map 1 CLASS EXCERCISE Study map 2. The continuous black lines are the geological boundaries separating the outcrops of the dipping strata, beds P, Q, R, S, T and U. Note that the geological boundaries are not parallel to the contour lines, but, in fact, intersect them. This shows that the beds are dipping. 1. Draw structure contours for each geological interface. 2. Calculate the direction and angle of dip. 3. Construct a cross-section along line N-S, and illustrate on it the dipping beds P to T. 4. Calculate the vertical thickness and true thickness of beds Q and S. Map 2 Map 3