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INTRODUCTION TO GEOLOGY
LECTURE
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TITLE
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Plate Tectonics and Earth Structure Minerals:
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The Building Blocks of Rocks Igneous Rocks:
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Plutonism and Volcanism Earthquakes
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Metamorphic Rocks
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Sedimentation and Sedimentary Rocks
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Sedimentary Environments 1 : Glaciers and Glaciations
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Sedimentary Environments 2 : Deserts and Wind
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Sedimentary Environments 3 : Fluvial and Groundwater
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Sedimentary Environments 4 : Shoreline and Pelagic
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Stratigraphy and Structure
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Geological Mapping and Construction of Cross-Sections
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The Formation and Occurrence of Petroleum
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Petroleum Migration, Reservoirs and Traps
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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