Download A2 Course Study Guide

A2 Level Geology Study
Guide
Name……………………………..
……………………………………….
Contents
Minerals ...........................................................................................................................3
Igneous Petrology ..........................................................................................................6
Sedimentary Petrology .............................................................................................. 58
Metamorphic Processes and Rocks ....................................................................... 106
The Fossil Record ...................................................................................................... 116
Global Causes Of Climate Change Through Geological Time ........................... 160
Evidence of Global Climate Change ....................................................................... 175
Rock Deformation ..................................................................................................... 179
Geology of the Lithosphere .................................................................................... 195
Geological Evolution of Britain ............................................................................... 240
Geological Mapwork .................................................................................................. 277
2
MINERALS
All substances consist of very small particles called atoms. Atoms are made up of still smaller
particles called protons, neutrons and electrons. Each proton has a single positive charge; a
neutron has no charge; and electrons each have a single negative charge. Since atoms are
electrically neutral the number of protons in an atom equals the number of electrons. The
protons and neutrons together make up a central mass called the nucleus of the atom. The
electrons move in orbits around the nucleus.
Elements are substances consisting of atoms which all have the same number of protons. The
simplest element is hydrogen; its atoms consist of one proton and one electron. Helium has two
protons, two neutrons and two electrons in its atom. Oxygen atoms have eight protons, eight
neutrons and eight electrons. As there are over 100 elements you might expect an almost
infinite number of different crystalline arrangements (minerals) to form from them. In fact,
the total number of minerals discovered is only about 3,500, and the number of commonly
occurring rock-forming minerals is much smaller. The reason for this is that even though there
are lots of elements most of them are very rare. Thus, the widely occurring rock-forming
minerals represent the combinations of a small number of readily available ingredients, that is,
the abundant crustal elements (O, Si, Al, Fe, Ca, Na, K and Mg).
CHEMISTRY OF MINERALS
Minerals are natural elements and compounds which usually form crystals. The arrangement of
atoms within a crystal can be found by shining a beam of X-rays through it. X-rays have such a
short wavelength that they appear to bounce off layers of atoms inside the crystal. The
diffracted X-rays then form a pattern on a photographic plate which can be interpreted to find
the way in which the atoms are arranged. It has been found that the atoms in crystals are
arranged in three-dimensional frameworks called lattices. Most minerals are made up of ions
and the way in which the ions are packed together depends partly on the sizes of the ions. The
sizes of some ions are given in the table below. The ionic radii are given in picometres (pm); a
picometre is 10-12 metre, or a metre divided by a million, million!
Name of Ion
Sodium
Potassium
Magnesium
Calcium
Iron
Aluminium
Silicon
Oxygen
Hydroxide
Chlorine
Symbol & Charge
Radius (Picometres)
Na+
97
K+
133
Mg2+
66
Ca2+
99
Fe2+
74
Al3+
51
Si4+
42
O2140
OH140
Cl181
Table to show radii of some common ions.
3
One of the simplest crystal structures is that of sodium chloride (Na +Cl-). Here, the sodium and
chloride ions are situated as if they were at the corners of cubes. Each sodium ion is
surrounded by six chlorine ions and each chlorine ion is surrounded by six sodium ions. Minerals
containing silicon combined with oxygen are silicate minerals and have a more complex crystal
structure. In silicate minerals one silicon ion is bonded to four oxygen ions at the corners of a
tetrahedron, called the silicate group (or SiO4 group).
Substances which have the same chemical composition but different atomic structures are said
to be polymorphs. Graphite and diamond are polymorphs of carbon. Graphite consists of sheets
of hexagonally-arranged carbon atoms. Diamond also consists of carbon but here the atoms are
arranged in a tetrahedron. (A tetrahedron is a pyramid shape with four faces which are all
equilateral triangles.) Polymorphism is very common among minerals. It also happens that
substances with different compositions have the same type of atomic structure and crystal
form. Such substances are said to be isomorphous. Isomorphism results from the replacement
of one ion by another in the crystal lattice without changing its atomic structure.
Minerals are often found whose compositions would seem to have resulted from the mixing of
other minerals, e.g. olivine. The name olivine actually refers to a continuous spread of chemical
compositions between Mg2SiO4 (the mineral forsterite) and Fe2SiO4 (the mineral fayalite). Its
chemical formula is usually written as (Mg,Fe)2SiO4, indicating that magnesium and iron can
substitute for each other. This kind of chemical mixing is called solid solution and occurs when
two ions have similar sizes (e.g., Mg2+ = 66 pm and Fe2+ = 74 pm), and charges (e.g., Mg2+ and
Fe2+), so that either ion could fit into the same crystal structure without changing it.
SILICATE MINERALS
Silicates are minerals in which metals such as magnesium, iron, calcium, sodium, potassium and
aluminium are combined with silicon and oxygen. Silicates are the most abundant rock-forming
minerals. Along with quartz (silicon dioxide) they make up all common rocks apart from
materials such as limestones.
In silicates the silicon and oxygen form tetrahedra with each small silicon ion surrounded by
four oxygen ions. Different silicates have their tetrahedra arranged in different ways. In the
simplest case the SiO4 tetrahedra are separate from each other. Olivine has a structure like
this. Minerals possessing a structure of separate or isolated SiO4 tetrahedra are very compact
so they have high densities (olivine density = 3.2 – 4.4), they are hard (olivine hardness = 6.5)
and they have no distinct cleavage.
In the pyroxenes (e.g. augite) the SiO4 tetrahedra form chains with each tetrahedron sharing
an oxygen ion with the tetrahedra on each side. The chains are held together by metallic ions.
Since the silicon-oxygen bonds within the chains are stronger than the bonds holding separate
chains, pyroxenes tend to split parallel to the chains in two directions which are nearly at right
angles to each other. Pyroxenes have relatively high densities (3.2 – 3.9) and they are relatively
hard (5 – 6) because their ions are quite densely packed.
4
The amphiboles (e.g. hornblende) have their SiO4 tetrahedra in double chains. The large
hexagonal holes in the chains are occupied by hydroxide ions (OH -) and the double chains are
held together by ions such as those as calcium and magnesium. The double chains are not
strongly held together so amphiboles have two pronounced cleavages at about 60° to each
other. The cleavages do not break the silicon-oxygen bonds of the chains. Amphiboles tend to
form long or fibrous crystals and they are relatively hard (5 – 6) and dense (2.9 – 3.6).
Minerals such as mica have their SiO4 tetrahedra in sheets. The sheets form double layers with
the tips of the tetrahedra pointing inwards. The spaces within the sheets are occupied by
hydroxide ions. Between the double sheets there are potassium ions which weakly bond the
sheets together giving mica its very good cleavage parallel to the sheets.
The framework silicates are the most abundant silicates in the Earth‘s crust. They have
complex crystal structures, with each SiO4 tetrahedra joined to four others, giving a
polymerized, three-dimensional framework. The framework of quartz, SiO2, is not too compact
so the relative density of quartz is only 2.65. The structure is a strong one, however, so quartz
has a hardness of 7 and it lacks cleavage. Feldspar is another framework silicate and makes up
approximately 70% of the Earth‘s crust. The name feldspar actually refers to a group of
silicate minerals, which share the same basic structure: some silicon replaced by aluminium,
with K, Na or Ca ions inside the cavities of the tetrahedral framework. The three most
important feldspars are: orthoclase (KAlSi3O8), albite (NaAlSi3O8), and anorthite (CaAl2Si2O8).
Orthoclase is also known as alkali feldspar, and albite and anorthite are known as plagioclase
feldspar.
At
very
high
temperatures there is a complete
solid solution between albite and
anorthite.
Solid
solution
in
plagioclase feldspars is rather more
complex than that in olivine because
some of the substituting ions (Na+
and Ca2+) have different charges
(although they do have similar ionic
radii (Na+ = 97 pm and Ca2+ = 99
pm).
A coupled substitution is
required
to
maintain
charge
balance; that is, two substitutions
occur
simultaneously:
Na+
2+
substitutes for Ca
and Si4+
substitutes for Al3+. Like quartz,
the framework of feldspars is not
too tightly packed so feldspars do
not have high densities (2.55 –
2.76). The rigid framework does,
however, make them quite hard (6 –
6.5). Cleavage is quite well
developed as a result of weaknesses
introduced by ionic substitution.
5
IGNEOUS PETROLOGY
Igneous petrology is the study of igneous rocks. Igneous rocks form by the crystallisation of
molten material called magma which comes from inside the Earth. Since igneous rocks form
from liquids their crystals can grow freely in any direction so their crystals usually show
random orientation. Magma is not just molten rock. In addition to the material which will
eventually form rock, magma contains a large number of volatiles. Volatiles are substances
which vaporize easily. They include water, carbon dioxide, hydrogen sulphide, hydrogen chloride,
hydrogen fluoride, sulphur, sulphur dioxide, chlorine and fluorine. Water is, by far, the major
volatile and it may constitute as much as 8% of a magma.
MAGMA – THE HOT STUFF
Magma is the volcanologist‘s raw material, the molten material that is ultimately erupted at the
surface in lava flows or pyroclastic eruption columns. Magma is not synonymous with lava.
Magma is an elusive term, difficult to define succinctly, but it is best regarded as fresh, mostly
molten rock, still vitalized with volatiles that it acquired in its source region.
Many changes affect molten rock during its transformation from underground magma to
surface effusion. Three components, or phases, are usually present in magma. First, a viscous
silicate melt; second, a variable proportion of crystals; and third, a volatile or gas phase. Each
of these phases influences the way in which the magma erupts at the surface. When subjected
to subtly different eruption mechanisms, a single magma may give rise to startlingly different
eruption products.
The melt phase is the most complex component. Some of the physics of magmas is still
uncertain – it is impossible after all to get into a magma chamber with experimental
instruments. Molten silicate rock is physically different from the liquid water that results when
ice is melted. Pure water consists only of simple molecules of H2O, and therefore ice melts (and
freezes) at atmospheric pressure at a single temperature: a temperature so sharply defined
that it forms the starting point of the Celsius temperature scale: 0°C. A molten magma, by
contrast, is chemically complex, consisting of silicate molecules in which a wide range of
elements is combined. This complexity has two consequences for volcanology: the melt does not
consist of free molecules, but is polymerized, and it does not have a single, clear-cut freezing
point.
Polymerization describes the way that molecules cluster to form larger complexes by repeated
linking of the same molecular groups, retaining an underlying chemical identity. It is a common
phenomenon. An everyday example of a polymer is polyethylene, composed of CH2=CH2 molecules
linked to form endless chains. Although less familiar, silicate polymers are similar. They are
important in volcanology because they affect the viscosity of a melt, and hence the way in
which it erupts. Polymerization in silica-rich magmas (rhyolitic) is due to the strong bonds that
exist between the silicon and oxygen atoms which form networks of inter-linked tetrahedra.
6
Silica-rich magmas contain more silicate tetrahedra, and so are more highly polymerized and
therefore more viscous than silica-poor magmas (basaltic).
To reduce the viscosity of a silicate melt, the silicon-oxygen bonds must be broken. One way of
doing this is to add water, which, by forming OH- ions, breaks the bonds and causes
depolymerisation (the same happens when water is added to glycerine). Conversely, adding
carbon dioxide to a melt may help polymerization, and therefore increase its viscosity.
Three factors influence the melting (and freezing) temperature of silicate materials:
composition, pressure and volatile content. Moving from basaltic compositions towards rhyolitic
compositions, there is a marked drop in melting temperature. At high pressures, a silicate of
fixed composition will melt at higher temperatures than at lower pressures. And a ―wet‖ silicate
(containing lots of volatiles) will melt (and freeze) at lower temperatures than a dry one. All
these variables, coupled with polymerization, mean that when a rock is heated, it does not begin
to melt at a fixed temperature and carry on melting at the same temperature, like ice, until it
is molten. Rather, for a given set of conditions (composition, pressure, and volatile content), it
will begin to soften at a certain temperature, and become progressively more molten as
temperature increases until at some higher fixed point it is entirely molten. These
temperatures increase with pressure, so on a temperature-pressure graph, two lines maybe
plotted. One marks the beginning of melting (the solidus), and the other marks the end (the
liquidus). Exactly the same applies in reverse when a melt cools.
Fortunately, volcanologists need only consider melts at or near the surface, so pressure is not a
major variable, and diagrams such as the one above can be safely left to petrologists. For most
of the discussion that follows, we can assume that the eruption temperatures of common lava
compositions are as shown in the table below.
7
Rock Type
Rhyolite
Dacite
Andesite
Basalt
Temperature (°C)
700-900
800-1100
950-1200
1000-1250
The second phase involves a melt with a variable amount of crystal formation. Many magmas
have begun crystallizing long before they are erupted. Thus, lavas erupted at the surface often
contain abundant phenocrysts. Typically, they are only millimetres across, but exceptionally
they may reach several centimetres. Phenocrystal minerals by their very nature are those that
crystallise out at the highest temperatures. In a basalt they are therefore typically olivine and
augite, while in rhyolites they are more commonly feldspar. Sometimes, however, so much
crystallisation takes place that olivine crystals sink downwards through the melt, accumulating
at lower levels to form rocks crammed with olivine. These lavas, of course, have compositions
very different from the original liquids that they crystallised from.
Study of phenocrysts provides useful volcanological insights. Phenocrysts begin crystallisation
before eruption, and so they may have complex histories. Plagioclase feldspar phenocrysts
often exhibit spectacular zoning. Zoning is described as normal when composition changes from
more calcium-rich to more sodium-rich towards the edge of the crystal, reverse when the
opposite is observed, and oscillatory when the composition varies erratically from one to the
other. These mineralogical variations can be used to track the evolution of physical conditions
within the magma chamber, showing that pressures have varied abruptly over short periods of
time, perhaps in response to surface eruptions. Other phenocrysts have mineral compositions at
variance with the composition of the lava containing them, suggesting that magmas of different
compositions must have mixed together. Phenocrysts derived from an alien source are termed
xenocrysts.
The third phase is a gas (volatile) phase. Wafts of pungent fumes are sure signs that you are
approaching a volcanic vent. Even if it is not erupting, a volcano may release thousands of tonnes
of sulphur dioxide every day. Obnoxious smells apart, gases play a dominant role in the eruption
of magmas. Determination of the amounts and compositions of gases present in a magma is
tricky, not only because of the physical difficulty of sampling the hot raw material but because
some gases which are stable within the magma react chemically the moment that they are
exposed to the air. Sulphur dioxide (SO2) is the most easily recognised volcanic gas (smells of
rotten eggs), but steam and carbon dioxide are more abundant. Hydrogen, hydrogen chloride,
hydrogen fluoride, hydrogen sulphide, carbon monoxide, and several other gases have also been
detected.
There is a marked range in volatile content from basalts to rhyolites. Typical ocean ridge
basalts often contain less the 0.5% water, whereas a rhyolite may contain 4 – 5%. Just as the
carbon dioxide used to give Coke its fizz is more soluble at high pressure than low, so magmatic
gases are more soluble at high pressures than low. Thus, gases exsolve from a magma as it is
depressurised near the surface, with explosive results. Gases are also more soluble at low
temperatures than high, although this has a less marked effect in terms of eruptive behaviour.
8
Volatiles diffuse relatively easily through magmas and are of low density; therefore they
congregate towards the top of the magma chambers, forming volatile-rich caps. These volatilerich layers are necessarily the first to erupt, a fact which explains why the early stages of
eruptions are often more violently explosive than the later stages.
Viscosity is one of those concepts which can be usefully banded about conversationally, but
which are mine fields of complexity when used quantitatively. In everyday terms, viscosity
describes the sluggishness of fluid; its resistance to being stirred. Notice the term fluid,
rather than liquid: the same principles apply to liquids, gases, pyroclastic flows and aerosols.
Viscosity is a crucially important parameter in many volcanic processes. It can be defined as the
internal resistance to flow by a substance when a shear stress is applied. A simple fluid like
water flows in response to the slightest stress; it has a low viscosity. When a glass of water is
tilted, the water instantly responds by flowing to find its new level. Honey is more viscous;
when a jar is tilted, it takes far longer to find the new level. In both cases, once flow has
started, the rate of flow (or strain rate) is directly proportional to the applied stress. Such
fluids are termed Newtonian fluids, after Sir Isaac who first described them.
More complex substances behave differently. At low stresses, they do not flow at all; they
appear to be solid. Once a certain minimum stress has been exceeded, however, they begin to
flow and at higher stress levels behave like Newtonian fluids, their rate of flow varying in
direct proportion to the applied stress. The initial stress required to make a fluid commence
flowing is its yield strength. Fluids which
exhibit a yield strength are termed
Bingham substances. Many fluids exhibit
behaviour
intermediate
between
Newtonian and Bingham; they do not
possess a definite yield strength, and
show a non-linear relationship between
shear stress and strain rate. They are
termed pseudo-plastic. Toothpaste is an
example; a small squeeze of it retains its
shape on the brush, but a large volume
deforms under its own weight! While a
few exceptional basaltic lavas approach
Newtonian behaviour, most have pseudoplastic, toothpaste-like qualities. For
simplicity, however, we can regard most
lavas as Bingham substances.
Like many other volcanological parameters, magma viscosity is exceedingly difficult to
determine directly. Some measurements have been made on lavas, however, by hardy souls
sticking instruments into them. Typical basaltic lavas have viscosities of between 102 & 103 Pa s
(Pascal second; kg/m/s; 1 Pa s= 10 poise). To put this in context, pure water has a viscosity of
about 10-3 Pa S, so these basalts are thousands of times more viscous than water. (Olive oil is
about 100 times more viscous than water.)
9
Fluids become more viscous as they get colder –
this is why automobile engines require
different grades of oil in winter and summer.
This is because when a liquid is heated the
cohesive forces between the molecules reduce
thus the forces of attraction between them
reduce, which eventually reduces the viscosity
of the liquids. In magmas, the effect is
dramatic – for rhyolitic melts, viscosity
increases by more than eight orders of
magnitude between 1300°C and 600°C. Basalt
shows a similar trend, although not continued as
far - basalts are mostly solid below 1000°C.
The graph at the side of the page makes an
important additional point: although melts of all
compositions become less viscous at higher
temperatures, rhyolitic melts are always more
viscous than basaltic ones at the same
temperature.
Relationship between water
content and viscosity at
different temperatures on a
rhyolitic magma.
Dissolved water also has an important effect on
magma viscosity because of its ability to render
magma less polymerized by breaking siliconoxygen bonds. This effect is illustrated in the
graph to the left: at 1000°C, the viscosity of a
rhyolitic melt is decreased by many orders of
magnitude when the dissolved water content is
increased. Basalt shows a similar, but less
marked decrease in viscosity. Again, the graph
below shows that rhyolitic magmas are more
viscous than basaltic ones at the same
temperature and with the same water content.
Relationship between water
content and viscosity at
different temperatures on a
basaltic magma.
10
Other phases present in a fluid also affect its viscosity. Solid materials such as phenocrysts
naturally increase viscosity, but the effects are difficult to quantify because of the wide
ranges in sizes and shapes of the crystals. Gas bubbles, the only other likely components, have
effects which are even more difficult to untangle, since a well-vesiculated magma is a froth
whose flow properties are controlled by factors such as surface tension and the thickness of
bubble walls.
Pumice is an example of a highly vesicular volcanic rock. Vesicles are small gas bubbles which
are found trapped in lavas of all compositions. They are often only a few millimetres across and
thinly scattered, but some lavas are so honeycombed with large bubbles that they resemble
Swiss cheese; more bubbles than solid. Vesicles form when dissolved gases come out of solution
from a magma as it depressurised; just as they do when a bottle of champagne is opened. If the
bubbling gas can escape freely, the lava that results will be degassed or ―flat‖ and the resulting
rock essentially free of bubbles. If, on the other hand, gas is still exsolving while the magma is
erupted, bubbles will continue to grow, and will be found frozen into the lava when it cools. The
ability of exsolving gas disrupting the magma is critical to the understanding of volcanic
eruption mechanisms. Two aspects require our attention: the formation of vesicles, and their
growth.
Formation of vesicles in a magma depends on the amounts of dissolved volatiles such as water
and carbon dioxide, and the vapour pressure that they exert relative to confining pressure of
the magma. Exsolution will commence when the vapour pressure equals or exceeds the confining
pressure, just as steam bubbles begin to grow in boiling water when their vapour pressure
balances the water pressure. Thus, vesicle formation will be initiated at a greater depth
(pressure) in a volatile-rich magma than in volatile-poor one. One way of initiating vesicle
formation, then, is to depressurise a magma, a process physically similar to uncorking a
champagne bottle. This is termed first boiling.
Vesiculation may also occur as a result of a second, more complex phenomenon. When
crystallisation takes place in a cooling magma, removal of crystals from the magma concentrates
volatiles in the remaining liquid, thus driving up their vapour pressure. Crystallisation of
minerals also liberates latent heat of fusion, keeping temperatures high, which in turn causes
vapour pressures to remain high, ultimately causing bubble formation. When runaway bubble
growth is triggered by crystallisation, the process is termed second boiling. Enormous
pressures can be generated in magma by the increase in volume that results, which may have
spectacular consequences.
Once initiated, the growth of vesicles to larger sizes is controlled by the volatile content of
the magma, the rate at which volatiles can diffuse through the magma into the bubbles and
other intrinsic variables such as density, viscosity, and surface tension of the magma. Initial
formation of vesicles has an interesting side-effect: as water vapour separates out and forms
bubbles, the magma increases in viscosity and its yield strength. This process is sometimes
likened to the stiffening that occurs when the white of a raw egg is whisked and bubbles form.
Also as magma nears the surface, where explosive eruptions become possible, the pressures
upon it are reduced and more and more bubbles can than separate out from its still-fluid parts.
11
The effect of these variables is that bubbles are likely to grow in diameters between 0.1 and
5cm in basaltic explosive eruptions, but only 0.001 to 0.1cm in rhyolitic eruptions. The bubbles
so formed not only increase quickly in number and size, but they also rise through the magma
and help to propel the magma more rapidly upwards.
If the volatile content is so high that many bubbles develop, and thus make the magma very
vesiculated (frothy), then the stiff films of magma forming the walls of the bubbles are
thinned and weakened so that the bubbles can expand – often greatly increasing the volume of
the magma as a whole. When the magma erupts, sudden surface decompression enables the
bubbles to burst in violent explosions that shatter the magma into small fragments that then
cool quickly. Near the surface, basic and acidic magmas behave differently, often because of
their different viscosities and volatile contents. Thus, because the walls of bubbles in fluid
basaltic magmas are relatively weak, they can burst more easily, less violently, and over a longer
period. Moreover, many basalts are so fluid that bubbles can move through them and reach the
surface without any explosions at all. But the properties of acidic magmas often conspire to
generate sudden, violently explosive eruptions, especially where the volatile content is high.
When bubbles become so numerous the whole upper mass of a magma chamber becomes a
frothy, foam-like liquid the volatiles suddenly overcome both the viscosity of the magma and
the reduced crustal pressures. The gases expand rapidly. The cork is blown from the
champagne bottle. The volcano erupts violently. The magma is pulverised into fine, suddenly
chilled lava fragments of ash, pumice and volcanic dust that are thrown by the enormous gas
discharge high into the air in a billowing eruption column.
CLASSIFICATION OF IGNEOUS ROCKS
Such considerations of viscosity alone may help explain the differences between volcanoes, but
we also need to enquire why there are different magma types in the first place and what the
ultimate origin of magmas may be. This involves studying phenomena at greater depths. Within
the Earth and considering how other igneous rocks are related to the lavas. There is often a
connection between magmas erupted as lavas and those which consolidate below ground. Dykes
are often made of igneous rock which is chemically similar to the lavas they intrude into, but
which differs from them in that the crystals had longer to form and are, therefore, larger.
The Cuillin Hills on the Isle of Skye in Scotland are composed of a very coarse grained igneous
rock known as gabbro. No one can observe such an intrusive rock being formed. It is believed to
have cooled considerably more deeply in the crust than a dyke, so it cannot be so easily linked
with surface volcanic activity. There are, however, so many chemical and mineralogical
similarities to lavas being erupted today that it is natural to seek some links. Obviously,
measurements of magma temperature, gas content and rate of flow are quite impossible once
the rock is formed, so we must turn to other criteria. In doing so, we hope to develop a method
of classification which, if it is to be of any value, must cover the whole range of igneous rocks,
from ancient to modern. Any system must be flexible enough to fulfil a wide range of needs.
Geological investigations may include the examination of hand specimens in the field, detailed
study of thin slices of rock beneath the microscope, chemical analyses of crushed rocks or of
extracted minerals, as well as observations of actual volcanic processes. Many of the products
12
of igneous activity have economic uses, so here is another need for careful definitions. You will
also be painfully aware of the fact that you are often expected to identify odd lumps of rock
straight from the drawer, without reference to any helpful aids such as knowledge of the field
relationships, rock chemistry or even thin sections for the microscope. Inevitably, any one
system of classification is bound to be in the nature of a compromise, which will need
adaptation by the various specialist interests, but the basic framework set out on the next few
pages has gained general acceptance among geologists.
SiO2
Al2O3
Fe2O3
FeO
MgO
CaO
Na2O
K2 O
TiO2
H2O
50.0
16.0
2.0
7.0
8.0
12.0
2.5
0.5
1.5
0.5
46.0
15.0
4.0
8.0
9.0
9.0
3.5
1.5
3.5
0.5
60.0
17.0
2.0
4.0
3.5
7.0
3.5
1.5
0.5
1.0
Peridotite
Rhyolite
Andesite
Alkali
basalt
Tholeiite
basalt
Component
Many differences between lavas are due to variations in their chemistry, which provide one
possible basis for classification. Although at one time laborious, it is now technically feasible
for a well-equipped laboratory to measure, quite speedily, the chemical composition of a
crushed sample of igneous rock. Results are usually expressed in terms of the oxides of the
various elements present, not because they occur in this form but because this is how they
were traditionally processed as part of the operation.
73.0
13.0
0.5
1.5
0.5
1.5
4.0
4.0
0.5
1.5
43.5
4.0
2.5
10.0
34.0
3.5
0.5
0.3
1.0
0.7
Table to show compositions of some igneous rocks (average weight %)
The table above gives the results of analyses of some typical igneous rocks. The importance of
silica content of a lava has already been stressed and it is this which has been chosen as the
main basis of chemical classification of all the igneous rocks. Many years ago, arbitrary
divisions were chosen and the spectrum of igneous rocks divided into four categories. The
names of each of these were based upon a misconception of the actual properties of silica, but
the names have remained, in spite of attempts to find better ones. The categories are shown in
the table below.
Name of category
Acidic
Intermediate
Basic
Ultrabasic
Percentage of SiO2 in bulk chemical composition
66% SiO2 and above
52% - 66% SiO2
45% - 52% SiO2
45% SiO2 and below
13
The main advantage of this system is that it enables the geologist to relate the solidified rock
to its probable magmatic source. For example, the lavas of oceanic ridges look very similar to
those of ocean hot spots and yet there are differences in their chemistry which reflect their
different origins (i.e. tholeiitic-basalt and alkali-basalt in the table showing the chemical
compositions of some igneous rocks. On the other hand the lavas, dykes and deep seated
intrusion already referred to look very different from each other and yet analyses of them
show that they are chemically almost identical.
70%
52%
The main disadvantage is one of inconvenience. We need names to apply to specimens as we find
them, without having to await the laboratory‘s report. Another complication is the rather
arbitrary way in which the divisions were drawn up in the first place. After many thousands of
analyses, we know that the division at 66% SiO2 coincides with a minimum in the frequency
curve of all the rocks analysed, so it is quite well chosen. However, the other two dividing lines
are badly selected, with the 52% line actually coinciding with a peak of abundance of rock
types! The second most frequent igneous rock is one with just over 70% SiO2, which is in the
middle of the acidic category.
In spite of these problems, the concept of a range of rocks from ultra-basic to acid is of
considerable value, although in practice the label is more frequently derived from other criteria
related only approximately to the silica percentage. Variation in other chemical components is
also important, yet it is not expressed by this method.
Rocks may be grouped on the basis of their chemistry after laboratory analysis. By contrast,
one of the most obvious properties of a rock which can easily be recorded is its colour. Igneous
rocks range from pale grey or white through to black. This often, although not always, reflects
significant differences in rock chemistry or mineral content and it may be used as a rough basis
for classification. The terms used are based on Greek words and are as follows:
Light coloured
Medium coloured
Dark coloured
14
– leucocratic 0 – 30% dark minerals
- mesocratic 30 – 60% dark minerals
- melanocratic over 60% dark minerals
The method provides a useful basis for rough divisions but it reveals nothing about the genesis
of the rock. Also, whilst most of the melanocratic rocks are basic or ultrabasic types there are
some acidic rocks which are black. A good example is the shiny black volcanic glass, obsidian.
More recently, other terms have entered common usage in an attempt to describe concisely the
dominant types of minerals present in the rock. Thus felsic describes a rock in which feldspar
and quartz (silica) are major constituents. Mafic is used for rocks which contain some feldspars
but which are rich in the ferromagnesian minerals (i.e. magnesium and iron (Fe) bearing).
Ultramafic labels a rock which is almost completely made of ferromagnesian minerals alone. The
felsic minerals are mostly light-coloured and the mafic ones dark, so for most practical
purposes the terms have come to be applied on the basis of the overall colour of the rock, and
are often loosely regarded as alternatives to leucocratic, mesocratic and melanocratic.
When describing the colour of a rock in the field a colour index can be used. This is based on
the estimate of the total percentage of mafic minerals present. It is well worthwhile using this
diagram in your fieldwork as there is a common tendency, because of the greater visual impact
of dark minerals, for colour index to be overestimated in field reports. The accuracy with
which the colour index can be estimated will depend on the grain size of the specimen.
At AS Geology you have seen and handled good specimens of a range of minerals. In doing so,
you have learned that most minerals have clearly defined properties which enable them to be
identified. The majority of igneous rocks are composed of minerals which have crystallised
tightly packed together. Identification of them is not as easy as it is for individual mineral
specimens but with a little practice it can be done, either in the hand specimen, or with a thinsection of the rock beneath a petrological microscope, or both. The relative abundance of each
mineral can also be estimated. Because of the wide range of chemical constituents comprising
rocks from acid to ultrabasic, there are an equally diverse number of minerals which may be
present. Here then is a potentially useful way of establishing divisions, both in the field and in
the laboratory.
Over 3,000 minerals are known, but, fortunately for the petrologist (person who studies rocks),
most of these are uncommon! In practice, the majority of igneous rocks can be described in
terms of a dozen or so which are usually referred to as the rock-forming minerals. The main
groups of minerals found are quartz, the feldspars, the micas, ferromagnesian minerals such as
the olivines (e.g. forsterite and fayalite), the pyroxenes (e.g. augite), the amphiboles (e.g.
hornblende), and the iron ores, notably magnetite.
15
The diagram on the next page has been constructed empirically from the results of many
thousands of mineralogical analyses of igneous rocks. It shows the variation in the proportion of
each of the rock-forming minerals and how this is related approximately to the chemical
division described before. The minerals listed are usually referred to as the essential minerals
and it can be seen that there is only a handful of such minerals in each category, e.g. basic
rocks consist essentially of plagioclase feldspar and a pyroxene such as augite. Other minerals
are commonly present, but usually in smaller quantities and these are known as accessory
minerals. Examples would include sphene, apalite or zircon in acid rocks. In the basic category,
olivine, although listed in the table below, is often regarded as an accessory mineral. When
appropriate, its presence is indicated by a hyphen, including the word as a prefix to the rock
name, e.g. olivine-basalt.
In identifying igneous rocks from their mineral content, it is usual to approach them
systematically:
1.
Is quartz present and if so in what proportion? The table above shows that acid rocks like
granite contain between 5% and 40% quartz, intermediate rocks between 0% and 10% and
basic rocks like basalt 0%. When quartz does appear in basic rocks it is known as an
accessory or a xenocryst). You must not confuse the % of the mineral quartz in the rock
with the % of silica in its bulk chemistry. Both share the formula SiO2, yet in chemical
analysis the SiO2% is the total of all the silica, occurring in its free state and in the
formation of other silicates minerals, e.g. olivine (Mg,Fe) SiO4. Thus basalt (basic/mafic
rock) may have a SiO2 content of 50% and yet have no free quartz because all silica is used
up in other minerals.
2.
What is the feldspar content and which feldspars are present? The proportion of the rock
which is composed of feldspar can usually be estimated in the hand specimen, as can the
distinction between potash feldspar (orthoclase) and the plagioclases. The determination
of the variety of plagioclase, however, is normally only possible with the aid of a
16
petrological microscope. A slice of rock is ground down until it is so thin that most of its
minerals are transparent. This is examined beneath a microscope equipped with polarised
light and a rotating stage. The optical properties of minerals in such thin sections are quite
characteristic and it is not difficult to identify them. It is important to be able to work in
this detail, since the classification partly depends on the type of plagioclase, ranging from
the calcium-rich types (anorthite) in the basic rocks to the sodium-rich (albite) in the acid
ones. Potash feldspar (orthoclase) and the albite variety of plagioclase feldspar are
commonly referred to as the alkali feldspars.
3.
Are micas present and if so which ones? The white mica, muscovite, is most often found in
the acid rocks and helps to give them their lighter colour. Biotite is often present in acid
rocks but is most abundant in the intermediate ones.
4.
Which ferromagnesian minerals are present and in what proportions? The ultrabasic rocks
are mainly composed of the darker, denser ferromagnesian minerals and it is perhaps in
this category that distinguishing between them is of greatest importance. The mineral
table above shows the most common ferromagnesian minerals in each group, and it also
shows that some minerals are mutually exclusive. For example, it would be very unusual to
find both olivine and quartz in the same rock; if sufficient silica is present to form free
quartz, then olivine crystals would not have survived for long in the magma, but would have
reacted with the excess silica to form an amphibole (e.g. hornblende) or pyroxene (e.g.
augite).
It must be evident from the above detail that the mineral content of a rock provides a
powerful basis for its classification. Nonetheless, there are some drawbacks. It does not
distinguish between rocks formed at different depths, thus fine-grained rocks often have to
await sectioning and microscope examination before a positive identification can be made. It is
also easy to ignore possible genetic relationships between rocks by putting them in neat
categories. For example, small quantities of acid lavas are found in association with basic ones
and may represent later derivations from the same magma. The fourth area of classification of
igneous rocks is its texture.
By the word texture we mean the grain size of the rock and the relationship between
neighbouring mineral constituents. When describing the texture of a rock in the field the
following need to be examined; the grain-size (reflects rate crystallisation); the fabric
(arrangement and shape of crystals); and homogeneity (whether the rock is uniform and
equigranular or banded and porphyritic). Magma cools by losing heat to its surroundings. The
rate at which heat is lost from a body of magma depends upon; the ratio of its surface area to
volume; the temperature of the magma; and the thermal conductivity of the materials around it.
We shall examine some particular textures later but it is sufficient here to note the
relationship between the size of the crystals in an igneous rock and the time which it took to
cool. Generally, a slower rate of cooling produces bigger crystals. In geological terms, slower
cooling rates are achieved deeper in the crust than at the surface, so the use of grain size, as a
criterion for classification, will usually reflect the geological environment in which the rock
crystallised. In the past, such principles were probably taken a little too far, and the terms
17
volcanic, hypabyssal and plutonic were applied to the hand specimen to indicate, respectively,
whether the rock was produced at the surface, at moderate depths, or deep in the crust. Such
ideas are most useful in the field, where larger scale structure may also be observed, but it is a
little risky in the hand specimen. For example, not all fine-grained basalts are lavas; basalt quite
frequently occurs as dykes, intruding through strata.
In practice, the most useful classification system is one which combines the mineral content of
a rock with its texture. The approximate relationship between mineral content and chemical
composition has already been shown. In the table on the next page, which shows an outline
classification of igneous rocks, texture has been plotted against the approximate chemical
categories only, for sake of simplicity. It must be stressed that such a fitting of rocks into
pigeonholes is a convenient but rather artificial way of seeking order in the natural world. All
geologists would agree that rocks do not fall into neat slots but form parts of a continuously
varying spectrum.
18
19
Since so much of our understanding of igneous rocks depends upon microscope examination of
thin sections, a series of drawings of such sections follows. The minerals have been shown by
partly stylised symbols, which approximate more closely to the views seen in plane polarised
light.
IGNEOUS ROCK FORMATION
Igneous rocks are those that have solidified from hot, molten material called magma. Magmas
have densities lower than those of the rocks above them, so they rise through passageways and
zones of weakness. As they do so they lose heat at their edges to the country rock (the
surrounding sedimentary rock), bringing about crystallisation.
The vast bulk of magmas crystallise at depth; only a small proportion reaches the land surface
or erupts as volcanic lavas and ashes. Magmatic material congealed at depth is collectively
referred to as intrusive. Uplift and erosion leads to the exposure of such intrusive rocks to the
surface.
To create magma you have to melt rocks. Rocks melt when the temperature is higher than the
melting points of minerals in the rock. It becomes entirely molten (liquid) if temperature is
higher than a melting point controlled by all the minerals in the rock.
Factors affecting the temperatures minerals melt at include;
 Pressure - the greater the pressure the higher the melting point of a mineral, and higher
temperatures are needed to melt it.
 Gas Content - the effect of pressure changes if enough gas (especially water vapour) is
present. As water pressure increases the melting point of minerals decreases, e.g. water
mixed with granite lowers its melting point from 900°C to 625 °C.
 Neighbouring minerals - some minerals melt at lower temperatures when mixed together than
they do when alone. e.g. pure quartz needs temperatures well over 1500°C to melt, but when
mixed with feldspar in a 42% quartz to 58% feldspar mixture it only takes a temp of 1000°C
for both minerals to melt.
LOCATION OF MAGMA PRODUCTION
Active volcanoes testify to the existence of bodies of magma at depth. This does not imply
that the entire interior of the Earth is molten. Seismic evidence shows that the Earth is solid
down to the outer core (2900 km). Magma production must be a localised phenomenon. The
melting of rock at depth to form magmas occurs in a number of distinct belts:
 Mid-oceanic ridges (constructive plate margins). Here the partial melting of mantle rocks
(peridotite) as the lithosphere thins forms basaltic magmas, e.g. Iceland.
 Rift Valleys (newly forming constructive plate margins). Here the partial melting of mantle
rocks (peridotite) as the lithosphere thins forms basaltic magmas, e.g. East African Rift
Valley.
20
 Subduction zones (destructive plate margins). Here the partial melting of subducted oceanic
lithosphere (basalt) forms andesitic magmas, e.g. Andes.
 Hot spots (intraplate areas). Here the partial melting of mantle rocks (peridotite) forms
basaltic magmas, e.g. Hawaii.
MAGMA PRODUCTION
In all the locations noted above it is the process of partial melting that forms magma. Partial
melting is the process whereby the melting of solid rock is incomplete. For example, from the
chart on the next page it can be seen that mafic rocks (basalt) at the surface (0.1 MPa
pressure) start to melt at 1075ºC (solidus curve) and are thus partially molten up to 1200ºC,
and completely molten after 1200ºC (liquidus curve).
As the rock begins to melt each mineral contributes to the magma at a different rate. The
more acidic minerals (quartz, muscovite and orthoclase feldspar) melt first and hence make up a
large proportion of the newly generated magma. Partial melting will then always produce a more
acidic magma (magma with higher silica content) than the original rock. For example, the partial
melting of ultramafic rocks (peridotite) forms mafic magmas which when solidified form mafic
igneous rocks such as basalt (if cooled at the surface) or gabbro (if cooled deep underground).
The partial melting of mafic rocks (basalt) forms intermediate magmas which when solidified
form intermediate rocks such as andesite. If a rock becomes completely molten the new magma
would retain the same composition of the original rock when solidified.
Rocks melt due to three reasons:
 An increase in temperature is the most obvious mechanism to cause rocks to melt, but rarely
occurs on its own.
21
 A decrease in pressure is a less obvious mechanism to induce melting. In the diagram above if
you take a rock at 900 MPa and 1100ºC it is solid. However, if this rock was to keep the same
temperature and decrease its pressure by taking it to the surface it would melt.
 Rocks can also melt to form magma if water or water vapour is added. Water lowers the
melting point of a rock and therefore induced melting without an increase in temperature.
This can be seen in the graph below, for example a dry rock at 900 MPa and 1100ºC will be
solid but if water is added its melting point is lowered to 650ºC and will melt.
The graph below shows the normal situation found in the lithosphere over most of the Earth.
The graph shows that temperature increases with depth at an average rate of 25 ºC/km (known
as the geothermal gradient) for the first 100 km or so and then this tapers off to a lower rate.
You can see that the temperature is not high enough to melt the rocks in the lithosphere or the
asthenosphere because the geotherm line does not cross the solidus curve (temperature at
which rocks start to melt). However the geotherm line is closest at about 125 km depth. Here
rocks are starting to get closer to their melting points and will become more mobile as one or
22
two crystals begin to melt. This is where the asthenosphere is situated and explains why it is
5% molten and not brittle and solid. However it is not magma!
It is only in certain places around the world, such as spreading centres, subduction zones etc.,
where the mantle is induced to melt and form magma. The reasons for this are outlined below:
1). Rift Valley.
Lithosphere is stretched and thinned allowing the asthenosphere to flow into the space
provided. In rising nearer to the surface where pressure is reduced, without losing its
temperature, the mobile asthenosphere crosses the melting point curve and starts to melt. This
process is known as decompression melting and accounts for the production of magma in
continental rift valleys.
Note that the geotherm only just crosses the melting curve so only a relatively small amount of
magma is produced.
2). Mid-Oceanic Ridge.
This is an extreme case of lithospheric extension, where the asthenosphere is allowed to rise
almost to the surface. As two oceanic plates pull apart the lithosphere is stretched and thinned
allowing the asthenosphere to flow into the space provided. In rising much nearer to the
surface where pressure is reduced, without losing its temperature, the geotherm crosses the
melting curve to a much greater extent than in continental rift valleys, where the crust is much
thicker. So at mid-oceanic ridges the asthenosphere retains its higher temperatures but at the
23
now lower pressures the melting point of the mantle rocks are easily exceeded and large
amounts of magma are produced. This process is again known as decompression melting.
3). Hot Spots.
Melting of the asthenosphere occurs due to convective plumes of anomalously hot mantle rising
from great depths (probably mantle-core boundary). The average upper mantle temperature is
1300 ºC, which at this depth and pressure is too low to melt. However, thermal plumes raise the
temperature by 300ºC which is enough to cross the melting curve for peridotite and the mantle
melts. Decompression melting at the top of these mantle plumes produces magma in places like
the Hawaiian Islands in the Pacific Ocean. The exceptionally vigorous volcanic activity in
Iceland is thought to be due to the superimposition of a mid-oceanic ridge on a hot mantle
plume!
The superimposition of a continental rift system on a hot mantle plume similarly produces large
volumes of basaltic magma.
4). Subduction Zones.
Here the lithospheric plate is returned to the asthenosphere. This ought to be the least
volcanic areas on Earth because subduction takes the cold rocks of the ocean floor down into
the mantle. These are however some of the most volcanically active areas in the world! Why?
Along with sea floor, subduction takes enormous volumes of seawater down into the mantle. The
water has the effect of lowering the melting point of the mantle by 400°C and causing it to
melt. This process is known as hydration melting.
24
From the diagrams and notes above it can be seen that the partial melting of solid rock to form
magma can occur due to either, a rise in temperature, a change in pressure or the addition of
water.
MAGMA MIGRATION
Once melting has occurred how can magma collect into a body and rise upwards through the
crust? Partial melting begins by producing a small amount of melt by the preferential melting of
the starting materials' most silica-rich minerals. The first parts of a rock to melt will be at the
edges of individual mineral grains. This forms a liquid (silicate melt) at the grain boundaries
within the still mostly solid rock.
This silicate melt is less dense than the rock of the same composition, and obviously less dense
than silica-poor rocks. For both these reasons, melt that has begun to form by partial melting is
buoyant and so has a tendency to rise upwards. However, melt is slowed down by its own
viscosity (the more viscous the melt, the harder it is to
Melting at grain boundaries flow) and the resistance offered by the surrounding rock.
If there is a fracture available (perhaps because the crust
is under tension), then magma can escape up it in a matter
of a thousand years or less. On the other hand, if there is
no easy pathway, the melt must collect into a body many
kilometres across before it is capable of forcing its way
upwards by means of its own buoyancy. This would require
about 30% of the source rock to melt, and take in excess of
1 million years!
As magma rises to shallower levels within the crust, it will lose heat to the surrounding rocks.
Whether or not this cooling causes it to solidify depends on its starting temperature and the
rate at which it loses heat as it rises. A magma body in the crust will usually keep rising until it
has become almost entirely solid. Even a "crystal mush" consisting of over 90% crystals and less
than 10% melt can continue to rise. If a magma body solidifies completely without reaching the
surface, it will form an intrusive igneous rock, whereas if it is erupted at the surface it
produces a volcanic rock (extrusive). Igneous rocks that have been intruded deep in the crust
commonly occur as bodies up to 30 km across and are referred to as plutons. When a number of
plutons merge together the intrusion becomes known as a batholith. There are a number of
mechanisms that allow plutons to rise further and further up into the crust, these are known as
emplacement mechanisms.
1). Diapiric Emplacement
Plutons at a depth of 20 km or so tend to consist of a number of elongated masses of magma.
They also tend to be concordant with the country rock, which at this depth tends to be
regionally metamorphosed rocks. The explanation for this structure is that at this depth the
country rock is hot enough (and therefore deformable enough) for a pluton-sized mass of
magma to force its way upwards by pushing the country rock aside. A pluton that rises in this
way is described as a diapir. A granitic diapir could rise from the lower crust to about 10 km
25
below the surface in about 100,000 years to 1,000,000 years, but at shallower depths the crust
is not deformable on a sufficiently short time-scale to allow further diapiric rise.
2). Dyke Ascent
Plutons emplaced at shallower levels in the crust (<10 km depth) can usually be seen to be
discordant. Their edges cut across the fabric of the country rock, which at this depth is usually
sedimentary bedding. When intruding into cold, brittle crust, a shallow-level pluton cannot push
the country rock aside and force its way up diapirically as may sometimes the case deeper
down. Nor can it simply "melt its way upwards" by assimilating all the country rock in its path,
because to do so would require more heat than the pluton contains. The most likely means of
accommodating a large magma body at shallow depth is by magma rising up near-vertical
fissures, a process known as dyke ascent. Continuous magma ascent through a 5m wide fissure
could supply an average-sized granitic pluton (5000 km³) in about 100,000 years.
The diagrams above show a ring fracture forming and the central block subsides into the
magma and the magma squeezes up the sides of the block, acting as a lubricant helping the
block to sink. The magma then occupies the cauldron so formed. These types of intrusions are
26
relatively common and good examples can be seen in Scotland. In fact the process can be
multiple with the cauldron sinking repeatedly to form a series of ring dykes, such as in Glen Coe
and Ardnamurchan.
3). Magmatic Stoping
This is an additional process that allows a pluton to slowly rise up through the brittle upper
crust. The heat of the pluton will often crack and fracture the cold and brittle country rocks
surrounding it. These fractures weaken the
overlying country rocks which subsequently
break away and sink into the pluton. This is
the process of magmatic stoping (yes stoping and NOT stopping!) and makes space
for the pluton to rise into. A chunk of
country rock enclosed within an intrusion is
called a xenolith (xeno, pronounced "zeno", is
Greek for "foreign", and xenoliths are
"foreign rocks" in the sense that they do not
belong in the intrusion). Xenoliths are
particularly common near the roofs and walls
of intrusions, where they tend to be angular
in shape. Some xenoliths may sink to the
bottom of the pluton; however they will
usually melt by then. Xenoliths can sometimes be identified near the centres of plutons, where
they tend to have more rounded shapes, and sometimes a recrystallised texture, indicating that
the heat from the surrounding magma was sufficient to cause them to become soft and mushy.
In extreme cases, xenoliths are no more than ghostly dark patches, showing that they have
become almost entirely assimilated into the magma.
IGNEOUS BODIES
1. Batholiths are large, elongated intrusions of coarse-grained plutonic rocks occurring in belts
50-150km in width and 500-1500km in length. They are found in mountain belts and they are
elongated parallel to the mountain ranges. Batholiths are usually composed of a large number of
cross-cutting smaller intrusions, including bodies 5-50km in size with circular outcrops known as
plutons. The igneous intrusions found in large intrusions are described as plutonic because Pluto
was the ancient God of the Underworld. These rocks are coarse-grained because they have
cooled slowly in large intrusions which have been deeply buried.
2. Dykes are the commonest form of minor intrusions. A dyke is a vertical or near-vertical
intrusion that cuts across horizontal or gently-dipping bedding in the surrounding country
rocks. Dykes are therefore discordant to structures in the country rock. Dykes are narrow,
linear intrusions which usually vary from a centimetre to many metres in width, but in general
the average width is probably in the range of 1 - 5 metres. Ring dykes are cylinder-like
intrusions whose diameters usually measure a few kilometres. Ring fractures are probably
caused by the upward push of underlying magma. Following up-doming, the crust stretches and
27
fractures. The block enclosed by the ring fracture sinks and magma rises up the fracture to
feed volcanoes on the surface. Cone-sheets also have near-circular outcrops but they taper
downwards. These cone sheets occupy fractures produced by the upward push of magma in a
chamber under the centre of the sheets. All dykes and cone-sheet intrusions require the
intruded country rocks to be in a state of tension. Dykes are often found in groups that radiate
away from individual igneous centres and are termed radial dykes. Where dykes are
concentrated parallel to a regional tectonic trend they are termed dyke swarms. These dyke
patterns reflect the distribution of tensional force in the intruded rocks. Radial dykes come
from volcanic centres or from batholiths and the dykes occupy fractures produced by the main
body of magma pushing up and cracking the overlying rocks. Parallel swarms indicate tension on
a larger scale. The degree of extension suffered by an area can be found by adding the
thicknesses of the parallel dykes. In the Isle of Arran, Scotland, the extension from the total
thickness of more than 500 dykes amounts to 1.6km.
3. Sills are wide, linear intrusion that parallels horizontal or gently-dipping structures. Sills are
generally concordant between beds of layered rock, but may show small transgressions from
this. Where a sill crosses from one bedding plane to another it is said to be transgressive.
There is great variation in the size of sills, the Palisades Sill at over 300 metres thick being a
particularly large example. One of the best-known British examples is the Great Whin Sill of
Northern England. This averages 30 metres in thickness and underlies a vast area from
Northumberland in the north to the Tees Valley in the South. It is most well-known where its
outcrop has formed a north-facing scarp slope along which Hadrian‘s Wall has been built.
One problem often confronting geologists when they see a layer of igneous rock lying
conformably in a succession of sedimentary rocks is to tell whether it is a sill or a lava flow.
The table below is a summary to help distinguish sills from lava flows in the field. The ticks
indicate the features you might see in cross-section in the field when studying an ancient
volcanic area. The crosses indicate features you would definitely not expect to find.
Chilled margin at top
Chilled margin at bottom
Columnar joints
Concordant top with local discordance
Concordant base with local discordance
Rubbly top and/or bottom
Baking of overlying rock immediately above the contact
Baking of overlying rock immediately below the contact
Weathered surface
Vesicles
Xenoliths
Included fragments
Sill





x


x
()

x
Lava Flow
()
()
()
x
()

x


()
x

Proper chilled margins occur in sills but not in lava flows. However, the glassy top of some kinds
of pahoehoe might be mistaken for a chilled margin, which is why there is a tick in brackets in
28
the lava flow column in the table. Similarly, the base of the lava flow might be chilled against
the underlying rock; therefore chilled margins are not infallible guides. Columnar joints can
occur in thick lava flows as well as sills. Sills are generally concordant, but can be locally
discordant at the top or the base. The base of a lava flow is discordant if there has been
erosion prior to its eruption, but otherwise it is concordant. However, deposits laid on top of
the flow will be concordant everywhere, with no local discordances. Rubbly tops and bottoms
are characteristic of aa flows and blocky flows, and offer one of the most reliable means in this
table of distinguishing a lava flow from a sill. The easiest way to distinguish between a sill and a
lava flow is to observe the baked margin formed. A sill will have a baked area above and below it
but with a lava flow one can only develop in the underlying rocks as there was nothing above it
at the time of extrusion.
Further indications are given by the nature of the upper surface. With a lava flow, this may be
uneven and have a soil developed upon it as it could have been exposed to subaerial processes of
weathering. In contrast, a sill could not show this, as it is by definition intrusive. Also this longterm exposure prior to burial would allow included fragments of the lava flow to be found in the
sedimentarys above. A sill could, however, contain xenoliths of the overlying strata whereas it
would be impossible for a lava flow to have these.
4. A volcanic plug or neck is a cylindrical intrusion representing the conduit which fed material
to the volcano. Erosion of composite volcanoes may reveal these sub-circular exposures of finegrained rock which are generally a few hundred metres in diameter, but may extend up to a
kilometre. Most plugs are composed of lava, but a small proportion consists of pyroclastic
material such as agglomerate (consolidated blocks and bombs).
NATURE OF EXTRUSIVE IGNEOUS ACTIVITY
Magma (molten rock) comes to the surface of the Earth as lava through volcanoes. In a central
vent type eruption the magma rises through a cylindrical volcanic pipe and it leaves the volcano
through an opening called a vent. Eruptions of this type may build large cones. In a fissure
eruption the lava comes from a vertical crack and any cones which form are due to some
sections of the fissure becoming choked with lava and tend to be small.
Volcanoes give off lavas, airborne fragments and gases. The main lavas are basalt, andesite,
dacite and the less common rhyolite. Sometimes very rapid cooling leads to glass formation
(obsidian). The gas in magma tends to escape as the magma rises and pressure on it falls. When
the gas does not escape fully the lava remains full of bubbles or vesicles. Basalt lavas are more
often vesicular than other types especially in the upper parts of the lava flow, where pressure
has been reduced allowing the gas to exsolve. Lava rich in mineral-filled vesicles (amygdales) is
described as being amygdaloidal.
At the time of eruption, basalt lava is at a temperature of about 1000°C. Hot basalt lava flows
very freely but as it cools it becomes more viscous and it solidifies with a ropy-looking surface
as pahoehoe lava or with a rubbly surface as aa lava. This rubbly layer also extends underneath
29
an aa lava as well. Rubbly material from the flow top tumbles down the flow front and is
overridden in a similar way to the rolling caterpillar track on a bulldozer. When a flow that has
been emplaced in this manner is seen in cross-section its interior is usually solid, whereas both
its upper and lower surfaces are rubbly. Pillow lava looks like a pile of closely-packed pillows
each about 1 metre in diameter. Pillows result from lava extrusion under water. As the basalt
lava flows its skin is cooled by the water but the hard skin frequently bursts and a blob of lava
leaks out like toothpaste from a tube to be cooled and hardened into a pillow.
Andesite lava is much more viscous than basalt lava so it tends to form thicker flows and does
not give pahoehoe lava. Instead, andesite flows have rough, blocky surfaces and are sometimes
called block lavas. Generally speaking the more viscous the flow, the larger the blocks. Dacite
lava flows represent a further increment in viscosity over andesite flows. They are so thick
that the term lava dome is used rather than lava flow. Lava domes form above vents as lava
added from underneath pushes up the outer cooled layers. Rhyolites are extremely viscous and
are much less abundant than dacites since magmas of this viscosity usually fragment into
pyroclastic material during an eruption. Rhyolites very often show flow banding.
Solid airborne fragments from a volcano are pyroclasts. Pyroclastic material usually comes from
eruptions in which a gas-rich viscous magma suddenly separates into liquid and gas. The gas
bubbles grow explosively giving a mixture of fragments and gas which is shot from the volcano.
If the gas separates deep in the volcano the volcanic pipe acts like a gun barrel and the glass,
mineral and rock fragments are shot high into the air. The fragments fall back to form a
pyroclastic fall deposit often called tephra. Tephra consists of ashes (less than 2 mm in
diameter); Lapilli (2 – 64 mm in diameter); and blocks and bombs (more than 64 mm in
diameter). Blocks are fragments which are solid before ejection. Bombs are lumps of molten or
partly molten lava which acquire elongated shapes something like lemons! This shape is not due
to aerodynamic drag as the bomb is thrown out of the vent but results from rapid stretching of
ejected fluid magma due to the initial explosion. Bread-crust bombs look like crusty loaves, with
a glassy (chilled) outer and highly vesicular interior. They form when gas continues to expand
and exsolve after the eruption has flung out the bomb into the air. On consolidation, ashes and
lapilli form tuffs while blocks and bombs form agglomerate. Scoria consists of variously sized
fragments of cinder-like, vesicular basic lava. Pumice is again highly vesicular but it is usually
derived from acidic magma.
When gas is suddenly released from magma close to the volcanic vent the material spills out
sideways as a pyroclastic flow which rolls down the side of the volcano as a hot gaseous cloud
similar to the base surge of a nuclear explosion. Pyroclastic flow terminology can be very
complex! What follows is just scratching the surface….. Pyroclastic flows range between two
end-varieties: those that involve vesiculated low-density pumice, and those that involve
unvesiculated, dense clasts. The first variety involving vesiculated, low-density pumice are
called pumice flows and produce ignimbrite deposits (a pumice and ash deposit). The second
type involving unvesiculated, dense clasts are called block and ash flows (or often nuée ardente
– meaning ―glowing cloud‖) which produce block and ash deposits.
30
LAVA FLOWS
Lavas may appear to be merely streams of molten rock presenting few technical difficulties for
volcanologists. Nothing could be further from the truth. Lavas have a wide range of
compositions, basaltic to rhyolitic. Composition apart, their physical properties are also
influenced by their volatile contents, crystal contents, and cooling histories. Thus, flows of
marginally varying compositions may behave quite differently. Only recently has the rheology
(the study of flowing materials) of lava flows come under close scrutiny, so it remains poorly
understood.
What controls the thickness of a lava flow? How far will a flow reach? Questions like these may
seem simple – and indeed they are – but it is still difficult to answer them satisfactorily. The
fluid dynamics of lavas are inherently difficult to study; however, as discussed earlier, magma
viscosity
varies
enormously
with
temperature, composition, and degree of
polymerization. While many measurements
Relationship between angle of
slope and lava thickness.
of lava viscosity have been made by hardy
souls inserting viscometers into active
flows, these are single, point source
measurements. Measuring temperature and
viscosity and their variations across all
parts of a basalt lava at the same time is
much more difficult, and probably
impossible. Andesite and rhyolitic flows
present even more intractable problems.
Superficially, a blob of viscous liquid such
as glycerine may seem to behave like lava
when it oozes slowly over a flat surface.
Over time, however, a blob of glycerine
flattens out to a thin film: it is a
Newtonian liquid, with zero yield strength. Two forces influence the rate at which a blob
spreads: gravity, and viscous resistance. The magnitude of the gravitational force depends, of
course, on the mass of the planet out blob of glycerine is oozing over, but viscosity is a
property intrinsic to the material. Lavas possess a yield strength, largely derived from the
chilled crust which immediately forms on their surface. Before a blob of lava can spread, its
yield strength must be overcome by shear stresses. Thus, lavas never form flat films; they
always retain some thickness, which is a measurement of their yield strength. Fluid basaltic
lavas such as those of Hawaii commonly flow in oozy dribbles as little as a few centimetres
thick; rhyolites, by contrast form flows which are never less than many metres thick.
31
Aspect ratios (thickness: area) for lavas of different compositions. Basalt lavas have ratios less than 1/100;
andesites, dacites and rhyolites plot over broader fields but have lower ratios.
Once a flow gets moving, things rapidly get extremely complicated, because of the large
number of variables. Apart from temperature and composition, lava viscosity is also a function
of the rate of shear. Thus, a basalt flow moving on a steep slope where shear rate is high will
have a lower viscosity than one moving on a gentle slope, other factors being equal. Also, lavas
with lower viscosities will flow faster than those with high viscosities.
It follows from their variations in viscosity and temperature that lavas of different
compositions will form flows with different morphologies. One useful way of expressing
variations in morphology is the aspect ratio. The aspect ratio is the ratio of the thickness of a
flow to the area it covers. Basalt lavas are thin and extensive, whereas rhyolite lavas are fat
and compact. These variations can be expressed on a thickness against area plot, where basalt
lavas occupy a distinctive field, with aspect ratios less than 1/100.
A lava flow begins to erupt from a vent high on the flanks of Mt. Etna or Mauna Loa. How far
will it flow? Will it engulf the bustling village of Zafferana, or incinerate the leafy suburbs of
Hilo? Such questions emphasise that lava flows present problems of more than purely academic
interest. Many factors influence the distance a flow will travel – composition, slope, rate of
cooling, temperature and so on. Other factors being equal, the rate of effusion is also critical.
On the largest scale, this is clearly correct: basalt lavas achieve far higher effusion rates than
silica lavas, and they flow farther; hundreds of kilometres in exceptional cases. In flood basalt
eruptions, effusion rates of thousands of cubic metres per second are possible, but in a typical
andesite eruption maximum effusion rates do not exceed a few cubic metres per second, and
flows rarely extend more than ten kilometres.
Among flows of similar composition, the problem of lava flow length is more complex. While
effusion rate remains significant, other factors are important. The effusion rate is a better
guide to the surface area finally covered by a flow, rather than its length. Since radiation from
32
the surface is the dominant cooling mechanism, the proportion of the lava‘s surface that
exposes glowing, hot lava rather than chilled crust will dictate how quickly the lava cools, and
hence how far it travels. This proportion is in turn affected by slope and topographic
roughness.
Radiation is much more important at high temperatures than low. At moderate to low
temperatures, convection is the chief cooling mechanism. A small percentage decrease in
temperature yields a huge decrease in radiant thermal energy. Like other rocks, lavas are poor
conductors of heat. When the surface temperature of a lava is high, there is a tremendously
rapid radiative heat loss, which cannot be balanced by thermal conduction from the core of the
lava flow. Thus, the crust of an active lava flow inevitably chills quickly as it moves away from
its source.
PROCESSES THAT FORM MAGMAS OF DIFFERENT COMPOSITIONS
The diagram below shows the huge variety of igneous rocks, ranging from basalt through to
rhyolite.
1). Partial Melting
Partial melting is a process whereby melting of a solid rock is incomplete and consequently
produces a melt with a higher silica content than the original rock. For example, if the
ultrabasic mantle rock peridotite were partially melted (rather than fully melted) it would
produce a melt (a liquid) which was basic in composition (i.e. basaltic). If a rock were completely
melted the subsequent melt would retain the same composition of the original rock and
therefore not form magma of a different composition.
33
So how does partial melting produce higher silica content than the original rock? In a solid rock
consisting of more than one mineral, all the minerals begin to melt at the same temperature but
as melting proceeds each mineral contributes to the melt at a different rate until it is entirely
used up. Minerals disappear sequentially in the order shown in the diagram over the page. So if
a rock containing quartz, mica, orthoclase feldspar, plagioclase feldspar and hornblende (an
intermediate rock) were heated the first mineral to melt would be quartz (contributing Si and
O into the liquid melt now formed) then muscovite mica (contributing potassium, aluminium and
more silicon dioxide into the liquid melt) then orthoclase feldspar (contributing more potassium,
aluminium and silicon dioxide into the liquid melt). The resulting liquid would be very rich in
silica and would cool and crystallize as an acidic igneous rock. The solid remains of the original
rock would now be depleted in silica and would be have a basic composition.
Partial melting is exactly the process which forms basalt at constructive plate margins. When
the solid ultrabasic peridotite (made up of olivine, augite & plagioclase) of the asthenosphere
starts to melt it is the augite and the plagioclase which melt first. This liquid (being more
buoyant and less dense then the surrounding rocks) starts to rise leaving behind the solid
olivine rich peridotite. This rising magma consisting of augite and plagioclase is now a basic
magma.
2). Fractional Crystallisation
As magma rises to shallower levels in the crust it will lose heat to the surrounding rocks and it
will begin to crystallise according to the sequence of crystallisation shown in the diagram over
the page. If these first formed crystals are carried along with the remaining melt until it has
completely solidified, the composition of the eventual rock will be identical to that of the initial
magma body, and the first-formed minerals will be phenocrysts. However, if the first-formed
crystals became separated from the melt, the escaping melt would become more acidic than the
original melt, whereas the separated crystals will form a rock more basic than the original melt.
Such a physical separation of the solid crystals from the liquid melt is termed fractional
34
crystallisation. By this process, basaltic magmas can potentially give rise to basaltic andesites,
which in turn can produce andesite and still felsic, fractionated (or evolved) magmas such as
dacite, and eventually rhyolite.
There are three ways in which this physical separation can occur. Firstly, it can occur by the
magma being squeezed out of the space between the crystals (more likely if magma is rising in
small quantities) and is known as filter pressing; secondly by the settling of crystals if magma
becomes stored for a long period of time in a magma chamber (more likely if magma is rising as
a pluton-sized body) and is known as gravity settling (an example of fractional crystallisation by
gravity settling in action can be seen in the Palisades Sill, New York, and will be covered in more
detail later); and lastly by crystals being carried away and deposited by convection currents in
the magma chamber. It won‘t take a rocket scientist to notice that the diagram below is the
same as the partial melting diagram on page 33, apart from the arrows go in the opposite
direction. But that is the point. Minerals melt in the sequence shown and crystallise in the exact
reverse order.
3). Assimilation
Another way in which magma composition can change is if the magma has enough heat to melt
some of the country rock it passes through. The process of magmatic stoping adds more felsic
country rock material into the magma thereby changing its composition. This is the process of
assimilation. A rising body of mafic magma can assimilate rocks of intermediate and felsic
composition rather than the other way around because felsic rocks have a much lower melting
point than mafic rocks and therefore are easier to melt. Assimilation causes the silica content
of magma to increase.
4). Magma Mixing
35
As the name implies, magma mixing is a process that causes two (or more) magmas to be mingled
together, forming a new magma of a composition between that of the original composition. The
two magmas may come together when one is injected into the base of another magma chamber.
Magma mixing may involve magmas with very different mineralogies. For example, a basic magma
containing olivine might conceivably mix with an acidic magma containing quartz. Normally,
olivine is in chemical equilibrium with a basic magma but quartz is not, whereas the reverse is
true for an acidic magma. So, mixing these magmas together would produce a mixture in which
olivine and quartz co-existed; a combination that cannot be explained by fractional
crystallisation and is therefore a good indication of magma mixing. These minerals that are not
in equilibrium and shouldn't be there, such as quartz in a basic magma, are known as xenocrysts
(i.e. "foreign crystals").
5). Magmatic Underplating
This process can help explain the great volumes of granitic plutonic rocks that occur in many
areas of the world and which cannot be explained by fractional crystallisation. The continental
crust contains high amounts of silica which have a relatively low melting point and could be
induced to partially melt if a significant heat source were introduced. Currently, geologists
think magmatic underplating plays an important role as the extra heat source needed to
generate granitic magmas in the lower continental crust. Initially, some of the mafic magma
coming from the asthenosphere works its way to the base of the continental crust where it
starts to pool or underplates. Heat from this cooling and crystallising mafic magma is conducted
upward to create larger volumes of felsic magma by partially melting more of the continental
crust. This felsic magma, in turn, separates from its solid component and works its way upward
in diapirs to a higher level of the crust where it slowly solidifies to form a batholith.
6). Thickening of the Crust
Magma can also be formed at a collision zone without subduction. In the final stages of
continental collision the crust has thickened so much, due to thrust faulting and overfolding,
that rocks which were initially nearer the surface can become deeply buried. The deeper a rock
36
is below the surface the hotter it will be, and consequently the more likely it is to melt. So
when rocks are buried to a depth of over 20 to 25 km during continental collision, the
temperature is raised to 650-700ºC, causing the felsic continental crust to partially melt. The
granite batholiths of the Cairngorms in Scotland resulted from the melting at the base of the
continental crust, which had been thickened because of folding and the piling up of thrust
sheets during continental collision 400 mya (Caledonian orogeny). This is the second process
that can explain the great volumes of granitic plutonic rocks that occur (i.e. too much to form
from fractional crystallisation processes).
DETERMINING THE HISTORY OF AN IGNEOUS ROCK
Our understanding of the origins of igneous rocks has been helped greatly by three things:
laboratory work on silicate melts; detailed examination of igneous rock textures; and
observations of the structure and field relationships of igneous rock masses.
1). Silicate Melts in the Laboratory
Observations of the order of crystallisation from silicate melts have led to the formulation of
an ideal sequence in which minerals could be expected to crystallise. This sequence has become
known as Bowen's Reaction Series, named after the American geologist N L Bowen who in 1928
devised the diagram.
Although nowadays seen as an over-simplification of what happens, this diagram neatly
summarises much of our understanding of how magmas form. The minerals at the top of the
diagram (olivine & plagioclase feldspar) have the highest melting point, and so crystallise first
as a magma cools. Following the arrows down will progressively take you to lower temperatures
of crystallisation. By turning the arrows around the diagram can be used to work out the
37
sequence in which minerals will melt as a rock is heated. This diagram plays such a key part in
understanding igneous processes that it should be used as mental "coat hanger" for your
learning about igneous processes.
The reaction series diagram shown above consists of two lines of descent for silicate minerals.
One comprises the ferromagnesian minerals which form a discontinuous reaction series (the
left-hand side). This means that the early-formed minerals react with the remaining melt or
magma to form new minerals as the temperature drops. This produces an effect rather like
that seen in a Smartie, and is known as a reaction rim or corona structure. The other consists
of the plagioclase feldspars, which form a continuous reaction series (the right-hand side). This
means that the early-formed minerals react with the remaining melt to give different varieties
of the same mineral. This produces a mineral that looks, like a gobstopper sweet and is known as
a zoned crystal. The two series are related to temperatures of crystallisation and unite at
their lower temperature ends. The minerals at the high temperature ends of the series (the
olivines and Ca-rich plagioclase feldspars) have the highest melting points of their respective
series and would therefore be expected to crystallise first as a basic magma cools. They are
also the densest minerals of their series and therefore tend to sink within the melt. These
would then be followed by augite and then hornblende with the most acidic minerals such as
muscovite and quartz crystallising last.
SOLID SOLUTION PHASE DIAGRAMS
A solid solution phase diagram explains the behaviour of a chemical solid solution series, such as
the transition from high temperature calcium-rich plagioclase (anorthite) to low temperature
sodium-rich plagioclase (albite), or the transition from high temperature magnesium-rich to low
temperature iron-rich crystals in ferromagnesian minerals such as olivine.
The conventions for the phase diagram include the following:
 Temperature is plotted along the vertical axis and the complete variation in chemical
compositions for either olivine or plagioclase along the horizontal axis.
 The diagram is divided into three fields, all liquid, liquid & crystal, and all crystal. The liquidus
line separates the all liquid phase from the liquid & crystal phase. The solidus line separates
the liquid & crystal phase from the all crystal phase. (The solidus and liquidus lines are
experimental, and have been determined by melting and cooling many melts at different
compositions).
There are a number of assumptions made when using these solid solution phase diagrams.
Firstly, that the system (i.e. magma) remains in equilibrium throughout its history so that all
reactions can take place and everything can come to stability. What these means is, as
temperatures drop a mineral becomes unstable and reacts with the surrounding melt to change
into a new variation of the original mineral which is now stable at the new temperature.
Secondly, everything in the original melt remains in communication throughout the melting or
crystallisation process. What this means is that no magma or solid minerals are lost from the
magma chamber as it crystallises.
38
CRYSTALLISATION OF THE OLIVINES
The pale green mineral found in igneous rocks which we know as olivine is in fact one member of
a whole family of such minerals. The olivines are silicates of iron (Fe 2+) and/or magnesium
(Mg2+), e.g. MgFeSiO4. The ions of magnesium and iron are each of valency 2 and their
dimensions are comparable (the ionic radii are: Fe2+ = 74 pm; Mg2+ = 66 pm). This means that
Mg2+ and Fe2+ are easily interchangeable within the well-ordered lattice structure of the
mineral and so a continuous range of composition is possible, from an all-magnesium variety of
olivine (Mg2SiO4 - forsterite) to an all-iron variety (Fe2SiO4 - fayalite). This phenomenon of
interchangeability of ions is known as isomorphous substitution (isomorphous means "same
shape", i.e. there is no change in the crystal structure throughout the series). This series is
better known as a solid solution series. The diagram below shows the cooling curve for the
olivine family from forsterite to fayalite. The melting point of pure forsterite is 1890°C and
that of pure fayalite is 1205°C. This melting-point curve is quite difficult to follow so here an
example, and shown on the diagram below:
39
Take magma of composition 50% forsterite (Fo), 50% fayalite (Fa) at temperature of 1800°C,
shown as point
. The magma will cool to about 1650°C, point
before any crystals appear.
When they do, they will have the composition shown by drawing a line to the solidus curve
,
then dropping another line down to the bottom of the graph, i.e. point
; here a composition
of Fo 80% and Fa 20% can be read off. Such crystals will now be in equilibrium with the melt at
that temperature.
As the temperature drops, however, the composition of the melt changes along the upper line
and that of newly-formed crystals changes along the lower line, the in the direction of the
arrows. This has some interesting implications. Unless the early-formed crystals have been
removed from the melt, perhaps by sinking to the bottom of the magma chamber, they will find
themselves surrounded by a liquid with which they are no longer in equilibrium. Hence, fresh
mineral growth around the margins of such crystals will be more iron-rich than the original core
of the crystal and it will appear "zoned". In a rock which contains larger, early-formed crystals
(phenocrysts) set in a finer-grained groundmass; the large crystals may be markedly richer in
magnesium than the later, smaller ones. If cooling is slow enough, the whole crystal may be
altered by such reaction.
40
CRYSTALLISATION OF THE PLAGIOCLASE FELDSPARS
The plagioclase feldspars also form a continuous series of isomorphous substitution (solid
solution) between albite (NaAl3SiO8) and anorthite (CaAl2Si2O8). The properties that enable
ions to substitute for one another to permit the change from anorthite to albite is similar to
that of olivine, i.e. the ions of calcium and sodium are of comparable radius (Ca 2+ = 99 pm : Na+ =
97 pm), although of different valency. There is therefore enough space in the silicate lattice to
allow substitution of one element for the other. The difference in valency is relatively easily
compensated for by variation in the proportion of aluminium (Al3+).
a. Equilibrium Crystallisation of Magma
In the solid solution phase diagram below the high temperature anorthite (Ca-rich) and low
temperature albite (Na-rich) are plotted along the horizontal axis. (The fact that anorthite
crystallises at a much higher temperature than albite is just what we would expect from
Bowen‘s reaction series. The larger Ca (0.99Å) is higher/hotter than the smaller Na (0.97Å).
41
If we take liquid melt (magma) with a composition of 30% anorthite (Ca-rich plagioclase)
and 70% albite (Na-rich plagioclase) and
cool it from a temperature of 1500°C to the
liquidus line at about 1380°C the first crystals begin to form.
To determine the
composition of the first crystal move horizontally across the diagram to the solidus line (this
line always indicates the crystal composition),
then drop from the solidus straight down to
the bottom scale. The first crystal is 72% anorthite. (The diagram is always read in this
manner, down-across-down, regardless of the starting composition).
As crystallisation occurs Ca is removed from the melt faster than Na. So a 30% anorthite melt
yields a first crystal not at 30% but at 72% anorthite. As a result as crystallisation proceeds,
Na concentration in the liquid melt gets higher and the Ca gets lower. Thus, as the temperature
lowers the composition of the liquid melt migrates down the liquidus line. But at the same time
the composition of the crystals forming are moving down the solidus line. This is because the
magma system is cooling slowly enough to stay in equilibrium at all times. The earlier formed
crystals react with the melt, exchanging Ca for Na, to come to a composition in equilibrium with
the temperature at that moment. The composition of the melt and crystals move down in
tandem, exactly opposite each other along a horizontal line.
If we were to slowly drop the temperature from 1380°C to 1290°C the resulting crystals
would have a composition of about 56% anorthite and the remaining liquid melt would be 14%
anorthite in composition (i.e. Na-rich). If the system remained in equilibrium at all times so that
all reactions can take place and the original melt remains in contact throughout the
crystallisation process the composition of the final crystals would be exactly the same as the
initial melt, i.e. 30% anorthite and 70% albite.
b. Equilibrium Melting of Rock
This process can be reversed by beginning with a solid rock and heating it slowly until it melts.
In this case the diagram is read up-across-up and can be seen in the graph on the next page.
If we were to take an igneous rock with an original composition of 60% anorthite (Ca-rich
plagioclase) and
heat it to 1310°C (i.e. until it hit the solidus line) the first drop of liquid
would form as the rock started to melt.
To find out the composition of the first melt draw
a line straight across horizontally to the liquidus line,
then drop straight down from the
liquidus line to the composition scale at the bottom and read the composition. It is about 16%
anorthite (Ca-rich plagioclase), and therefore 84% albite (Na-rich plagioclase).
So the first drop of melt (magma) is always more Na-rich than the original crystal. As heating
continues more solid crystals will melt but at ever higher Ca compositions. This is because as
the temperature rises the higher Ca crystals are the ones becoming more unstable, and thus
melting. (Actually, entire crystals are not melting so much as the more Na-rich fractions of the
crystal melt first).
42
For example if we continue to heat the rock to say 1440°C (as on the graph above) the
magma composition will have changed to about 47% anorthite (Ca-rich plagioclase) and the
remaining solid crystals will now be over 80% Ca-rich.
So you can see that just as with crystallisation, during melting of a rock magma composition and
crystal composition will move in tandem horizontally across from each other, but this time up
the solidus and liquidus lines until all the crystal is melted. The composition of the final magma
will be the same as the composition of the original rock, in this case 60% anorthite and 40%
albite. At the point no crystal is left the line moves straight up the graph as the temperature
continues to rise.
c. Fractionation by Partial Melting
Melting or crystallisation under equilibrium conditions (where everything in the original magma
or rock remains in contact throughout the crystallisation or melting process) is not very helpful
43
or interesting, as you always end up with what you started with! In these situations magmas
cannot evolve into new magmas of different compositions. In real magmatic systems
crystallisation and melting do not often remain in equilibrium. As an example of how magma can
evolve into a new magma of different composition by chemical fractionation through the
process of partial melting, look at the phase diagram below.
Take an igneous solid rock of composition 50% anorthite and 50% albite (i.e. 50% Ca and
50% Na-rich plagioclase) and
slowly heat it until it begins to melt. The crystals do not melt
uniformly, rather the Na fraction (lower in the reaction series) begins to melt before the Ca
fraction (higher in the reaction series).
The composition of the first melt is found by drawing a vertical line up to the solidus line
(at about 1260°C), and then across to the liquidus line, and
then back down to find the
composition of the first melt. This is about 9% anorthite (i.e. 91% albite).
So far we have not done anything different. But now imagine that this first melt is removed
from the system so it can no longer react with the remaining crystals (this could be done by
44
filter pressing for example). We have now got a 9% anorthite (Na-rich) melt which, if allowed
to cool and crystallise, will form an igneous rock very different from the original rock. We could
of course separate the melt from the unmelted crystal at any time during the process and
fractionate the system. In each case, the melt fraction would be more Na-rich and the
remaining solid rock would be more Ca-rich. For example, if we had taken the temperature of
our 50% anorthite to 1350°C and then fractionated the system, we would have liquid fraction
with a composition of about 25% anorthite (75% albite) and a remaining solid fraction of
composition 65% anorthite (35% albite).
d. Fractionation by Fractional Crystallisation
Fractionation can also occur from fractional crystallisation. As with partial melting what is
required is separating the crystal or solid fraction from the melt or liquid fraction before
complete crystallisation has occurred.
If we take a magma composition of 40% anorthite and
cool it to 1400°C it will split
into two fractions (one liquid and one of solid crystal).
To read these compositions off the
phase diagram we need to start at the top and zip-zag our way down the graph, which will give
45
the
remaining liquid fraction a composition of 35% anorthite (65% albite) and the
new
solid fraction a composition of 73% anorthite (27% albite). If this solid fraction were to be
separated from the remaining magma by gravity settling for example the 35% anorthite magma
would cool and crystallise into an igneous rock of a different composition from the original
magma. This magma is said to have evolved, and will always form a rock of higher acidity than
the magma it initially evolved from.
2). Textures of Igneous Rocks
A considerable amount may be learned about the origin and history of an igneous rock from its
texture. The importance of grain size in classification and as an approximate guide to the depth
at which the rock crystallised, is well known to you. However, there exists a wide variety of
detailed textures, many of which have frightening Greek names and some of which are not
easily understood! The table below outlines the main characteristics of each texture:
TEXTURE
Glassy
texture
EXAMPLES
Obsidian
DESCRIPTION
Usually black and shiny in
hand specimen, with
conchoidal fracture.
INTERPRETATION
Resulting from supercooling of
magmas where cooling was too
rapid for crystallization and the
magma set as a glass. With time
the
glass
crystallises,
or
devitrifies into tiny crystals.
Coarse
grained
Crystals are easily seen
by the naked eye.
Crystals had plenty of time to
grow around a limited number of
nuclei. Usually typical of plutonic
rocks.
Medium
grained
Crystals give a speckled
appearance.
Crystals formed quickly around a
greater number of nuclei. Typical
of minor intrusions (dykes &
sills).
Fine grained
Crystals not
distinguishable with the
naked eye.
Crystals formed very quickly
around many nuclei. Typical of
lava flows and chilled margins of
minor intrusions.
Grain size
46
Euhedral
Well-formed crystals
showing perfect or nearperfect crystals.
Usually the first crystals to form
in a magma and therefore
unrestricted.
Subhedral
Crystals show an
imperfect but still
recognisable crystal
form.
Formed at a time intermediate
between the early and late
stages of crystallisation.
Anhedral
Crystals show no regular
crystal form.
Usually the last crystals to form,
filling up gaps. Many rocks
consist largely of equidimensional
anhedral or subhedral crystals.
Their texture is referred to as
equigranular.
Zoned
crystals
Common in plagioclase
feldspar. Crystals show
concentric "rings", the
central portion generally
being more Ca-rich than
the outer zones which
are relatively Na-rich.
Reaction has taken place between
crystal and liquid in a continuous
solid solution series under
relatively quick cooling. There is
not time for the whole of the
early-formed crystal to react
with the changed composition of
the melt.
Corona
structure
This appears as a type of
zoning, except that the
zones are composed of
different minerals, e.g.
olivine in the core, with a
rim of augite, followed by
an outer rim of
hornblende. Also known
as a reaction rim.
Also results from reaction
between early-formed crystals
and the melt as its composition
changes, so that the two are no
longer in equilibrium. Occurs in
minerals
forming
on
the
discontinuous
series,
i.e.
different
minerals
develop
rather than variations of the
same mineral.
Grain shape
Reaction
textures
47
Other
textures
48
Porphyritic
texture
Large crystals, usually
euhedral forms, are set
in finer-grained
groundmass. The large
crystals are termed
phenocrysts. Common in
lavas and in some
granites (e.g. Shap
Granite).
Phenocrysts
started
to
crystallise slowly at considerable
depth below ground. Then the
magma rose to a higher crustal
level, or was erupted as lava,
when the rest set quickly to
produce a finer-grained ground
mass. However, in some plutonic
rocks, it is possible that the
phenocrysts grew during the late
stages of the emplacement of
the magma, under the influence
of enriched volatile components.
Vesicles are created by the
exsolution of volatiles (gas) when
the confining pressure on a
magma is removed (i.e. when it
reaches the surface) but were
unable to escape before the lava
flow solidified.
After a lava has solidified
groundwater
can
percolate
through the vesicles where
evaporation occurs precipitating
any dissolved minerals such as
calcite to form amygdales.
Vesicular
texture
A lava with bubbles or
gas cavities known as
vesicles.
Amygdaloidal
texture
Vesicles
filled
with
secondary minerals such
as calcite or silica, known
as amygdales.
Flow banding
texture
Layering
in
lavas
especially thick, viscous
lava such as andesite and
rhyolite. Occur parallel
to the flowing lava
surface.
Layering due to flow movements
within the liquid rock before it
solidified.
Flow
brecciation
Angular
and
poorly
sorted, with a finergrained matrix. Matrix
harder than fragments
so
stand
out
on
weathering.
The breaking up a body of lava
into angular fragments by the
moving lava as it cools and
solidifies. Viscous magma flows
slowly, allowing the lava surface
to
solidify.
This
is
then
fragmented by the continued
movement of the still liquid
interior of the flow.
Eutaxitic
texture
Streaky texture
characteristic of a
pyroclastic flow deposits
made up of squashed
pumice fragments
surrounded by ash.
Corona Structure
Often the ash and pumice
fragments in a pyroclastic flow
are still hot when they come to
rest, and may become welded
together. Particularly likely near
base of the flow where the
process is aided by compaction
under the weight of the upper
part of the deposit. The pumice
fragments are squashed while
they are still hot and pliable into
forms described as fiamme
("fee-am-eh"), which is Italian
for flame.
Zoned Mineral
Ca-rich
plagioclase
feldspar
Olivine
Augite
Na-rich
plagioclase
feldspar
Hornblende
The corona structure is an example of incomplete reactions on the discontinuous arm of the
Bowen‘s reaction series. Whereas zoned minerals are examples of incomplete solid solution
series reactions on the continuous arm of the Bowen‘s reaction series.
3). Observations of the Structure and Field Relationships of Igneous Rock Masses.
Much can be learnt about the way magma behaves during crystallisation and about its origins by
studying major intrusions. Most of the major intrusions fall conveniently into two broad groups,
the basic and the granitic, using the terms in their broadest sense, and these will be considered
separately.
There are a number of very large basic intrusions which can be quite informative about the way
in which magma crystallises. In dealing with major basic intrusions it should not be forgotten
49
that the rate of cooling is completely different from extrusive events and some major
intrusions can take thousands or even millions of years to finally consolidate.
The significance of the laboratory work by N L Bowen and others has already been outlined. In
a sense, a large basic intrusion may be thought of as a gigantic natural laboratory where the
results of processes acting over a long period of time may be preserved. Observations of
reaction rims and the order of crystallisation from silicate melts have led to the formulation of
an ideal sequence in which minerals could be expected to crystallise. This sequence, which you
should be well aware of by now, has become known as ―Bowen‘s Reaction Series‖. It consists of
two lines of descent for silicate minerals which are related to temperatures of crystallisation
and unite at their lower temperature ends. The minerals at the high temperature ends of the
series (the olivines and calcium-rich plagioclase feldspars) have the highest melting points of
their respective series and would therefore be expected to crystallise first as a basic magma
cools. They are also the densest minerals of their series and therefore tend to sink within the
melt. These would be followed by the pyroxenes and then the amphiboles with the most acidic
minerals such as muscovite and quartz crystallising last.
It is obviously not possible to watch an igneous intrusion crystallising, so geologists use indirect
evidence to determine its cooling history. Much of this comes from studies of detailed features
such as the zoning of crystals which can be related to the experimental work on magmas
already described. There are also a few intrusions which show a systematic distribution of
minerals within them which can be related to Bowen‘s Reaction Series.
The Palisades Sill of New Jersey, USA, provides geologists with such an example. This basaltic
sill is over 300m thick which was intruded as magma horizontally into sandstone country rocks.
The variation in mineral composition from top to bottom of the sill makes it a classic example of
fractional crystallisation. After the intrusion was emplaced at a temperature near 1200 ºC the
magma in contact with the surrounding country rock was cooled rapidly to become a finegrained rock, effectively preserving the chemical composition of the original magmas. The still
molten interior cooled more slowly, so that larger crystals formed. The first mineral to
crystallise was olivine, which being a ferromagnesian mineral (Fe & Mg) is heavy and sank
through the molten magma to the base of the intrusion, where it can be found today in the
olivine-rich layer just above the chilled basalt zone. Some of these crystals may have taken as
long as 200-300 years to sink down through the intrusion. Minerals which sink and accumulate
at the bottom of intrusions are known as cumulate minerals. The sinking (known as gravity
settling) is one way in which an originally homogeneous magma may begin to crystallise in various
fractions, each of which may be significantly different from another. The second mineral to
form was augite (also a ferromagnesian mineral) and accumulated in the lower third of the
intrusion due to gravity settling.
With the crystallisation of calcium-rich plagioclase next this explains its concentration below
the sodium-rich plagioclase which crystallises at lower temperatures. Because of the early
settling out of the more basic (mafic) components of the intrusion the residual liquid becomes
more acidic (felsic) so that the last part of the intrusion to crystallise, which is about 40
metres from the top, consists of acidic lenses containing quartz and orthoclase feldspar. These
have crystallised from aqueous solutions which were concentrated in the last part of the
50
intrusion after the crystallisation of the anhydrous phase. The presence of these minor lenses
of acid (felsic) rock in the Palisades Sill is predictable from Bowen‘s reaction Series. They tend
to be coarse-grained, which probably reflects the concentration of water and other volatiles as
the liquid phase of the magma was gradually displaced upwards by the sinking crystals.
It is worth noting that recent studies have shown that this explanation may be something of an
oversimplification. In fact there are 16 metres of normal dolerite between the top of the
chilled margin and the olivine layer. This has led speculation that there may have been more
than one pulse of magma intruded, with the olivine settling out at the base of the second pulse.
The presence of minor lenses of acid rock in the Palisades Sill is predictable from Bowen‘s
Reaction Series but we need to examine an example of a larger intrusion to see if significant
quantities of acid rocks can be generated in this manner.
The Skaergaard Intrusion of east Greenland (see sketch diagram on next page) is one of a
series of Tertiary intrusions and is an excellent example of a so-called layered basic intrusion.
It has the shape of an inverted cone and, since it has been closely studied by many workers, it
is very well documented, although still not perfectly understood.
Around the edges of the intrusion there is a fine-grained olivine-gabbro series known as the
marginal border group, which is believed to represent the chilled margin of the intrusion, and
so, just as with the chilled margin of the Palisades Sill, its composition is thought to be the
same as that of the original magma. The marginal border group has suffered contamination due
to melting of the surrounding rocks, in particular where it is in contact with gneiss, which has a
lower melting point than the Tertiary basalt lavas. However, sufficient samples which are
considered to be uncontaminated have been obtained to state with some confidence that the
original magma was tholeiitic in character.
51
Sketch section through the
Skaergaard Intrusion
The main part of the intrusion is made up of the layered series. As the name suggests this is
composed of a succession of layers, each of which is turned up at the edges and the series
represents cooling from the base upwards. By no means is all of the layered series exposed as
can be seen from the diagram. Indeed, it has been estimated that over 70% of the intrusion is
not exposed and therefore some of the deductions and conclusions reached are, of necessity,
rather tenuous.
Within each of the layers, a clear stratification is seen with the dark, heavy minerals, pyroxene
and olivine, at the base and the less dense, lighter coloured plagioclase feldspar concentrated
near the top of each layer. This layering is thought to have an origin very much like that of
graded bedding in sedimentary rocks. As convection currents swept down through the intrusion
they transported crystals with them and deposited them as a ―mush‖ on the floor of the
intrusion. Gravitational settling of the heavy, dark minerals within these cumulate layers would
then be responsible for the light and dark bands. Within the layered series, troughs can be
seen which have been scoured out by the convection currents.
Less obvious, but of greater significance, is the trend of changing mineral compositions
throughout the layered series. Passing upwards, the minerals tend towards the low temperature
end members of their respective solid solution series. Thus the ferromagnesian minerals such
as olivine and pyroxene become enriched in iron and the plagioclase feldspar becomes richer in
sodium. The complete sequence of minerals shown in the Bowen‘s Reaction Series is not seen in
Skaergaard as amphiboles and biotite are absent. The intrusion exhibits differentiation to
ferrogabbros rather than granites although there are minor quantities of acid rocks in the
form of granophyre (type of microgranite).
The diagram over the page summarises the variations in chemical composition of the cumulus
minerals. Plagioclase feldspar shows the most complete sequence varying continuously from An69
(i.e. 69% anorthite, 31% albite) at the base of the layered series to An33 at the top.
Plagioclases of composition An77 have been found in the marginal border group, and it is
assumed that this represents the composition of the first plagioclases to crystallise. If this is
52
so, plagioclase feldspar of composition An77 should be found near the base of the hidden part
of the layered series. In fact, drilling has confirmed this prediction.
The sequence of chemical variations in the cumulate olivines is not complete since olivine is
absent in the middle zone of the layered series. At the base of the lower zone of the exposed
layered series the olivine has composition Fo67 (i.e. 67% forsterite, 33% fayalite), and it
reaches Fo53 by the top of the lower zone. When it reappears in the upper zone its composition
is Fo36 and by the top of the upper zone of the layered series it is Fo 2. When the top of the
layered series it is reached, the iron content of the rock is great enough for it to be called
ferrogabbro.
The third part of the intrusion is the upper border group which, because of erosion, is only
seen in part. This is thought to represent cooling from the top downwards since the reverse
trend in plagioclase composition compared to that in the layered series has been recorded.
The last part of the intrusion to crystallise was the Sandwich Horizon between the upper
border group and the layered series. This is the most acidic part of the intrusion and it
contains lenses of granophyre with some quartz and orthoclase feldspar, set in quartzferrogabbros.
53
Thus, just as with the Palisades Sill, rocks of acidic composition can be obtained from a basic
magma. A composition of chemical analyses of Skaergaard with those given for the Palisades
Sill shows similar trends although the final products of differentiation are more extreme in
Skaergaard. However, in both cases, the amounts of acid rocks produced are very minor. Very
slow cooling is necessary for this to happen and it has been calculated that over 12,000 years
were needed for the exposed parts of the layered series alone to crystallise. Because of their
slow cooling, each of these intrusions has preserved a record of the order of crystallisation of
the minerals and has led to a greater understanding of the processes of differentiation working
in large plutonic basic intrusions.
Granites and associated rocks such as granodiorite form the largest group of intrusive igneous
rocks, comprising over 90% of the total. This contrasts sharply with extrusive rocks where
basalt is the most common rock type. Part of the reason for this contrast between the
composition of the main intrusive and extrusive rocks has been explained in a previous section
on the crystallisation of magma, but part of the answer also lies in the origin of the respective
magmas.
Basic magma is produced mainly by partial melting of the ultrabasic mantle. Does the acidic
magma which forms granite also come from the mantle or does it have a separate source? It is
possible for small quantities of granitic rock to be formed by differentiation of the primary
basic magma as has been illustrated in both the Palisades Sill and the Skaergaard Intrusion.
However, in order to produce all the granitic rocks by differentiation, there would have to be
at least ten times as much basic magma as acidic magma. This is clearly not the case, as the
acidic intrusions are nearly twenty times as abundant as all other intrusions.
As basic magma rises through the continental crust, partial melting may occur and the products
become assimilated into the magma thus contaminating it. However, it is very unlikely that this
process would result in any significant quantities of granites.
If most granites are not formed from primary basic magma, how do they originate? A clue to
their origin is provided when the distribution of granites is studied. We find that the vast
majority are associated with the continental crust, usually within former orogenic belts. The
few small intrusions which are in oceanic areas can be accounted for by differentiation of a
basic magma. The chemical composition of the upper continental crust closely approximates to
that of granite, and the orogenic areas are regions where temperatures and pressures are
abnormally high. Theoretically, therefore, granites could have been formed by the melting of
the upper continental crust, but is this possible in practice?
Anhydrous (dry) granite begins to melt at about 950°C – 1,000°C at the pressures encountered
in the crust. As the geothermal gradient, at the time of orogenic activity, could have been as
high as 30°C per km, the temperatures required to produce melting are found at a depth of less
than 35 km, which is well within the continental crust in orogenic areas. Hydrous (wet) granitic
melt can be produced at even shallower levels in the crust as the temperature required is only
about 650°C, which is reached at a depth of 20-25km.
Although we can see that it is perfectly possible for granites to be produced by the heating up
54
and eventual melting of continental crust, the idea has given rise to a long debate between
geologists. On one hand there are the ―transformationists‖ who saw that granite could indeed
be produced by transformation of continental crust. To the ―transformationists‖, granite is the
end product of metamorphism, produced first as a ―mixed‖ rock, part gneiss, part granite,
known as migmatite (this term actually means ―mixed-rock‖) which itself undergoes complete
transformation at a later stage, becoming granite. On the other hand there are the
―magmatists‖ who point to the many examples of granites that clearly have formed from an
intrusion of magma, and are thus truly igneous. There are points to be made on both sides of
the argument and the following examples illustrate features which suggest a magmatic origin in
some cases and an igneous, or magmatic, origin in others. There then remains the question of
whether the genesis of these two types of granite can be connected in some way.
As has already been pointed out, it is very common to find granites in orogenic belts associated
with areas of high grade regional metamorphism. The sequence of metamorphic rocks – slate –
phyllite – schist – gneiss produced from shales or mudstones with increasing grade of
metamorphism is well known. Not quite so well known, perhaps, are the properties of the group
of rocks known as migmatites.
These migmatites are most often associated with schists and gneisses and they are in fact
―mixed‖ rocks where granitic material has come to be mingled with host rock which originated
by high grade regional metamorphism. The granitic material seems to have migrated through
the host, either as magma or aqueous fluids of granitic composition, although it is possible that
it took the form of a diffusion of ions migrating through pore fluids. These are rather loosely
referred to as ―emanations‖. It is often claimed that these emanations have their origin outside
the host, perhaps rising from deeper levels in the crust, although other authors believe that in
some cases the granitic material may have originated as a segregation from the host itself.
It is quite easy to see that there are two distinct components in hand specimens, even though
metasomatic reactions between them will have caused some modifications in composition. The
original composition of the host will have an important part to play in the composition of the
final product, but with increasing migmatisation a rock is produced which approaches granite in
composition.
Therefore, it has been quite convincingly argued that migmatites provide a link between the
true metamorphic rocks and granites and that the complete metamorphic series should read:
slate – phyllite – schist – gneiss – migmatites – granite. In Britain the main areas where
migmatites are found are in the Highlands of Scotland and in Ireland. Here, during the
Caledonian Orogenesis, a broad zone of regional metamorphism was formed and migmatites can
be seen in such places as the Central Sutherland Complex in Scotland. Here, as in many other
examples there is gradation from an area where the host rock, in this case Moine Schists, is
traversed by a few veins of granite, though zones where the granitic component becomes more
conspicuous to local occurrences of almost pure granite conforming to the strike of the country
rock.
In contrast, granites which are clearly of magmatic origin tend to be smaller, more
homogeneous bodies than the granites that are believed to have been generated ―in situ‖ by the
55
processes of migmatisation.
Although not strictly a granite, because of its lower silica content, the main part of the Criffell
granodiorite of the Southern Uplands of Scotland provides an excellent example of an
intrusion. It is one of a series of Caledonian granite outcrops in the area and it has been
forcefully intruded into folded Silurian shales and greywackes. There are three parts to this
intrusion, for the main granodiorite grades into a porphyritic granodiorite with feldspar
phenocrysts in its central portion. The third part is older than the main complex and is seen at
its south-western end. There, finer grained granodiorites are associated with some quartzdiorite rocks.
The Criffell granodiorite shows a large number of features which could only have originated by
the intrusion of hot magma into colder country rocks. Most significantly, in contrast to the
magmatic granites previously described the contacts with the country rocks are sharp and they
dip steeply outwards. This has been confirmed by geophysical surveys across the intrusion.
There is also no gradual transition from high grade regional metamorphic rocks or migmatites,
since the country rocks have only previously been metamorphosed to slate grade. The intrusion
has, however, formed a clear metamorphic aureole within the country rocks which have been
hornfelsed and in some places mineralised. In the outer parts of the intrusion xenoliths are
common and veins of the granodiorite invade the country rock in places. A flow foliation has
been recorded in the outer parts of the intrusion which indicates that the origin must have
been from a mobile magma. Although the outcrop of the granite is slightly elongated in the
direction of the strike of the country rocks it is completely discordant to it at its western end
which again contrasts with a migmatitic granite.
All these features make it clear that this particular intrusion originated from a hot, mobile
magma. It is thought that the mass intruded to a relatively high level in the Earth‘s crust,
perhaps rising to within 1,000m of the Earth‘s surface before finally solidifying. The examples
illustrate that some granites result from crustal melting – the so-called ―metamorphic‖ or
magmatic granites – whilst others have a definite magmatic source – the so-called ―igneous‖
granites. How can these two greatly contrasting origins be resolved, if indeed they can be at
all?
In 1957, a ―Granite Series‖ was proposed which provided a link between the two types of
granite. In this series, the metamorphic granites represent deeper levels in the Earth‘s crust
where partial melting has taken place, a process known as anatexis.
The melt has then consolidated ―in situ‖. Sometimes, this melt becomes mobile and, because it
has a relatively low density compared to the rocks around it, it will then move upwards to
intrude into the higher levels of the crust where it will crystallise to form a magmatic granite.
The diagram on the previous page summarises the various stages and links between them in the
Granite Series.
One question needs consideration. If most granites can be referred back to crustal melting,
how is it that they show little variation in composition when the crust is very variable?
56
Experiments on mixtures of quartz, orthoclase and plagioclase have shown that they begin to
melt at varying temperatures, depending upon the relative proportions of the minerals in the
mixture. The mixture which remains liquid to the lowest temperatures contains approximately
equal proportions of the three components. Analyses of a great number of natural granites,
which were poor in ferromagnesian minerals, showed that they contained quartz, orthoclase and
plagioclase in much the same proportions as in the artificial mixture. This is because the natural
granites will contain the most easily melted fraction, i.e. melted at the lowest temperatures, of
the original crustal rocks and so variations in the composition of the crustal rocks will be
evened out.
Granites are also found outside orogenic areas where the thinner crust and the lower
geothermal gradient means that it is not possible for them to have been produced by crustal
melting. For these granites we have to revert to postulating an origin in the mantle. The
peridotite mantle must have suffered extreme differentiation to produce a granitic melt. This
theoretically possible, although very large quantities of material would be required, the ratio of
original peridotite to granitic melt produced being approximately 100 to 1. Evidence for this
hypothesis has recently been found. In crustal granites, and those thought to originate from
the mantle, the ratios of two isotopes of strontium have been found to be different, reflecting
their different origins. This new work on the strontium ratios has led to increased speculation
about the origin of granites. The migmatitic granites are obviously formed in the crust as they
are derived from partial melting of crustal rocks. However, magmatic granites could either be
derived from partially melted crustal material which becomes mobile and moves upwards as
outlined on the Granite Series above or they could alternatively have been derived from the
mantle.
57
SEDIMENTARY PETROLOGY
Sedimentary rocks are formed at the Earth‘s surface by deposition. They may be formed from
the fragments derived from pre-existing rocks, by the chemical precipitation of dissolved
substances or from the remains of plants and animals.
SEDIMENTS TO ROCK
Sediment is unconsolidated material deposited by water, ice, or wind. A particle in general
language means a piece. In geological language it is synonymous with the term grain. A clast is a
particle broken off a pre-existing rock, and is usually applied to a particle within a rock, rather
than a particle within sediment. The term clastic comes from the word clast, and means a group
of sedimentary rocks composed of particles which have been broken off pre-existing rocks.
Sandstone is an example of a clastic sedimentary rock. A fragment is a large clast
Diagenesis is the group of processes which change sediment into a sedimentary rock after
deposition has occurred, because of this they are referred to as post-depositional processes.
Diagenesis takes place at relatively low temperatures (just above 150-200ºC) and pressures at
or near the Earth‘s surface.
Lithification is a diagenetic process in which loose, unconsolidated sediments is converted into
sedimentary rocks. The turning of sediment into rock involves two main processes, compaction
and cementation. These processes begin as soon as sediments are deposited and their burial
commences.
1. Compaction
Compaction is an important first step in lithification. In a texturally mature sand deposit 30%
or more of its volume occurs as intergranular pore spaces. These pore spaces are filled with
water if the sediment was deposited in water or exists below the water table. If the sediment
is grain-supported immediately after deposition, a grain is in contact with one of its neighbours
at one point only. During compaction, the pressure of the overlying sediments packs the grains
closer together and more efficiently, reducing the volume of pore space and squeezing out the
pore water. As quartz grains are progressively buried, the pressure at the grain contacts
increases until the quartz begins to melt slightly and dissolve. This pressure dissolution is more
common in texturally mature sandstones than in those containing a significant amount of finegrained matrix. The matrix helps to spread the load thereby reducing the pressure at the grain
contacts. Pressure dissolution produces fused quartz grain contacts. A grain may simply become
indented where the adjacent grain has pushed into it. Or if dissolution is more advanced then
the contact may have a wavy, sutured appearance showing where one grain has become
intergrown with the next on dissolution.
When clay-rich sediments are first deposited, they may contain as much as 70 - 90% water by
volume. On burial, both water loss and compaction are relatively rapid and much of the pore
water is lost by the time sediments have been buried to about 1 km depth. During compaction,
58
any elongated or flaky grains such as clay or mica will become aligned parallel to the bedding
plain. This alignment of clay minerals may lead to mudstone and shales splitting easily into
layers and being known as fissile. No melting occurs to fuse the clay grains together because
they have flat surfaces and no sharp edges to concentrate pressure and cause melting. Clay
minerals however are naturally cohesive and stick together.
2. Cementation
Cementation is the second stage of lithification, and involves the gluing together of compacted
grains to form a rock. Often compaction alone will not produce a lithified rock. Sediments can
be cemented with a matrix or a cement. A matrix is the clay material that holds the other
grains together and comes from the mix of material deposited at the time. Cements are
precipitated from solution within sediments. Common cementing materials are silica (as quartz),
calcite, and more rarely iron oxide. Which of these occurs depends very much on the
circumstances under which lithification occurs. Iron oxide cement is indicative of deposition
under oxidising conditions such as those of deserts, alluvial fans and flood plains.
Diagenesis often results in the formation of new minerals which grow in the sediment or
sedimentary rock e.g. the green mica glauconite may grow in shallow marine sediments.
Greensand is a sandstone with a high proportion of glauconite. Limestones may show numerous
changes in mineralogy since aragonite changes to calcite and calcite may be altered to dolomite.
Other mineral changes which take place in sedimentary rocks include the alteration of feldspar
to clay minerals or to white mica.
Diagenetic processes are important for several reasons. They can considerably modify a
sediment, both in terms of its composition and texture, and in rarer cases, original structures
are completely destroyed. Diagenetic events also affect a sediment‘s porosity and permeability,
properties which control a sediment‘s potential as a reservoir for oil, gas or water.
POROSITY AND PERMEABILITY
Porosity is the measure of the amount of water that can be held in the pore spaces of a rock.
The porosity of a rock or sediment can be calculated using the following formula:
Porosity = volume of pore space
volume of sample
x100%
The porosities of some common materials are given in the tables below.
Material
Limestone
Sandstone
Shale
Clay
Approximate
Porosity (%)
10
18
18
45
Material
Sand
Gravel
Granite
Basalt
Approximate
Porosity (%)
35
25
1
1
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Porosity does not depend on grain size, but it does depend on factors such as grain shape, how
the grains are packed together, the amount of cement, the degree of compaction, the amount
of smaller grains between larger grains (grain sorting) and particle fall velocity.
Grain shape is an important factor affecting porosity. The diagrams below show the
differences in porosity produced by equidimensional spheres, cubes and plates. Very high
porosities may be produced in the spheres, and very low porosities in cubes and plates.
The arrangement, or packing, of these grain shapes also has an impact on porosity. The effect
of packing on porosity is shown in the following two diagrams. In diagram (a) the grains are all
spheres packed on top of the other. In a material like this, known as cubic packing, the porosity
is 48%. Where spherical grains are packed more closely with each grain sitting in the
depression formed the two grains underneath, as in diagram (b), the porosity is 26%. This form
of packing is known as rhombohedral packing.
(b) Rhombohedral packing
(a) Cubic packing
Porosity also depends on the amount of cement. Cementation is the principal process of porosity
loss in rocks such as sandstones. Silica, calcite and clay can all be precipitated as cements,
lining and filling pores and decreasing both porosity and permeability.
Cement
Pore spaces
Loss of porosity
Compaction during diagenesis also reduces porosity. Compaction takes place from a few metres
below the sediment surface, and results in closer packing of the grains, and eventually at
depths of hundreds to thousands of metres to pressure dissolution and interpenetration of
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grains. Rock fragments are particularly susceptible to deformation and crushing by increasing
overburden.
The porosity of natural sands with similar packing is independent of grain size, but it does vary
according to sorting. Experiments reveal a 25% difference in porosity between well-sorted
sands and very poorly sorted sands of the same mean grain size. The diagram on the next page
shows how porosity is reduced by the presence of small grains between the larger grains.
Well- sorted
Poorly sorted
Loss of porosity
When sediment is deposited the particles fall through the medium they are being transported
by, such as water, and sink until they reach the floor of the river or ocean. The speed at which
these particles fall has an affect on porosity. Increasing particle fall velocity tends to
decrease porosity and encourage close packing of the grains, since particle-bed impacts
transfer kinetic energy to underlying layers, causing particle shear jostling and rearrangements
that reduce pore space. This vibration-induced resettling is an important way to make grains
pack together more efficiently.
Permeability is a measure of the ease with which a fluid can pass through a rock or sediment.
Permeable rocks allow liquids and air to pass through fairly easily, while impermeable rocks do
not allow fluids to pass through. Materials with the largest pore spaces have the greatest
permeabilities, so rocks of equal porosity do not necessarily have equal permeabilities, e.g. shale
and sandstone may both have porosities of about 20%, but sandstone is much more permeable
because its larger pore spaces offer less resistance to fluid movement. Gravel is very
permeable because it has very large pore spaces.
The formation of cavities and fractures (joints and faults) in rocks may render them more
porous and more permeable. Limestones may become very permeable because water enlarges
cavities by solution as it passes through. Also igneous rocks may be made permeable by the
presence of joints.
WEATHERING PROCESSES
Weathering is a set of processes (physical, chemical and biological) that change the physical
and chemical character of the rocks at the Earth's surface, due to water, air and temperature.
Weathering helps break down a solid rock into loose particles that can easily be picked up or
physically removed by streams and eventually form new sedimentary rocks.
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1. Physical Weathering
This leads to rock breakdown by purely mechanical means (i.e. the rock disintegrates) as a
consequence of temperature changes or other causes of stress within rock bodies.
Frost shattering (freeze/thaw) is one example of physical weathering. Water collects in cracks
and joints within rocks, and when it freezes it expands by 9.2% as changes from a liquid to a
solid (primarily because ice takes up more room than water in its liquid state). The pressure
attained by water freezing in rock crevices can reach 14 MPa (remember 0.1 MPa is
atmospheric pressure), whereas the tensile strength of granite (the pressure at which it will
yield by fracturing) is only 4 MPa.
In mountainous areas, frost shattering can create spectacular landscapes, spiked with jagged
peaks and littered with rock fragments. These rock fragments are usually pebbled-sized or
larger, and angular to sub-angular in shape. The typical landforms they produce are screes.
These piles of unconsolidated rock fragments have surface slopes of between 34° and 40°. The
sorting of debris in screes is very common. The largest boulders are more frequently found at
the base of the scree slope, because of their greater momentum and the action of smaller
particles as ―bearings‖ over which the boulders may slide.
It takes many cycles of freezing and thawing to shatter rocks completely, however the process
is quickened by well-cleaved and jointed rocks as these provide convenient pathways for
percolating water. Freeze/thaw processes are most likely to be active in climatic zones where
temperatures fluctuate across the freezing point of water for significant periods of the year,
and when rocks are well cleaved or jointed. This means that both climate and the state of the
rocks themselves determine the effectiveness of frost shattering.
Exfoliation is second process of physical weathering. Rocks are poor conductors of heat, and so
if they are heated by intense sunlight they warm up and expand on their outer edges but not on
their insides. These different rates of expansion between the outside and inside of a rock
cause stresses to be set up, causing the rock to fracture. This is thought to be especially
important in desert environments with their hot days and cold nights causing the rocks to
expand slightly during the day and contract at night, again fracturing the rocks. If a rock has
mafic (dark) and felsic (light) minerals within it these will also heat up at different rates
causing different areas on the rock surface to expand at different rates. This further
stresses the rock causing it to fracture.
Salt weathering is a form of physical weathering of rock. No chemical alteration of rock
constituents is involved in salt weathering. The salt derives from an external source (capillary
rising ground water, aeolian origin, sea water along rock coasts, atmospheric pollution). Salt
weathering is favoured by dry conditions, such as are found in warm and cold (arctic) arid
climates. Salt weathering also occurs on buildings and monuments in arid climates as well as
under dry microclimatic conditions in humid climates. Rock breakdown by salt weathering takes
place through salt crystal growth in rock pores and/or through hydration (absorption of water)
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of hydrate forming salts. The expanding salt crystals exert a pressure on the walls of the rock
pores that exceeds the tensile strength of the rock.
2. Chemical Weathering
Physical weathering opens the way for chemical weathering to proceed more easily. This is
because it increases the surface area of the rock. Minerals making up igneous rocks crystallise
at high temperatures and pressures. They are stable when they form but become unstable
during exposure to the lower temperatures and pressures found at the Earth's surface.
Rock-forming minerals weather at rates that are proportional to their positions in Bowen's
reaction series. For example, high temperature minerals like olivine or augite form under
conditions very different from surface conditions. When exposed to the surface they become
unstable and weather faster than minerals that are formed at much lower temperatures and
pressures such as quartz.
Oxidation is an important process of chemical weathering because many of the rock forming
minerals contain iron in its reduced state as Fe2+, these are the ferromagnesian minerals
(olivine, hornblende, augite and biotite mica). When Fe2+, in olivine for example, encounters
atmospheric oxygen dissolved in water it will be oxidised rapidly to its more oxidised state,
Fe3+, to form ferric oxide. Ferric oxide, more commonly known as rust, is highly insoluble and
forms a brown or reddish-coloured material, coating rock surfaces.
Carbonation is the process whereby rainwater mixes with atmospheric carbon dioxide to make
the rain into a weak carbonic acid. This acid reacts chemically with alkaline materials such as
calcium carbonate (limestones & chalks). The calcium carbonate is dissolved and runs away in
solution into streams and rivers.
Hydrolysis is a more complex chemical reaction between the hydrogen elements in water and
minerals such as feldspar, mica and the ferromagnesian minerals. The hydrogen in the water
chemically reacts with minerals by replacing the metallic ions in a mineral such as K (potassium)
in feldspars. With the potassium replaced by hydrogen a new substance is formed, this is a
white powdery clay material known as kaolin. The potassium and some silica is dissolved and
washed away in solution.
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PRODUCTS OF WEATHERING
The products of weathering consist of rock fragments produced by physical weathering,
residual quartz grains that have been released by chemical weathering (although resistant
themselves), and other products of chemical weathering including new minerals (clay) and
various elements in solution (Ca, K, Na). Crystals of mica and feldspar may also be released as
residual products if chemical weathering is not very advanced. The table below shows the end
products of the chemical weathering of the main rock-forming minerals.
As an example the gabbro (a basic igneous rock) contains a little olivine, but mainly augite and
Ca-rich plagioclase feldspar. The olivine and augite in the gabbro would chemically weather to
produce Fe3+ (insoluble ferric oxide) as a new material and silica in solution. Clay minerals
(kaolin) would also form from the decomposition of the plagioclase feldspar. However, an
orthoquartzite (a medium-grained sedimentary rock) composed entirely of quartz, will not
decompose chemically and so would form an insoluble residue.
Once rock fragments and mineral grains have been released by a combination of physical and
chemical weathering, they begin their journey to their eventual site of deposition to form
sedimentary rocks (more of this later).
RATES OF WEATHERING
The rate a rock weathers at very much depends on its mineralogy. Many rocks are formed under
high temperatures and pressures and the minerals which they contain are stable under the
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conditions of rock formation. Under new conditions present at the Earth‘s surface these
minerals are unstable so they tend to react to form stable weathering products. In general,
silicate minerals with high proportions of silicon are more stable than minerals rich in metal
ions, e.g. olivine (Mg2SiO4) is less stable than pyroxene (MgSiO3) and anorthite (CaAl2SiO8) is
less stable than albite (NaAlSi3O8). This means that minerals formed at high temperatures are
less resistant to weathering than minerals formed at low temperatures. Minerals can be
arranged in order of their resistance to chemical weathering as follows:
Olivine
Anorthite
Augite
Hornblende
Albite
Biotite
Orthoclase
Muscovite
Quartz
Increasing
stability
How does this sequence compare with the crystallization sequence of minerals in igneous rocks
(i.e. Bowen‘s reaction series)? Carbonate minerals, such as calcite, have very little resistance to
chemical weathering because they are readily attacked by carbonic acid.
The surface area available for weathering to take place will also affect the rate of weathering.
In general, igneous and metamorphic rocks are more resistant to weathering than sedimentary
rocks because crystalline rocks are not porous so water cannot penetrate them. Well-bedded
and fractured rocks tend to weather faster than more massive, unfractured rocks for the
same reasons. In addition, coarse-grained rocks are generally less resistant than fine-grained
rocks. Where igneous rocks such as basalt and dolerite are cracked or jointed into blocks, the
blocks are often attacked by chemical weathering in such a way that shells of decomposed rock
flake off in a chemically caused form of exfoliation known as spheroidal or onion skin
weathering. Rocks can be roughly arranged in order of resistance to weathering as follows:
Metaquartzite
Granite
Basalt
Sandstone
Siltstone
Limestone
Increasing
resistance
The last factor affecting the rate of weathering is climate. Usually, all types of weathering
operate together to break down rocks. Climatic differences, though, dictate the dominant
weathering processes. Since the rates of chemical reaction double for every increase of 10°C
and since chemical weathering requires water, it follows that humid, tropical areas will show the
effects of chemical weathering most strongly. In temperate regions the end-product of
feldspar decomposition is clay but in the wet tropics clays are broken down further. Their silica
is carried away in solution, by the process of leaching, to leave insoluble, hydrated aluminium
65
oxides along with hydrated iron oxides to form a soil called laterite.
In tropical deserts physical processes dominate, such as exfoliation. However, any water which
is present tends to be drawn to the rock surface. Here, it evaporates and leaves its dissolved
salts whose growing crystals cause the rock to disintegrate.
In temperate areas physical and chemical weathering operate together so that, in general,
neither is dominant. Precipitation is, however, very variable from place to place so in dry areas
physical processes may dominate while in wet areas chemical processes may be the major
factors in rock breakdown. In temperate regions, chemical processes reach only a few metres
below the surface, whereas in the humid tropics such weathering may penetrate over 100m. In
winter, frost shattering is the dominant process.
Under Arctic conditions and on cold mountain tops the low temperatures mean that the
dominant weathering processes are physical. Frost shattering is the most important process.
MATURITY OF SEDIMENTARY ROCKS
The mineralogical composition of a sedimentary rock can be used to infer something about the
extent to which the original sediments have been chemically weathered. For example, suppose
you are examining a sandstone, containing 70% quartz & 30% fresh K-feldspar grains, both
derived from the same source. What might you infer about the chemical weathering
experienced by the source rock, and the resulting sediment during transport?
Well, K-feldspar (orthoclase) is a lot less resistant to chemical weathering than quartz, but is
found in pristine, unweathered state, not much chemical weathering can have taken place or
much transportation. Although, weathering must have been sufficient to decompose any other
minerals present in source rock.
Compositional maturity is a concept used when describing sediments composed of silicate
minerals. The degree of compositional maturity is a description of the extent to which sediment
has suffered chemical weathering.

Compositionally mature sediments contain almost nothing except quartz or clay
minerals (the end products of chemical weathering).

Compositionally immature sediments contain a good deal of undecomposed rock
fragments, fresh feldspar & other silicate minerals such as ferromagnesian minerals
we would expect to decompose rapidly.
A medium-grained, fragmental rock containing 75% iron-stained quartz & 25% feldspar held
together by a clay matrix would be described as compositionally moderately immature, and
would be classed as ARKOSE. A fine-grained, fragmental rock containing only soft clay minerals
would be described as compositionally very mature, and would be classed as a MUDSTONE or
SHALE.
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Textural maturity describes the degree of roundness, the size and the sorting of the grains
within the sediment or rock. The degree of textural maturity indicates how long the sediment
has been transported and whether this was by wind, water or ice.

Texturally mature sediments have grains that are well rounded, well sorted and
small in size, inferring they have been transported a long way, allowing attrition and
other erosive processes to get to work.

Texturally immature sediments have angular, large and poorly sorted grains,
inferring a short transportational history.
TRANSPORTATION OF SEDIMENTS
If you examine sand grains from various sources under a microscope you will find that the
grains differ in shape, composition, surface appearance and size. During transport and
reworking of the sediment the grains bump into each other and into the rock over which they
are travelling. The result of this is that the corners of the grains are constantly being chipped
away so the grains tend to become smaller and rounder. In describing grain shapes use is made
of terms such as rounded and angular. (Note that rounded does not necessarily mean spherical
e.g. an egg-shaped grain is well rounded). The chart on the next page shows how sediments can
be described in terms of their shape.
Transport by ice, wind or water has different effects on the grains. When sand is carried by
glaciers the grains are often splintered into shiny, angular fragments (glacial sand is sometimes
called ―sharp sand‖). Wind-blown sand grains collide violently because air does not cushion the
impacts. The grains often become almost spherical in shape but frequent impacts leave the
grains chipped and frosted. Grains transported by water tend to develop polished, glassylooking surfaces. They are generally less well-rounded than wind-blown grains. However, water
transported beach sands will be more rounded than water transported river sands. Immature
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sedimentary rocks such as greywacke and arkose often consist of angular grains because they
have not been transported very far or for very long by water.
The shape of sedimentary particles is partly controlled by its size. Large particles such as
pebbles and boulders quickly become rounded because their collisions are very violent. Sandsized particles require much more transport to become rounded because their collisions are
less severe. Very small particles never become rounded because their collisions are effectively
cushioned by the internal structure of the mineral or rock being transported, e.g. quartz lacks
cleavage so it tends to form round grains. Mica has one very good cleavage so it forms flaky
particles. Igneous rocks readily form round pebbles and boulders while rocks which have welldefined layering or preferred grain orientation tend to form flat or elongate pebbles and
boulders.
Loose sediments are transported by fluids such as water and wind. Obviously the sediment
cannot be moved unless the fluid moves. Two ways fluids can flow: Laminar flow (individual
parcels of fluid move in straight lines parallel to each other). Turbulent flow (parcels of fluid
move in a more chaotic fashion, up, down & in eddies) occurs as fluid increases speed & viscosity
decreases.
The factors affecting the transport of sediment in a fluid include fluid speed, grain size, grain
density, grain shape and turbulence. The more turbulent, the faster the flow, the smaller the
grains, the smaller the density difference between particle & water and flatter the grains the
more sediment will be moved.
1. Water has a higher viscosity (50 times greater) than air and a much higher density (1000
times denser) than air, which means it can move larger particles/sediment than air. Water flow
becomes turbulent once speeds of over a few cm/second occur and this also causes bigger
particles to be picked up in water than air.
Bedload (larger particles continuously or intermittently in contact with the bed during
transport) are carried by traction (rolling & sliding) or saltation (bouncing). Suspended load
(smaller particles such as clay particles & very small quartz grains less than 0.1 -0.2 mm) caught
up in the turbulent eddies carried in suspension.
Sorting
These various methods of transportation move particles at different rates. Rolling & sliding is
slower than saltation, which is slower than sediment carried along in suspension. The
consequence of this is that finer-grained material will be moved faster & further than coarsergrained material, so it will be separated out or sorted.
The coarser the grain size the greater the current speed required to start it moving, firstly as
bedload and only lifted into suspension when the current speed reaches a certain threshold.
However this is complicated slightly by finer-grained silts & clays. Fine-grained sediments,
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especially clays which are flaky, are very cohesive (tendency to stick together); this means a
higher current speed is required to start them moving.
Shape
Water-transported grains are usually rounded to well-rounded and normally have a glassy
appearance, like the crystals of quartz in granite.
2. Wind - Air has a low viscosity and low density which means it cannot move as large particles
as water. Air is less viscous than water so all air flow is turbulent which enables it to move some
smaller particles. The wind cannot transport as big particles as water flowing at the same rate
due to air's reduced density & viscosity. Wind needs higher speed to move particles of a larger
size.
Bedload transport by saltation is the dominant process (pebbles too big to slide or roll). Air
offers much less frictional resistance to saltating grains due to lower viscosity so they can leap
higher than they would if cushioned by water. This forms a moving cloud up to as much as 1m
above the surface. Higher saltating grains of sand hit the grains of sand beneath them at much
higher speeds when they descend. The hit grains are pushed forward as they are struck, not
enough to throw them into the air but sufficient to cause a slow creep forward. An impacting
sand grain can move another grain up to 6 times its own diameter through process of surface
creep. Suspension (only finer-grained silts & clays less than 0.1mm lifted & held permanently in
suspension)
Sorting
These various methods of transportation move particles at different rates. Surface creep is
slower than saltation, which is slower than sediment carried along in suspension. The
consequence of this is that finer-grained material will be moved faster & further than coarsergrained material, so it will be separated out or well sorted.
69
Silt-sized particles are lifted directly into suspension by the wind. Fine- to medium-grained
sand will begin moving by saltation, and then be lifted into suspension as the wind speed
increases (although it would take a gale force wind of 20m/s or 72 km/h to lift sand grains 1mm
in diameter into suspension). Coarse-grained sand begins moving by surface creep & gravel
coarser than 4mm does not move at all; it is left behind as a lag deposit. So as wind moves
across loose sediment finer-grained material is removed leaving coarse-grained material behind:
a process known as winnowing. This selectivity
and the different rates at which transport by
surface creep, saltation & suspension occur
causes the wind is so effective at sorting
sediments (much more so than water).
Shape
Aggressively bouncing grains in wind flow means
every time they strike other particles or rock
surfaces attrition occurs. Angular corners &
sharp edges worn down so become progressively
70
more rounded. As air has a lower viscosity the grains are not cushioned as they are in water.
The impacts with each other & the bed beneath are much harder & so rounding occurs very
much more rapidly in air. Also it takes a higher wind speed than water current speed to set
grains of a given size in motion so the speed of impact is correspondingly greater leading to
more rapid rounding. Wind-blown sands are rounded to well-rounded.
Their degree of roundness is a textural property which becomes a measure of the time they
spent in transport (longer grains are bounced the more rounded they become. Aeolian quartz
grains have a matt finish, similar in appearance to the frosting used to decorate some glass
bottles due to surface pitting caused by high speed collisions among grains in air.
GRAIN SIZE ANALYSIS
While the sizes of pebbles and boulders have to be individually measured, you can find out a
great deal about sand samples by passing them through stacks of sieves whose mesh sizes show
a regular decrease from top to bottom. The mesh sizes decrease by half from sieve to sieve so
the complete stack would normally have mesh sizes 4, 2, 1, 0.5, 0.125 and 0.063 mm. and
beneath the finest sieve a pan catches the particles which come through.
You can study sand from places such as beaches, coastal dunes, rivers and sand pits. Some
sandstone can be separated by careful grinding or by the use of acid. Before sieving, the sand
sample is weighed out to about 100g, and the sieves are cleaned and stacked in the correct
order (largest mesh size at top). The sand sample is tipped in and the sieves are shaken till
each fraction reaches a constant weight – probably about 5 minutes. After sieving, each
fraction is weighed and expressed as a percentage of the total weight of sand in the sieves and
pan. The results can be plotted as a histogram and analysed as shown below.
Skewness is a measure of the
symmetry of the distribution
and is seen visually from a
histogram. Apart from being
a useful descriptive term for
a sediment sample, skewness
is also a reflection of the
depositional
process.
In
general, sediment becomes
more negatively skewed (and
finer grained) along its
sediment transport path.
71
Alternatively, the results can be shown as a cumulative frequency curve by plotting the total
percentage of sand as running or cumulative total against the mesh size of the sieve. The
cumulative percentage is found by adding up the mass percentages as you go along from the
coarsest to the finest fractions. Each point on the curve is plotted directly above each sieve
size. These graphs illustrate some characteristics of sediments, e.g. the histogram will show
the range of grain size and which grain sizes are the most common in terms of weight.
Statistical use of the cumulative frequency curve unfortunately requires that the mm size scale
be changed to a more useful geometric scale called phi (ø) scale. A mesh size of 0.5mm, ø = 1.
When sediment is sifted to and fro by water currents or by wind, particles of the same size
tend to end up in the same place. Sediment which consists of similarly sized grains is said to be
well sorted. Where there has been little sifting of transported grains the sediment consists of
a wide range of particle sizes. Such sediment is poorly sorted. A numerical value of sorting can
be found by using the ø values which correspond to the 16 and 84 % levels. The coefficient
(numerical value) of sorting is given by the following equation:
Sorting = ø84 – ø16
2
Coefficient of sorting is a
numerical measure of the level of
sorting shown by a sediment. The
lower the value the better the
sediment is sorted.
<0.35
0.35-0.5
0.5-0.71
0.71-1.0
1.0-2.0
>2.0
very well sorted
well sorted
moderately well sorted
moderately sorted
poorly sorted
very poorly sorted
On a steep cumulative curve the values of ø84 and ø16 and are close to each other so ø84 – ø16 is
a small number. In other words, the lower the numerical value of sorting the better the
sediment is sorted. On a gently sloping cumulative frequency curve the values of ø84 and ø16 are
widely separated and the high numerical sorting value reflects the poor sorting. In general,
sediments with sorting values less than 0.5 are well sorted while sediments with sorting values
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above 1.0 are poorly sorted. Sediments with sorting values between 0.5 and 1.0 are moderately
sorted. For example, in the histograms and cumulative frequency curves shown in the diagrams
above, sediment A is moderately well sorted (coefficient of sorting of 0.6), while sediment B is
very poorly sorted (coefficient of sorting of 2.4).
Beach and dunes sands are usually well sorted because they have been sifted by water and wind.
River sediments tend to be poorly sorted because current velocities are very variable at
different times and places. Turbidites and glacial deposits are also poorly sorted. The diagram
below shows how sediments can be described in terms of how well sorted they are.
The grain size of a sediment gives some indication of the conditions of deposition, e.g. boulders
are deposited by the strong currents of high-energy environments such as storm beaches while
fine-grained sediment is deposited in low-energy environments such as lagoons and sheltered
bays. The diagram below shows the relationship between velocity of moving air and water and
the diameter of quartz grains carried in suspension.
From the diagram you can get an idea of
the strength of currents which operated
at the time of deposition. It can be seen
that once a particle is eroded from the
channel bed it can be transported by a
lower velocity and thus it can be carried
quite a long way before the velocity
decreases and causes the particle to be
deposited. The graph also shows that the
small particles of less than 250 μm
(0.25mm) in diameter start to require
disproportionately larger velocities to
raise them from the channel bed. This is
because of the cohesion of the particles
towards each other. In addition, as they lie
on the channel bed, they offer less
resistance to water flow than larger
particles and thus require a more
73
energetic stream to lift them.
You will also see that well-sorted sediments tend to be deposited over a narrow range of
current velocities while poorly sorted sediments are deposited over a wide range of current
velocities. In the case of a poorly sorted sediment this may mean that grains of various sizes
were carried by a strong current which suddenly slowed down so that all the material was
deposited at once. To get a get a fuller picture of the conditions of deposition grain size
analysis must be used along with field evidence, study of sedimentary structures and
examination of grain shapes and surface textures. Fossils, too, are very useful indications of
environmental conditions.
MODERN ENVIRONMENTS OF SEDIMENTARY DEPOSITION
A sedimentary environment is an area of the earth's surface where sediment is deposited. It
can be distinguished from other areas on the basis of its physical, chemical, and biological
characteristics. Before studying ancient sedimentary environments, it is helpful to consider the
types of sedimentary environments present on the earth today.
1. Continental Environments
Continental environments are those environments which are present on the continents.
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
Alluvial fans are fan-shaped deposits formed at the base of mountains. Alluvial fans
are most common in arid and semi-arid regions where rainfall is infrequent but
torrential, and erosion is rapid. Alluvial fan sediment is typically coarse, poorlysorted gravel and sand.

Fluvial environments include braided and meandering river and stream systems. River
channels, bars, levees, and floodplains are parts (or sub-environments) of the fluvial
environment. Channel deposits consist of coarse, rounded gravel, and sand. Bars are
made of sand or gravel. Levees are made of fine sand or silt. Floodplains are covered
by silt and clay.

Lacustrine environments (or lakes) are diverse; they may be large or small, shallow or
deep, and filled with terrigenous, carbonate, or evaporitic sediments. Fine sediment
and organic matter settling in some lakes produced laminated oil shales.

Deserts (Aeolian environments) usually contain vast areas where sand is deposited in
dunes. Dune sands are cross-bedded, well sorted, and well rounded, without
associated gravel or clay.

Swamps (Paludal environments) Standing water with trees. Coal is deposited.
2. Transitional Environments
Transitional environments are those environments at or near the transition between the land
and the sea.

Deltas are fan-shaped deposits formed where a river flows into a standing body of
water, such as a lake or sea. Coarser sediment (sand) tends to be deposited near the
mouth of the river; here sediment is carried seaward and deposited in deeper water.
Some well-known deltas include the Mississippi River delta and the Nile River delta.

Beaches and barrier islands are shoreline deposits exposed to wave energy and
dominated by sand with a marine fauna. Barrier islands are separated from the
mainland by a lagoon. They are commonly associated with tidal flat deposits.

Lagoons are bodies of water on the landward side of barrier islands. They are
protected from the pounding of the ocean waves by the barrier islands, and contain
finer sediment than the beaches (usually silt and mud). Lagoons are also present
behind reefs, or in the centre of atolls.

Tidal flats border lagoons. They are periodically flooded and drained by tides (usually
twice each day). Tidal flats are areas of low relief, cut by meandering tidal channels.
Laminated or rippled clay, silt, and fine sand (either terrigenous or carbonate) may be
deposited. Intense burrowing is common. Stromatolites may be present if conditions
are appropriate.
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3. Marine Environments
Marine environments are those environments in the seas or oceans.
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
Reefs are wave-resistant, mound-like structures made of the calcareous skeletons of
organisms such as corals and certain types of algae. Most modern reefs are in warm,
clear, shallow, tropical seas, between the latitudes of 30oN and 30oS of the equator.
Sunlight is required for reef growth because of the presence of symbiotic algae
called zooxanthellae which live in the tissues of corals. Atolls are ring-like reefs
surrounding a central lagoon (such as Bikini Atoll in the Pacific Ocean).

The continental shelf is the flooded edge of the continent. The continental shelf is
relatively flat (slope < 0.1o), shallow (less than 200 m or 600 ft deep), and may be up
to hundreds of miles wide. (The flooding of the edges of the continents occurred
when the glaciers melted at the end of the last Ice Age, about 10,000 years ago.)
Continental shelves are exposed to waves, tides, and currents, and are covered by
sand, silt, and mud.

The continental slope and continental rise are located seaward of the continental
shelf. The continental slope is the steep (5- 25 degrees) "drop-off" at the edge of
the continent. The continental slope passes seaward into the continental rise, which
has a more gradual slope. The continental rise is the site of deposition of thick
accumulations of sediment, much of which is in submarine fans, deposited by turbidity
currents.

The abyssal plain is the deep ocean floor. It is basically flat, and is covered by very
fine-grained sediment, consisting primarily of clay and the shells of microscopic
organisms (such as foraminifera, radiolarians, and diatoms).
IDENTIFYING ANCIENT SEDIMENTARY ENVIRONMENTS
Sedimentary rocks, which are exposed in many areas, contain clues that help us to determine
the sedimentary environment in which they were deposited millions of years ago. By an
examination of the physical, chemical, and biological characteristics of the rock, we can
determine the environment of deposition. Each sedimentary environment has its unique
combination of physical, chemical, and biological features. These features help us to identify
the sedimentary environment in which a rock was deposited. In class, you will be examining hand
specimens of sedimentary rocks, describing their physical, chemical, and biological features,
and then, interpreting their possible sedimentary environments of deposition. Geologists
consider the characteristics that we will study in class, but they also study the geometry of
the sedimentary deposits, the vertical sequence in which the rocks occur, and the
palaeocurrent directions.
Certain generalizations can be made, which help in identifying the depositional environment. For
example, fluvial sequences become finer upward, whereas delta and lacustrine sequences
coarsen upward. These predictable changes occur because the environments migrate over one
another as sea level changes, or as a basin fills with sediment. As a general rule, grain size is
coarser in shallow water "high energy" environments where waves or currents are present.
Waves and currents transport finer sediment offshore into "low energy" environments,
generally in deep, quiet water. Fine grain size indicates deposition in a "low energy", quiet water
environment. In some areas far from shore (or far from a source of terrigenous input) only the
shells of planktonic micro-organisms contribute to the sediment. These microscopic shells
accumulate to form rocks such as chalk or diatomite.
The lithology (composition and texture) of a particular sedimentary rock, its fossils and
sedimentary structures all taken together describe a particular sedimentary facies. The
sedimentary facies concept is useful because it can be linked to a particular set of depositional
processes. Sedimentary facies are a more holistic way to describe sedimentary rocks. For
instance, imagine two lithologically identical sandstones, both being composed of quartz,
medium-grained and well-sorted with moderately well-rounded grains. In other words, they are
sandstones that are lithologically identical. However, one of these sandstones contains a few
bivalves but shows no sedimentary structures and the other is stratified (layered) but devoid
of fossils. These differences define them as two distinct sedimentary facies, deposited by
slightly different processes.
A graphic log is a visual representation of the important and common features of a vertical
succession of sedimentary facies. Graphic logs show the vertical thickness of the individual
beds, the nature of the contacts between each bed, and the lithology, sedimentary structures
and fossils within the beds (i.e. the sedimentary facies). Graphic logs are an excellent way of
recording data because they are form of shorthand which shows the main features of the
sedimentary rocks much more readily than lots of written notes. When a graphic log and its
symbol key is complete, the author, or anyone else with a trained eye, should be able to
recognise easily any of the individual beds in the field. The different features of these
sedimentary successions are shown on a graphic log by the following:
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 Bed thickness – vertical scale in metres shows the thickness of the individual beds.
 Bed contacts – sharp or gradual, or straight or undulating contacts are shown by the nature
of the line drawn.
 Lithology – shown as a symbol within the bed (see key above).
 Grain size – horizontal scale with finest (clays) on left and coarsest on right. Bed of rock
drawn to width indicated on the horizontal scale. Indicates energy of the environment.
 Fossils – shown as symbols just to the right of the lithology column (see key above).
 Sedimentary structures – simplified and accurate version drawn onto the bed (see key
above).
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Sedimentary rocks and their contained structures and fossils (i.e. the sedimentary facies) can
be used to interpret environments of deposition e.g. fluvial, deltaic, shallow marine and deep
marine. The principle features of each environment will be shown on a graphic log.
1. Alluvial Environments
River systems dominate the surface processes that are seen in most of Europe today. They are
important because they are the main routes for the transport of sediments to the sea from
areas of continental weathering and erosion. River systems represent a complex system of
erosion, sediment transport and deposition. Not only is the river channel important but also the
land around the river which floods when the river swells, the catchment area for the rainwater,
and the source of the sediments. The area that the river floods during high water levels, which
is generally flat either side of the river channel, is called the flood plain.
Sedimentary environments that include rivers and their flood plains are referred to as alluvial
environments. The term alluvial refers to the channel itself plus all the area under its influence,
i.e. the river channel and its flood plain, whereas the term fluvial is used to describe the river
channel itself.
The pattern of river development, and the related sedimentation we see in western Europe, is
only one example of a number of patterns that may be found in different parts of the world.
Fluvial erosion and alluvial deposition operate to some extent in all types of continental
environments, including deserts and glaciated regions, although the patterns of development
vary.
There are two types of alluvial environment based on the most common river channel
morphologies. These are; meandering river channels where the river is confined to a single
channel form, and braided systems where the channel is split into a number of smaller channels
which may join or divide again. Braided systems are the sorts of channels more commonly found
in glacial environments.
The sediment derived from weathering is transported down the valley sides into a river channel
by sediment gravity flow processes, known as mass movement. Mass movement is the downslope
movement of slope material under the influence of the gravitational force of the material itself
and without the assistance of moving water, ice or air. There have been numerous attempts to
classify the diverse modes of mass movement. Five fundamental types of movement can be
identified – creep, flow, slide, heave and fall.
Creep is the slow, plastic deformation of rock or soil in response to stress generated by the
weight of overburden. It begins once the shear strength of the slope is exceeded. It occurs at
very slow rates, typically 1mm to 10m/year and is likely to be especially active where weakly
competent materials, such as clays, are overlain by more competent beds. Creep is often a
precursor of slide-type movements. It is important to distinguish the type of creep described
79
here from soil creep and talus creep. Creep acts solely under gravity, whereas soil creep and
talus creep involve heave processes.
In a pure flow, shear occurs throughout the moving mass of material and there is no welldefined shear plane. Flows are categorised as avalanches, debris flows, earthflows, lahars or
mudflows depending on whether they consist of predominantly snow and ice, rock fragments,
sand-sized material, volcanic ash or clay. The slowest type of flow is solifluction which involves
the downslope movement of saturated soil. Solifluction can occur at slope angles as low as 1°
and is particularly active in periglacial environments.
Slides are an extreme widespread form of mass movement, and the term landslide is part of
our everyday vocabulary. In a pure slide failure occurs along a well-defined shear plane. Slides
are nearly always long in relation to their width and depth. They can be subdivided into
transitional slides, which have predominantly straight shear surfaces, and rotational slides in
which the shear plane is curved. Rotational slides are most common where slopes consist of
thick, homogeneous materials, such as clays.
Heave involves cycles of expansion and contraction of material. When this expansion and
contraction is on a slope there is a very slow downslope movement of material.
Expansion (heaves
particles up, 90° to
slope surface)
Slope surface
Contraction
(particles fall back
down vertically)
Expansion causes the surface of the slope to heave at right angles to the original surface and
particles are lifted along the expansion path. Gravity affects the return path of the particle.
The most probable position for the particle on its return is vertically back down. The net result
is a downward movement of the particle on the slope. Expansion and contraction can be caused
by wetting and drying, freezing and thawing, temperature changes and the burrowing activity of
worms and other organisms. Two types of heave can be distinguished on the basis of the size of
the constituent particles – soil creep and talus creep; the latter involves coarser material than
the former. The rate of soil or talus creep depends on a number of factors. It will become
greater with increasing slope angle since this increases the downslope component of movement.
It will also be high on soils containing abundant quantities of clay, which expand significantly on
wetting.
Fall involves the downward motion of rock or, more rarely, soil through the air. Rock can become
detached as a result of various physical weathering processes, including pressure release, and
joint widening by frost action. Rock falls are common in terrain characterised by high, steep
rock slopes and cliffs.
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The source of energy to move water in a river and, indirectly the sediment, is the gravitational
potential energy that results from the difference in height between the upstream and
downstream sections of a river channel. As the water flows down-slope, much of this is
converted to kinetic energy (the energy of the motion of water) which transports sediments
and erodes the river channel. There are three ways in which the sediments move in water
according to their size. The largest grains (bedload) move by sliding or rolling (known as
traction) , the medium-sized to smaller-sized grains (also part of the bedload) move by bouncing
along the bed ( a movement called saltation), the smallest grains (suspension load) are carried in
suspension. All the material transported by the river is known as the load. This load, along with
the kinetic energy of the river is capable of wearing away and removing material from the river
channel; a process known as erosion.
Erosion can be accomplished in three ways. Hydraulic action erodes material in the channel
through the impact of the moving water alone and its frictional drag on the particles lying on
the bed. Usually it is capable of removing only unconsolidated sediment such as sand or fine
gravel. The velocities attained in normal river flow have little effect on hard rocks. River banks
are often undermined by hydraulic action and eventually they collapse into the river.
Abrasion occurs when rock particles already carried by the river as its load, impact into the
river banks and channel bed. Their action is like that of a hammer chipping the rock or like a
file smoothing down the surface and producing very small particles, which are easily
transported. If the particles in the load are large, erosion is more rapid. Small particles tend to
smooth or polish the surface. Since the river is using solid rocks and sediment as tools, abrasion
is possibly the most effective method of erosion and is normally responsible for most of the
downcutting in a river channel. If there is very little load in solid form, abrasion is ineffective.
Rivers at baseflow levels carry very little load in comparison with their ―brown‖ sediment-laden
state during flood conditions. More abrasion occurs during these short violent periods than in
months of low-flow conditions.
Attrition is the reduction in size of the particles in transport as they strike one another or the
bed of the channel. Thus as any particle moves downstream, there is a progressive reduction in
size. In addition, the sharp edges and angles of the particle become rounded, since they are
more exposed. The upper reaches of a river therefore tend to contain large, angular sediment,
whereas the downstream parts have fine, rounded particles.
Meandering river systems comprise a distinct sinuous channel and river banks. As a meandering
river develops through time, the number and size of the meanders will increase, in effect
lessening its slope compared to a river that flows straight down a slope. River meanders
gradually increase in amplitude and migrate downstream, because a helical or corkscrew flow of
water is superimposed on the overall downstream movement of water. This corkscrew flow is
the movement of surface water across the meander towards the outside of the bend and its
return across the base of the channel towards the inside of the bend. The result is that the
main current does not flow straight down the river channel but impinges against each bank in
turn causing erosion of the outside bend. The point of maximum erosion actually occurs slightly
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downstream of the mid-point of the meander thus extending the development of the meander
further downstream.
On the outside of a bend, water flows at a faster speed than on the inside. So, while the
outside bank of a meander is actively eroded, material is deposited on the inside bend of the
next meander downstream where the flow is slower. The sediment is added slowly sideways on
the inside of the bend and forms a bank of sediment called a point bar. At the same time, the
outside bend is being eroded further. The overall result is that lateral erosion on the outside
of the bends is compensated for by the lateral deposition on the inside of the bends, and the
amplitude of the meander increases.
A meander is usually only a temporary landform. If the rate of downstream migration of one
meander is greater than the one further downstream, then the upstream meander may get
closer to the downstream meander leaving only a narrow neck of land between. This neck of land
is easily breached at periods of high discharge so that the river cuts a new channel. A meander
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may also be breached at periods of normal discharge by progressive erosion through the point
bar deposits. In either case, the entrance and exit to an old meander soon become plugged with
deposited sediment so that it is abandoned and forms a lake termed an ox-bow lake. The oxbow lake will gradually fill with fine-grained sediments brought in during flooding and with
decayed vegetation.
At times of increased water level in a river channel, water may flow over the top of the channel
and any sediment carried in it will be deposited due to the sudden reduction in flow velocity
outside the channel. This sediment builds up to form levées (small ridges) on the edge of the
river channel. If the water level rises a lot, the levées may be breached and widespread
flooding occurs; this is called overbank flooding and sediment and water are carried out over
the flood plain. In humid areas, meandering river banks and their flood plains are usually areas
of lush vegetation and good soil development. Plant debris from natural decay, flooding and
fires often becomes incorporated into the sediments being deposited. A variety of animal life,
both invertebrate and vertebrate, will live in the confines of the river system. Animals that live
in the sediment on the point bars, for instance bivalves, might easily be preserved in life
position when they die. The point bar sediments may also preserve burrows made by various
animals. The remains of vertebrates living on the river banks and flood plain will probably only
be preserved during times of flooding when large volumes of sediment are deposited and they
become buried.
Now that you have read about the three-dimensional nature of alluvial systems within a
meandering river system, you are ready to construct an idealized vertical succession of
sedimentary deposits that might be found and to examine the types of succession present in
the geological record.
The most volumetrically significant deposits are those associated with the migration of the
channel. Remember that, as the current is eroding the bank on the outer side of the bend,
sediment is being deposited on the inner bank of the meander and on the next meander
downstream, forming point bars. The net effect, as shown in the diagram below, is that as the
outer eroding bank gradually migrates laterally, so the inner depositing bank gradually builds
out laterally in the same direction as more and more sediment is deposited on the point bar. So,
overall, in cross-section the channel maintains the same width and depth, but moves laterally.
The movement of the channel is slightly greater in some periods than in others due to
variability in discharge and sediment content, resulting in slight pauses in deposition followed
83
by more continuous deposition. These various episodes in the point bar deposits can be seen as
surfaces that mimic the former shape and positions of half the channel; these are called lateral
accretion surfaces (see in the diagram below).
What about the sediments that make up the point bar? As the outer bank is eroded and
collapses, fine-grained sediment is carried off in suspension and the medium-grained sediment
by saltation. The coarsest-grained sediment will only move by sliding or rolling at times of high
discharge so most of it remains at the base of the channel, forming a lag deposit. The lag
deposit may be added to by large clasts that have rolled down the river channel. The currents
in the channel and at the base of the point bar may build up subaqueous dune structures, which
form cross-bedding sedimentary structures. The finest-grained sediments will be deposited on
the upper part of the point bar where the currents are slower and weaker. They often form
current-formed ripples which give rise to small-scale cross-bedded sandstones and siltstones.
Thus, point bars are preserved as fining-upward successions with a lag of pebbles at the base
overlain by cross-bedded sandstones, capped by small-scale cross-bedded sandstones and
siltstones.
If the meander neck breaks to form an ox-bow lake, the oxbow lake will gradually infill with fine-grained sediment and
plant material during flooding. The levées will be composed
mainly of the suspension load fraction of the river deposited
at the edge of the river channel when the water level was high.
Overbank flooding will carry the finest-grained sediments over
the flood plain and as the flow slows down, a thin veneer of
laminated silts and clays is deposited. When the bank breaks
at a confined point, sediments are deposited in a crevasse
splay; these are usually composed of sediment with a mixture
of grain sizes and deposited quite rapidly so they show no
systematic change in grain size. After the flood subsides the
crevasse splay may become colonized by vegetation.
The diagram opposite is an idealised graphic log of the type of
succession that might be produced by a meandering river.
Overall, meandering river deposits typically comprise
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mudstones from the flood plain interbedded with thick sandstones that represent the point
bars. The textures would include poor to moderate sorting and angular to sub-rounded in shape,
making the deposits immature to moderately mature texturally. The sedimentary structures
would include cross-bedding, current-formed ripples and laminae. This meandering river
succession shows sharp-based, cross-bedded, fining-upwards units representing the point bar.
These are interbedded with fine-grained sedimentary rocks deposited during flooding of the
flood plain. The main fossils would be plants, fish bones and freshwater bivalves. However,
these would typically be poorly preserved and be made up of fossil fragments.
2. Deltaic
The diagram opposite shows a graphic log of a typical
deltaic sequence. Overall this succession coarsens
upwards from mudstones and siltstones of the prodelta
to slightly coarser-grained sediment of the delta front
as bedload sediments are deposited at the mouths of
the river channels. These will be capped erosively by
the coarsest-grained sediment of all, alluvial sediments
of the delta plain. On top of these sediment layers
there are soil and peat horizons which have been
deposited in marsh areas of the delta. Sedimentary
structures include inverse graded bedding, crossbedding and current-formed ripple marks. Freshwater
fossils brought down by the river, as well as estuarine
fossils will be in evidence. These will often be poorly
preserved due to the deposition of coarse sands and
the high energy levels in the river channels. The
migration of river channels on the delta plain resulting
in meander cut-off leads to the formation of ox-bow
lakes. The lakes progressively fill with mud during
flooding of the river channels and thus become swamps or marshes, where the favourable
conditions of water and nutrients from a constant replenishment of sediment will support
extensive plant and animal life. The prolific growth of new vegetation means that a vast amount
of dead vegetation accumulates. This dead vegetation will be trapped beneath the new growth
on the marsh or swamp, and the lack of oxygen will lead to the formation of peat. Sediment will
no longer be supplied to the swamp or marsh area because the river channel has changed its
course. As the sediment and plant matter become compacted, sedimentary rocks rich in organic
matter are formed and on burial the peat may become coal.
Deltas form along coastlines where alluvial processes are dominant such that the sediment
supply being discharged by a river into a lake or the sea is so plentiful that it cannot be totally
dispersed into the sea or a lake, by wave, current or tidal action. Rivers deposit their sediment
load as they enter the sea because the mixing of the river water and seawater aids
flocculation. Flocculation is a result of charge-based attractive forces between particles of
clay. In freshwater, the clay particles are negatively charged and repel each other, but in salty
85
waters these charges are neutralised by combination with other positively charged particles
(such as organic matter from bacteria and other dissolved ions) and so the clay particles
flocculate into dense masses and are deposited. As a result of flocculation, larger aggregated
particles are formed which will settle out at even higher current speeds.
In addition, there is a sudden increase in area over which the river water and its load can flow.
As the river moves from its restricted channel to the sea (or lake), the flow expands laterally,
energy is dissipated and therefore the flow decelerates so it is no longer fast enough to move
the bedload and flocculated particles. In addition to the deceleration of river flow as it enters
the sea, there is another important factor that affects deposition of sediment to form a delta;
this is the mixing of water masses of differing densities. River water has a lower salinity and
therefore a lower density than seawater, even with dense sediment being carried in it. Thus, as
a river enters the sea, the river water flows out across the top of the seawater. The result of
these two effects is that the Bedload of the river is deposited close to the coast, but the
suspended sediment is carried well out to sea. The bedload being carried by the river is
deposited fairly rapidly because the frictional resistance caused by the river and seawater
meeting will result in the current speed becoming less and mixing occurring. However, the
suspension load in the river water will be carried further out to sea as part of the surface flow.
The continued frictional resistance between the river water and seawater and their mixing will
eventually slow the currents down and aid flocculation, thus depositing the suspension load.
Deltas have three parts: the delta plain which comprises a flat area dominated by alluvial
deposition. The resulting vertical deposits include alluvial channel fills, overbank muds and the
fine-grained sediment infill of lakes. At the distal edge of the delta plain is the delta front;
sediments are deposited in mouth bars as the rivers emerge into the sea. The most distal part
of the delta comprises the prodelta where the finest-grained sediments are deposited.
Delta plain
Turbidity
current
Delta front
Prodelta
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3. Shallow marine
Shallow marine environments (coasts) can be sub-divided into two main types; one type
dominated by siliciclastic (sand) deposition and another dominated by carbonate (limestone)
deposition. Siliciclastic and carbonate sediments are not mutually exclusive, but one is usually
dominant because the presence of siliciclastic sediments inhibits carbonate production. Mixed
siliciclastic and carbonate successions are found in the geological record, where they are
thought to represent changing conditions, or where sediments are derived from two sources.
1.
Siliciclastic Coastal Environments
Siliciclastic coastlines are accumulations of either sand or gravel running parallel to the
coastline e.g. around the coast of Britain today. There are a number of physical, chemical and
biological processes and conditions necessary for the formation of these coastal environments.
The physical processes include waves, tides and currents (currents may be derived from fluvial
processes, from waves or from tides. Chemical processes along coastline include flocculation
and the formation of soils. Biological processes include the growth of plants, which stabilise
sediment, and the actions of animals.
Siliciclastic coastal environments may be split up into the backshore and foreshore
(constituting the beach), shoreface, offshore transition zone and offshore zone. These areas
are determined by the position of mean high tide and low tide, fairweather wave-base and
storm wave-base.
Vertical cross-section to show the morphological zones in a coastal area.
Above the mean high tide is the supratidal area, which is the area affected only by the highest
spring tides or occasional storm waves (backshore zone). Since this area is subaerial (not
underwater) for most of the time, the sand dries out and is then reworked by aeolian processes
into aeolian coastal dunes. The sediments formed tend to be very well sorted, very well
rounded, fine-medium quartz grains (sands) and have a frosted surface texture. The
sedimentary structures include asymmetrical dunes with internal cross-stratification. The area
87
is also stabilised by salt-tolerant grasses and often fossilise as rootlets in the geological
record.
The mean high tide is the average of the high spring tide and neap tides and, similarly, the
mean low tide is the average of the low spring and neap tides. The mean high and low tidal
marks are important because they define the intertidal area of the shore, which is the area
influenced directly by breaking waves (foreshore zone). The foreshore is characterised by
planar-stratified sands, due to the high energy of the waves, and some-wave formed ripples,
with chevron cross-stratification, due to the oscillating movement of the water as the waves
pass over the beach. As the foreshore is constantly being reworked the medium-coarse sands
of the foreshore zone are typically well-sorted, well rounded, glassy and usually compositionally
mature. The foreshore is often colonised by a range of biota, depending on whether the surface
they are living on is hard or soft. Epifaunal or infaunal bivalves and gastropods are common,
together with various worms and arthropods (e.g. lobsters and crabs). In ancient foreshore
deposits, you might expect to find the hard parts of similar animals associated with their trace
fossils (e.g. worm burrows). The high energy conditions of the foreshore result in many of the
body fossils being broken.
The shoreface zone lies below the effects of breaking waves, but above fairweather wavebase, so sediment is moved by the oscillatory motion of the water as each waveform passes.
Wave-formed ripples are common in the shoreface, resulting in small-scale chevron crossstratification. The texture of the sediment in this area is very similar to the foreshore zone
(well-sorted, well rounded and glassy), but due to the lower energy levels, the grain size will be
slightly smaller (medium-fine sands). A great abundance of biota lives in the shoreface zone.
This means that the primary sedimentary structures may often be destroyed by bioturbation
as a consequence of the organisms burrowing and moving through the sediment. However, the
frequency of reworking by wave currents means that much of this bioturbation may not be
preserved. Biota from many phyla are found in both modern and ancient shoreface areas. They
include benthonic communities of arthropods (e.g. crabs, lobsters and trilobites), molluscs
(bivalves and gastropods), brachiopods, cnidaria (e.g. sea anemones and corals), as well as
pelagic organisms such as molluscs (e.g. cephalopods, such as ammonites) and vertebrates (e.g.
fish). Some of these fossils may well be broken due to the effects of waves.
The area immediately below the shoreface and therefore below fairweather wave-base (the
maximum depth at which there is water movement caused by waves during fairweather
conditions) is the offshore transition zone. The fact that this zone is above storm wave-base
(the maximum depth at which there is water movement caused by waves during storm
conditions) means that sediment is moved around, but only during storm conditions. The
offshore transition zone is generally characterised by finer-grained sediments (fine silts) than
those that make up the adjoining shoreface because the offshore transition zone is a lower
energy environment. This zone is also heavily colonised by many organisms (just like the
shoreface zone but without any broken fossils) and the sediment is usually intensely
bioturbated because of the infrequent disruption by waves. However, the occasional storm
waves can cause hummocks and basins to form with their characteristic hummocky crossstratification.
88
Beyond the offshore transition zone is the offshore zone; this comprises the rest of the
continental shelf below the storm wave-base. The bulk of the sediments deposited here are
silts and clays that are carried there in suspension. Final deposition is aided by flocculation and
filter-feeding organisms ingesting the mud and redepositing it as faecal pellets.
The main sedimentary facies of the siliciclastic coastal environment can be summarised in the
form of an idealised graphic log, as shown below.
Backshore
Zone
Foreshore
Zone
Shoreface
Zone
Offshore
Transition
Zone
Offshore
Zone
2.
Carbonate Coastal Environment
Modern carbonate coastal environments are common in the tropics and subtropics between 30°
north and 30° south of the Equator. Shallow-marine carbonates differ from siliciclastics in the
following ways: sediment is produced mainly by biological and chemical processes in situ; grain
size is not necessarily related to the amount and/or type of transport; sediments are commonly
cemented and dissolved in situ; and most shallow-water carbonates are deposited in warm,
tropical waters.
The chief chemical process to produce carbonate build up is the direct chemical precipitation
of calcium carbonate from seawater. The calcite material precipitated forms micrite
(microcrystalline calcite) and ooids. The main biological processes include: deposition of the
calcareous parts of dead organisms (bioclast); ingestion of carbonate sediment by organisms
which then excrete the waste products ovoid-shaped carbonate pellets (peloids); extraction of
calcium ions (Ca2+) and carbonate ions (CO32-) from seawater and combining them into solid
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crystals of calcium carbonate (CaCO3) to form shells (bivalves and brachiopods) or exoskeletons
(corals) and; erosion by organisms (such as parrot fish and gastropods) as they graze on a
carbonate host (e.g. a coral reef), which in the process breaks down the carbonate into finegrained carbonate sediment to form micrite.
For maximum carbonate production and accumulation, the conditions have to be just right: not
too deep, not too shallow, not too warm or cold, the correct salinity and nutrient supply, not too
much siliciclastic supply and, probably most importantly, the right types of organisms must be
present.
The diagram below shows a graphic log of an idealised shallow marine environment. The main
rock types are limestones: micrite, shelly limestone and oolitic limestone. Large numbers of
fossils will be in evidence, including brachiopods, corals and crinoids. These fossils are usually a
good indication of shallow marine conditions. Corals also require certain environmental
conditions such as warm seas of between 25° - 28°C, a water depth of less than 50m and a
salinity of about 3.5%. Sedimentary structures include cross-bedding and wave-formed ripple
marks.
4. Deep marine
The diagram over the page is a graphic log of an idealised deep marine environment. The log
shows two main rock types; greywackes and black shales. The greywackes are poorly-sorted,
coarse to medium-grained sandstones formed by turbidity currents. Above these are the black
shales composed of fine-grained clay minerals carried in suspension from the continent and
deposited in a low energy environment.
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The black colour is from organic material from dead organisms which have not been decomposed
because of the lack of bacteria and scavengers due to the anoxic conditions (lack of oxygen) in
a deep sea environment. Sedimentary structures include laminae, graded bedding and sole marks
(flute casts and tool marks). Fossils include graptolites which have settled out onto the deep
ocean floor after death and are well preserved due to the fine-grained sediment, low energy
and anoxic conditions.
The key feature of the deep marine environment is the slow build-up of fine-grained sediment,
punctuated by rapidly deposited coarse to medium-grained material. In this low energy
environment there is a constant deposition of fine-grained sediments some of which have been
carried in suspension from the continent. These include weathered volcanic ash, wind-blown
dust and clay minerals derived from the continent; these are collectively referred to as
terrigenous sediment. Importantly, sediment is also produced biologically in the surface waters
of the oceans and then falls as biogenic rain to the ocean floor. Sediments that contain less
than 25% terrigenous material, and which are therefore almost entirely made up of sediments
produced and deposited within the ocean are termed pelagic (e.g.) chalk. Some sediment
deposited on the deep ocean floor will contain a mixture of pelagic and terrigenous material.
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The relatively steep gradients found on the continental slope means sediment gravity flows are
the most common sediment transport processes in the deep sea. Sediment gravity flows are
driven downslope by gravity, and once in motion it moves as a fluid. These flows can travel long
distances: for instance, they may start by collapse of a delta front then travel across the
continental shelf and down a submarine canyon, before finally coming to rest on the deep
abyssal plain. The most common types of gravity flow are turbidity currents. Turbidity
currents consist of a mass of turbid sediment and water. The turbidity current is initiated
when sediment on the continental slope becomes unstable, often due to tectonic activity and
earthquakes. As grains begin to move relative to each other, the water in the pore spaces
between them seeps across the grain boundaries. This lubricates the grains, reducing the
friction between them. The sediment pile is now no longer fully grain-supported because of the
thin films of water separating the grains, and it becomes fluidised. The whole sediment mass
becomes effectively a high density, turbid fluid within which individual grains are suspended by
the turbulence of the flow. Because of its high density it can begin to flow down-slope beneath
the surrounding lower density water, reaching speeds of 50 km/hour or more. A turbidity
current is not uniform along its length. It can be divided into ―head‖, ―body‖ and ―tail‖ sections.
In the early stages of the flow the head and body travel at different speeds. The body travels
faster than the head, so water and sediment are constantly being forced into the head area
from the body. The pressure created in the head forces the additional fluid up and around,
returning it back to the body. The body, therefore, is a mass of turbulent eddies which
maintain the sediment in suspension.
It can be imagined how awesome a turbidity current must be, with a head maybe hundreds of
metres high, thundering down the continental slope at speeds in excess of 50 km/hour. The
head is a region of devastating erosion as the water, carrying sand and gravel derived from the
continental shelf, is swept up and around. It rips up the underlying mud and shapes it into balls
which are incorporated within the first sediments to be deposited as the current slows. Other
common erosional features are flute marks and various tool marks. Flute marks are formed
when spiral eddies develop in the flow as it passes over the irregular sea-bed. The eddies scour
elongated hollows in the underlying muds which become shallower and wider down-current. As
sediment is deposited from the current, it fills the hollows so that the flute marks are
preserved as reversed-relief flute casts on the bottom bedding planes of the deposited
sediments. For this reason they are an example of a sole mark, a structure on the base or sole
of a bedding plane. Tool marks are also common. These are indentations of the underlying
sediment produced by objects or ―tools‖ being carried along in the bedload by the current.
These may range from prod and bounce marks to linear grooves, caused by anything from small
pebbles and shells. Once again, these tool marks may be preserved as casts on the base of a
bedding plane.
A turbidity current transports a mixture of grain sizes, including pebbles, coarse-grained sand,
clay, fine-grained sand and granules. As the current slows, the coarsest-grained sediments will
be deposited first because they are heavier. The succession from the base would be pebbles,
granules, coarse-grained sand, fine-grained sand with clay at the top. Deceleration of a
turbidity current begins as soon as the gradient of the slope begins to decrease. As a turbidity
current decelerates deposition begins from the body section, immediately behind the erosive
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head. As the current speed decreases, fining-upward succession will form with a predictable
pattern of sedimentary structures in a succession of beds known as a turbidite. Each bed
represents a particular part of the flow, and the size of the clasts and sedimentary structures
are diagnostic of particular parts of the flow and the speed of the turbidity current. This
characteristic succession of beds was first recognised and described by Arnold Bouma and is
thus called a Bouma Sequence.
As a fast-flowing turbidity current passes over a surface, its high energy means that it will
usually erode structures into the underlying bed. Because these structures are found as casts
on the base of the bed, they are collectively termed sole marks or sole casts. They include tool
marks, formed as sharp or large clasts scrape and dig into the underlying bed, and flute marks,
which form due to small eddy currents scouring the surface. At this stage, the turbulent
turbidity current maintains all of its sediment load in suspension by fluid turbulence so no
deposition occurs.
As the speed of the turbidity current begins to decrease, the larger grains settle towards the
base of the flow and start to collide with each other, slowing the base of the flow down
further. Sole marks will continue to form. Meanwhile, the grains in the upper part of the flow
continue moving in the fast turbulent flow. Eventually, the concentration of grains and the
slowing down of the base of the turbidity flow will cause the lowermost layer to be deposited.
This forms layer A in the succession above; and it contains the coarsest grains, and is generally
massively bedded although it may show some fining-upwards.
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As the current slows down further, the medium-sized particles start to settle out. The current
is still quite fast-moving (upper flow regime), and the sediments are deposited in straight
layers. This forms layer B in the succession above. With further slowing down of the current
into the lower flow regime, current-formed ripples will form and so create cross-bedding.
Layers A, B and C represent a greywacke, which tend to be immature both texturally and
compositionally. Continued deceleration allows both fine-grained and medium-grained sediment
to be deposited in laminae forming straight layers. This forms layer D. As the flow comes to a
stop, the final stage is to deposit the finest-grained sediment still held in suspension; this
forms layer E(t). Layer E(h) is the background sedimentation of fine-grained pelagic sediment
deposited from suspension and which has not been carried by the turbidity current. Layers D
and E represent a shale.
A whole series of Bouma Sequences stacked on top of each other, each representing a single
turbidity current. The diagram above shows a complete Bouma Sequence, however, in reality,
complete sequences are not always found because of the changing conditions of the flow and
the variety of grain sizes carried in different turbidity currents.
Some sedimentary facies, fossils assemblages and depositional environments indicate a
particular climatic zone e.g. desert environments, glacial environments and tropical shallow
seas.
1. Desert Environments
Deserts conjure up an image of a vast area of sand dunes, but in fact this is only one of several
sub-environments found in deserts. Deserts are defined as areas where the average rate of
evaporation exceeds the average rate of precipitation. This lack of water is the reason why
they contain very little plant and animal life. Lack of vegetation and water also means deserts
are dominated by physical weathering and that when precipitation does come erosion is rapid,
thus sediments in deserts can be easily moved around. Deserts occur in both hot and cold
climatic areas of the world. Hot deserts cover a larger area of the Earth and it is here that the
constant heating during the day and cooling at night set up stresses within the rocks which lead
to increased physical weathering.
The dominant agents of transport and deposition – wind and water – can be used to divide
desert features and sedimentary deposits into two categories. The first category is those
features and sedimentary deposits formed mainly by wind; these include aeolian sand dunes,
extensive area of which may be called sand seas, together with large areas of bare rock and
rock debris. The second category is the features and sediments formed by the transient
passage of water; these include alluvial fans, river valleys and temporary lakes (playa lakes).
Aeolian processes are in many ways similar to water transportational processes. However, the
difference in density and viscosity between air and water leads to some differences in
sediment behaviour. Saltation of fine to medium-grained sands and suspension of silts are the
dominant processes during aeolian (wind) transport, whereas rolling and sliding, as well as
saltation and suspension transport are important processes during water transport. As sand
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grains saltate to greater heights in the air than water, and strike the grains on the surface
beneath them at higher speeds, they cause coarser grains to move forwards by surface creep.
This process does not occur in water. As might be expected, these differences in mode of
transport lead to differences in the nature of the aeolian bedforms, and how they develop. The
main controls on aeolian bedform morphology are wind speed, wind direction, and sediment
supply.
Despite various experiments with sands in wind tunnels, the process of aeolian ripple formation
is not well understood. One suggestion is as follows. Sand grains begin to saltate as soon as the
wind speed at a particular point on the bed is sufficient to start them moving. As aeolian sands
are well sorted, the grains will all be fairly similar in size and so they will tend to saltate around
the same distance, descending to hit the surface again in more or less the same place. On
impact, they set in motion the grains that were too coarse-grained to saltate. So the coarser
grains move by surface creep, and the impact makes the finer grains saltate again. As this
process is repeated, a series of regularly spaced impact sites develops on the bed. Sand is
pushed forward at these sites by surface creep and so impact ripples begin to form. These are
rarely preserved in the sedimentary record.
Finer-grained sand migrates rapidly over the ripples by saltation and some is deposited and
trapped in the ripple troughs, but coarser-grained sand moves more slowly as it is pushed up
the up-wind side of the ripples by surface creep. This process results in coarser-grained sand
becoming concentrated at the ripple crests. This contrasts with current-formed ripples where
coarser-grained material is located in the ripple troughs. Impact ripples are slightly
asymmetrical, like current-ripples, but they are flatter and their crests are long and relatively
straight with occasional bifurcations.
Unlike subaqueous dunes, aeolian dunes do not grow from ripples in response to increasing wind
speed. As wind speeds increase, ripples are washed out and eventually replaced by a flat bed
again. Aeolian dunes take much longer to develop than ripples and, provided there is enough
sand available, may grow to spectacularly greater heights than the subaqueous varieties,
because they are not constrained by flow depth. Furthermore, actively migrating sand dunes
may have migrating impact ripples superimposed on them, and smaller dunes may grow on the
backs of larger ones. For aeolian dunes to grow it seems that there has to be some obstacle in
the path of the sand to begin with, such as an isolated clump of vegetation or a rock. As the
airstream flows over the obstacle, a wind shadow is created on the down-wind side, containing
localised eddies where sand is deposited to build up drifts. When sufficient sand has
accumulated the drift becomes independent of the original obstacle and starts migrating as a
dune.
Saltating sand grains bounce their way up the shallow slope of a dune to the crest where they
periodically avalanche down the steep lee slope by a process known as grain flow. On this steep
lee slope of the dune shallow gullies occur, these are the sites of small avalanches where grain
flow is takes. The sand on this steep slope often exceeds its stable angle of rest, and so begins
to move under the influence of gravity. Once in motion, the flow is sustained by grains colliding
and ricocheting, and rolling over each other. During this process the larger grains are
preferentially moved upwards to the top of the flow and the finer grains move downwards.
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Once the flow has stopped the grains, even though of very similar size, have been sorted, giving
rise to cross-bedding. Over time more and more sand is transported up the shallow windward
slope and deposited on the steep lee slope, where it becomes unstable and avalanches forming
another set of cross-beds on top of the previous set. Thus, the dune gradually migrates in the
direction of wind flow.
Aeolian dunes in the sedimentary record look very similar to those formed in water. There are a
number of pieces of evidence to distinguish between the two. To begin with, the aeolian crossbeds are much thicker, commonly in the order of 5-10m, whereas subaqueous cross-beds are
normally less than 2m thick. Secondly, although the cross-beds in aeolian dunes at the base of
each set make only a low angle with the horizontal, they curve upwards more steeply than
subaqueous cross-beds. This is because the steep slope of an aeolian dune may be stable up to
an angle of rest of about 35°, whereas the angle of rest of subaqueous dune slopes is usually
much less. Thirdly, in aeolian dunes cross-bedding is much better developed due to the fact
that the process of grain flow is much more efficient in air than in water, so the sorting
process is more pronounced. Other pieces of evidence would be the frosted surfaces to the
aeolian sand grains, due to the lake of water to cushion the contacts between grains. The
aeolian sandstones would also be texturally very mature, although sands which have been
subjected to long periods of wave transport would also be texturally mature. A lack of marine
fossils would also be a useful clue.
In order to preserve the dune sands, so that eventually they will become sandstones in the
geological record, there has to be net accumulation of sand. So the sands have to be protected
from the erosive and transportational power of winds. This can be achieved in several ways. (i)
Through deposition at or near the water table (wet sands simply do not blow away), combined
with subsidence, allowing more and more sand to accumulate. (ii) The growth of vegetation or
the formation of surface cements, protecting the sands and preventing them from blowing
away. The red iron oxide (haematite) coating around the grains of desert sandstones could be a
form of early cement that may well have helped to stabilize these desert sandstones so they
could be preserved. (iii) If the supply of sediment is exceptionally high, then the sands may
accumulate because there is insufficient wind to remove them.
So, in summary, wind processes are important in shaping, transporting and depositing sand
grains, together with moving all loose sediment away from an area and leaving behind bare rock
and clasts that are too large to be moved. The wind is also important in a desert in helping any
water that may be present to evaporate. Water causes some of the most spectacular and
dramatic features of desert erosion because most of the sediments are not bound together by
vegetation, fine-grained sediment or cement. Rainfall is sporadic, but when it occurs over a
short seasonal period or every few years it may well be torrential. In desert regions within or
near to mountain ranges, a rainstorm may be catastrophic. The sudden rainfall carries large
volumes of accumulated sediment down from the mountains and into the desert area and may
erode a steep-sided, narrow valley called a wadi (Arabic for watercourse). As the flood waters
subside by infiltration and evaporation, sediment is deposited. It is usually poorly sorted and
may well show only crudely defined bedding. For much of the time, most wadis are dry valleys
that are only periodically occupied by water. There are two other morphological features
formed by the presence of water in a desert; alluvial fans and playa lakes.
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Alluvial fans are localised deposits that form where a river or stream loaded with sediment
emerges from a confined mountain valley (such as a wadi) onto a flat lowland plain. Alluvial fans
have a basic semi-conical shape. The apex of the cone points up the valley it has come from. As
the water flows out from the steep valley and onto the flat plain, the sudden decrease in
gradient is accompanied by extensive deposition from sediment gravity flows. Although alluvial
fans are most common in semi-arid and glacial climatic regions where vegetation is sparse and
run-off is seasonal, they can also form in humid tropical regimes. Alluvial fan deposits are
generally poorly sorted, matrix-supported and both texturally and compositionally immature,
but they may show varying degrees of layering or graded bedding depending on how much water
remains in the flow. The actual degree of grain roundness is a function of the sediment source,
how far the sediment gravity flow travelled and whether it was fairly viscous so the grains
were cushioned. The fan is traversed by a network of rapidly shifting braided streams and so
the sediments in the channels comprise gravels, sandy gravels and sands. The sands may be
arkosic if the nearby mountains contain feldspar-bearing rocks and chemical weathering and
grain transport are limited. A feature common to all alluvial fans is that the sediments become
finer-grained with increasing distance from the apex of the fan. Alluvial fans are a prominent
feature of most braided river systems in mountainous areas and where mountains meet flat
desert areas.
Temporary lakes, often termed playa lakes, can form in any depressions in the desert surface
as it rains and are common at the base of alluvial fans. The playa lakes become infilled with
fine-grained sediments that have been transported in suspension either by wind or water. If
sediment has been carried by water down alluvial fans, the finest-grained sediments will be
carried the furthest and will be deposited at the distal (furthest from the apex) part of the
fan in the flat area where the playa lake forms. The resulting sedimentary rocks are typically
laminated mudstones and siltstones. Desiccation cracks are very common and form as the finegrained sediments dry out. The water may support organisms, such that both trace and body
fossils may be found. If the water does not drain away, but is allowed to evaporate slowly,
evaporite minerals such as gypsum and halite will be deposited and a salt pan is formed.
2. Glacial Environments
Glaciers act as a conveyor belts for large volumes of eroded material. Unlike most other
erosional depositional systems, the thickest sediments accumulate farthest from the source
and, furthermore, even the most distal sediments contain coarse boulder material. The manner
in which sediment is transported out of the glacier system is largely dependent on the thermal
regime. Cold, sliding glaciers carry a heavy basal debris load which is deposited as till. In
contrast, highly dynamic temperate glaciers have relatively little basal debris but high
supraglacial load, most of which is modified by meltwater soon after deposition. An
understanding of the transport paths through a glacier provides a basis for interpreting the
depositional processes and deriving the source areas of the sediment.
Debris carried at a high level (supraglacial and englacial) in a glacier is derived from a number
of sources, however, of these, rockfall is volumetrically by far the most important component
on the surface of most valley glaciers. It tends to accumulate as lateral moraines, forming
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medial moraines where ice-flow units combine. Commonly, these merge towards the snout, giving
rise to a totally debris-covered surface. Rockfall debris is largely derived from frostshattered processes and is predominantly coarse-grained and angular character. Once this has
fallen onto the glacier it will be transported passively, with little modification during transport.
Incorporation of debris into the basal ice of a glacier will be subjected to erosional processes.
Erosion at the bed of a glacier comprises various processes, including crushing and abrasion,
which combine to produce a sediment very different from that carried on the surface. Debris
eroded in this fashion is initially transported in a basal zone of traction, where particles
frequently come into contact with the glacier bed, so that large forces are imparted on both
particle and bed. The clasts carried in basal ice have a significantly higher roundness
(predominantly subrounded and subangular) than those carried in high-level transport. In
addition, the boulders tend to have relatively smooth faceted surfaces, although sharp-edged
fractures on otherwise smooth boulders are common. Striated abrasional facets are also
common. Also, in this basal debris the grain-size distribution is polymodal, but overall is
depleted in the coarse fraction and enriched in the fine fraction in comparison with debris
undergoing high-level transport. The high proportion of fine material is the result of
comminution of larger particles in the zone of traction, a process characteristic of a crushing
mill, and one that produces the rock flour that gives glacial meltwater its milky appearance.
In many areas, more sediment is transferred out of the glacier system by meltwater than
directly by the ice. Huge quantities of sediment may be involved, but the rate at which it is
moved is strongly dependent upon temperature, which in turn is a function of the season and
the time of day. Meltwater streams are powerful agents of erosion, especially beneath the ice
where the water may be under high pressure, and consequently have a large impact on the
sediment they transport. The basic mechanisms of deposition from meltwater are essentially
the same as those for ordinary fluvial systems. However, there are some important
differences; for instance, tunnels within ice are not usually free to migrate laterally as fluvial
channels may do and the rapid fluctuations in meltwater discharge can produce equally rapid
vertical changes in the calibre of sediment deposited which may range from large boulders to
sand. The key characteristics which distinguish water-laid fluvioglacial sediments from glacial
deposits is the absence of fine particles due to sorting and the presence of stratification.
Meltwater deposits may be less easy to differentiate from ordinary fluvial sediments, although
the glacial sediment source and the typically short distance of transport before deposition
means that the former are usually less rounded.
Deposits of glacial origin take characteristic forms. The most well-known is called boulder clay,
or more correctly, till. Boulder clay is a rock flour produced by the grinding action of ice sheets
and glaciers as they pass over the solid bed rock. The ice also picks up boulders and rock
fragment which it uses at its base as grinding tools, and these boulders themselves can
eventually be reduced to flour. When the ice retreats it leaves behind a mixture of clay and
boulders, hence the term boulder clay. The boulders can vary in size and vary in shape from
rounded to angular, but normally subangular to subrounded varieties predominate. Some are
smooth and occasionally marked with scratches. Many boulders have travelled considerable
distances from their original source where they were first picked up by the ice, and are called
glacial erratics. In East Anglia there are erratics from Norway, and in Shropshire many
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different granite erratics have come from Scotland and the Lake District. Till is the term used
more often nowadays than boulder clay. Frequently it is sandy, or a mixture of sand and clay.
Clay tills (true boulder clays) are sticky, blue clays (orange when weathered) which can be a
gardener‘s nightmare in parts of Lancashire!
Glacial moraines are formed on the top of a glacier as material falls on to its surface from the
valley sides, and also at its snout. When the ice melts they are left behind on the ground
surface over which the glacier or ice sheet has been moving. Moraines are usually irregular
dumps of poorly sorted, angular and coarse-grained material. Examples of morainic gravels can
be found in the mountain areas of the Scottish Highlands and the Lake District. Drumlins are
well-known landforms left behind by moving ice sheets. Glacial moraine deposited by the ice is
further moulded into smooth elongated hills with the blunter end towards the ice. They have
the shape of half an egg, typical dimensions being 400m by 100m, and often occur in large
numbers in drumlin fields, and produce a type of hummocky land surface called a basket of
eggs, well known in Ribblesdale and the Eden Valley in northern England.
As well as glacial tills and moraines which are all deposited directly from the ice, melting ice
sheets and glaciers produce large amounts of fluvio-glacial sands and gravels which are washed
out of the base and the edge of the ice by meltwater streams. Kames are masses of gravel
dumped at the edge of an ice sheet and eskers are long, narrow sinuous ridges of sand and
gravel formed in tunnels within the ice by melt waters flowing between the ice and the ground
surface. The finest examples of these are in north-east Scotland near Inverness where the
highest esker in Britain, the Kildrummie esker, is found at Torvean (70m high). Fluvio-glacial
deposits are generally sorted, rounded and stratified (layered) deposits.
Areas adjacent to or above ice sheets during the Ice Age were affected by what is called a
periglacial climate (peri means around the edge of, as in perimeter). Here, extreme cold at
night can be followed by relatively high temperatures in the day. This causes freeze-thaw
action when water which has been trapped in joints of crags expands at night, causing the rocks
to become shattered as the process is repeated day after day. Large scree fields form around
the frost shattered crags. The screes above Wastwater in Wasdale in the Lake District are a
classic example. In areas where bare rock does not occur at the surface, freeze-thaw action
causes the whole hill surface to become broken up to a depth of a few metres, into a coarse
angular deposit called head. Large areas of ground are affected by very slow gravity
movements of these fractured surface deposits, a process called solifluction.
3. Tropical Shallow Marine Environments
Shallow-marine carbonate sediments (such as limestones and chalk) are an important feature of
the sedimentary record and were more common in the past, chiefly in warmer climatic
conditions and periods of high sea-level when shallow seas covered extensive areas of the
continent, forming shallow-marine environments. Some shallow-marine carbonate are forming
today between 30°N and 30°S of the equator, around the Bahamas, South Florida and the
Arabian Gulf. There are many factors controlling the growth of modern coral reefs and it is
likely that these same factors exerted an influence on coral in the past. For coral-reef growth,
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these factors are: (i) water temperature – optimum growth is around 25ºC, (ii) water depth –
most growth takes place within 10 metres of the surface, (iii) salinity – corals cannot tolerate
great fluctuations and, (iv) turbidity and wave action – coral growth is favoured by intense
wave action and an absence of terrigenous (sediment from the land sources) silt and clay.
Shallow-marine carbonate sediment is produced mainly by biological and chemical processes in
situ and is not transported over large distances. Climate and organism biology are the two main
controls on the type of sediment produced in these areas. Sediments which owe their origins to
biological processes are known as biogenic sediments and the most important of these,
geologically speaking, are some types of limestones. Limestones are composed predominantly of
calcium carbonate (CaCO3) and the process of their formation begins when Ca 2+ is released into
solution by chemical weathering of silicate minerals such as plagioclase feldspar and
hornblende, and when pre-existing limestones are dissolved. The weathering of limestone also
releases bicarbonate (HCo3-) ions into solution. Some of the bicarbonate ions partially
dissociate to produce carbonate ions (CO32-). Under appropriate conditions, marine organisms
can extract Ca2+ and CO32- from seawater and combine them into CaCO3 to construct their
shells. Once these animals die their discarded hard parts accumulate on the sea-floor to form
bioclasts where they may make up the carbonate sediments that will form the next generation
of limestones. Another way biogenic sediments can be formed is from the excrement of certain
marine animals such as shrimps and snails. These may be deposited in situ or, if sufficiently
hardened, they may be reworked by waves or currents. These faecal pellets form rounded or
ellipsoidal grains, typically 0.1 – 0.5mm in diameter and are usually devoid of any internal
structure, and are termed peloids.
Carbonate sediments can also be produced chemically by direct precipitation from solution. One
reason for this is that in warm tropical environments seawater may begin to evaporate. As it
does so, it becomes supersaturated with various dissolved ions, such as Ca 2+ and CO32-, and
CaCO3 is spontaneously precipitated as mud-sized particles that are typically less than 10
microns in size. These grains, along with deposited bioclasts, are moved around on the seabed to
produce fine-grained carbonate sediment. After deposition and burial, the carbonate mud
recrystallises to produce very fine-grained calcite crystals, < 4 microns in size, known as
micrite (an abbreviation of two words: microcrystalline calcite). The very fine grain size of
micrite implies that the limestone must have been deposited from suspension in low energy
conditions. The spontaneous precipitation of CaCO3 may also occur directly onto small sediment
particles, such as small bioclasts, quartz grains or peloids, which act as a nuclei for the process.
As these nuclei are rolled to and fro over the sea-bed by water movement, layers of calcium
carbonate build up round them through constant precipitation so that they grow in size. The
end product of this process is a collection of creamy white, spherical grains, typically around
0.5mm in diameter. Because of their shape, these grains are called ooids, a word meaning egglike. Limestones composed predominantly of ooids are described as oolitic limestones. The
internal structure of the ooids consists of a series of concentric layers. Within each layer are
fine lines radiating outwards. By contrast, peloids are more or less uniform in appearance,
devoid of any internal structure. The rounded shape of ooids results from the way in which
they are rolled around on the sea-bed, indicating a shallow-marine environment. This shows that
the shape of the ooid is unrelated to the length of time they have been in transport, unlike
100
fragmental or clastic grains transported by rivers and waves. The observation that ooids reach
a maximum size of about 0.5mm suggests that at this critical stage of growth they probably
begin to suffer abrasion as fast as they form, which may help to explain why oolitic limestones
may appear so well sorted. From this discussion it can be seen that it is inappropriate to apply
the concept of textural maturity to all limestones. Many limestones accumulate in situ and
contain particles ranging from mud to granules in size, or even coarser. Some even contain large
fossil organisms, preserved in their life positions. It is best not to describe a limestone in
terms of textural maturity unless there is good evidence that it consists predominantly of
fragmented bioclasts that have been well sorted and rounded during transport.
Wave-formed ripples and current-formed ripples can also develop in carbonate sediments,
especially those made up of sand-sized grains, such as ooids. This means that some limestones
may also show cross-bedding, by the same processes that produce these sedimentary
structures in clastic sediments. The fossils found in tropical shallow sea are shown below.
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SEDIMENTARY ROCK CLASSIFICATION
For the identification of sedimentary rock types in the field, the two principal features to note
are composition/mineralogy and grain size. On the basis of origin, sedimentary rocks can be
grouped broadly into 4 categories.
Clastic/Fragmental
Conglomerates
Breccias
Sandstones
Mudstones/Shales
Biochemical/Biogenic
Limestones
Chalk
Coal
Chemical
precipitates
Evaporites
Volcaniclastic
Tephra
Tuff
Conglomerates and breccias are both poorly sorted and consist of large clasts larger than 2mm
(from 2mm to 4mm the clasts are called granules, from 4mm to 64mm the clasts are called
pebbles, from 64mm to 256mm the clasts are called cobbles and above 256mm boulders), more
rounded in conglomerates, more angular in breccias, with or without a sandy or muddy matrix.
Conglomerates and breccias are deposited in a range of environments, but particularly glacial,
alluvial fans and braided streams. Clasts in a glacial environment may have striations and
scratches. Fluvial conglomerates may be reddened due to oxidation. Conglomerates deposited in
beach and shallow-marine environments may contain marine fossils.
Sandstones are composed of medium sized grains between 2mm and 63 microns (μm) in
diameter. Although within this range sandstones can be subdivided into very coarse (1mm to
2mm), coarse (1mm to 500 microns), medium (500 microns to 250 microns), fine (250 microns to
125 microns) and very fine (125 microns to 63 microns). Bedding is usually obvious and
sedimentary structures are common within the beds and upon the bedding surfaces. The
accepted classification of sandstones is based on the percentages of quartz, feldspar, rock
fragments and clay matrix. With a hand-lens it is possible to estimate the amount of matrix
present in a sandstone and thus determine whether it is an arenite (a clean sandstone such as
orthoquartzite or desert sandstone) or a wacke (>15% clay matrix such as greywacke).
Arenites are typical of high-energy shallow-marine environments, and also aeolian (wind-blown)
sand-dunes in deserts.
Orthoquartzite (quartz sandstone), as the name suggests, are made almost entirely of quartz
grains. These rocks are well sorted and contain well rounded grains. A typical environment for
an orthoquartzite would be a shallow sea where current action constantly reworks the grains by
rolling them back and forth. Deposition tends to be slow so there is plenty of time for minerals
more susceptible to weathering and erosion, such as feldspar, to be broken down completely
rather than buried in the sediment. Only the strong and less susceptible quartz grains remain
and they become rounded and sorted as the movement continues. The current may also produce
ripple marks and cross lamination. Orthoquartzites are especially likely to form where a supply
of rounded quartz grains is arriving from the erosion of older sandstone rocks. Since only
quartz is present, the colour of orthoquartzite is commonly white or pale grey, especially those
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of shallow-marine environments. Most orthoquartzites are cemented by quartz which
crystallised to fill spaces between the grains.
Desert sandstones formed in the desert environments of the past are quite common in Britain.
These rocks are very well sorted and contain very well rounded grains. Since desert sandstone
is made almost entirely of quartz grains, it is really a special form of orthoquartzite. However,
a distinctive feature of desert sandstone is the red/brown cement of iron minerals such as
haematite that coats the quartz grains due to oxidation. This cement is rather weak, so desert
sandstones tend to be friable. Outcrops of desert sandstone often show large-scale crosslamination because the sand built up into dunes.
Greywacke actually means ―grey grit‖. They are impure sediments with a high proportion of fine
clay matrix. These sand-sized grains are angular, moderately poorly sorted and of mixed type.
Greywackes are mostly hard rocks, which are light to dark grey in colour due to the abundant
clay matrix. Feldspar and small rock grains are also common. Many greywackes were deposited
by turbidity currents in relatively deep-water basins and show sedimentary structures typical
of turbidites (sole structures, graded bedding and internal laminae). Greywackes commonly
grade upwards into mudstones as part of a turbidite sequence.
Another type of sandstone is arkose. Arkoses can be recognised by a high percentage of
feldspar grains. Many arkoses are red or pink, in part due to the presence of pink feldspars but
also through iron oxide pigmentation. In many, grains are sub-angular to sub-rounded and
sorting is moderate; a considerable amount of matrix may be present between grains. Relatively
rapid erosion and deposition under a semi-arid climate produce many arkoses (chemical
weathering would cause the feldspar to decompose and form clay minerals). Fluvial systems
(alluvial fans & braided streams) are typical depositional environments for arkoses, especially if
granites are exposes in the source area.
Siltstones are well-sorted, fine-grained rocks, which are made up of the smallest and finest
quartz grains (63-4 microns). A siltstone can be distinguished from a mudstone by its gritty
feel when rubbed on your teeth! Many siltstones show exceptionally thin bedding which is called
lamination.
Shales are very well-sorted and fine-grained sedimentary rocks dominated by clay minerals (<4
microns), although other minerals may be present. Clay minerals belong to the mica family and
each tiny crystal (about 0.005 mm across) has a flat plate-like shape. Shales are formed when
these clay minerals are deposited slowly, for example in deep quiet waters, where they settle
from suspension and come to rest with their flat ―plates‖ lying horizontally. During burial,
pressure from overlying sediment above squeezes the sediment and aligns the minerals even
more. The resulting rock is called shale. Since it is made of layers of clay minerals, shale is
easily broken into thin flat layers and is said to be fissile. It is important not to confuse shale
with slate. Slate is a metamorphic rock and splits into much stronger and harder sheets. Many
shales are black in colour due to the build-up of organic matter. Organic matter (dead marine
organisms not decayed due to deep-water, anoxic conditions) may reach several percent or
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higher, and with increasing carbon content the shale becomes darker and eventually black in
colour.
Mudstone is also a very well-sorted and fine-grained sedimentary rock which is produced when
clay minerals are deposited more rapidly and collect in a ―jumbled‖ fashion. Provided there is no
great pressure during burial the clay minerals do not become aligned so the rock is not fissile
and breaks unevenly. Mudstones can be deposited in practically any low energy environment,
particularly river floodplains, lakes, low-energy shorelines, lagoons, deltas, marine shelves and
deep-ocean basins.
Limestones are composed of more than 50% CaCO3 and so the standard test is to apply dilute
HCl; the rock will fizz. Many limestones are a shade of grey, but white, black, red, cream and
yellow are also common colours. Fossils are commonly present, in some cases in large numbers.
The principal components of limestones are carbonate grains (aragonite & calcite), lime
mud/micrite (micro-crystalline calcite), cement (calcite), ooliths and bioclasts (skeletal
grains/fossils). Skeletal grains are the dominant constituent of many limestones. The types of
skeletal grain present depend on environmental factors during sedimentation (water depth,
water temperature and salinity) as well as on the state of invertebrate evolution and diversity
at the time. The main organism groups contributing skeletal material are molluscs (bivalves),
brachiopods, corals, and crinoids. Ooliths are spherical to sub-spherical grains, generally in the
size range 0.2mm to 0.5mm, but reaching several millimetres in diameter. Ooliths consist of
concentric coatings around a nucleus, usually a carbonate particle or quartz grain. These grains
are rolled back and forth by the sea consequently picking up CaCO 3 to from oolitic limestone.
Although the majority of limestones are shallow-marine in origin, some limestones are deposited
in deeper water, and in lakes. Micrite is the matrix to many bioclastic limestones and the main
constituent of fine-grained limestones. It consists of carbonate particles mostly less than 4
microns in diameter. This carbonate mud is formed through the disintegration of carbonate
skeletons. (See also page 85 – Tropical shallow marine environments‖)
Evaporite deposits include gypsum, anhydrite and halite. Gypsum is quite common whereas
anhydrite and halite are only found in very arid areas.
Volcaniclastic deposits can also be classed as sedimentary rocks. Volcaniclastic sediments are
generally known as tephra and are formed by the fragmentation of material due to explosive
volcanic activity. Tephra mostly consists of pyroclastic fragments that include fresh lumps of
lava, and glass which fragment to give glass shards. On the basis of grain-size these pyroclastic
fragments can be divided into ash (<2mm) which forms tuff when lithified, lapilli (2mm to
64mm) which forms lapilli stone, and blocks and bombs (>64mm). Blocks are ejected as a solid
and form volcanic breccia (fine to medium-grained) and agglomerate (coarse-grained). Volcanic
bombs consist of rounded blobs of lava which are deformed into strange shapes if they were
still soft when they landed. Based largely on shape, spindle, bread-crust and cow-dung bombs
can be distinguished.
Sedimentary rocks form in layers. The thinnest layers of sedimentary rocks, less than 1cm in
thickness, are termed laminae (singular: lamina). Layers thicker than this are called beds. Beds
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are rock units with a sheet-like geometry, and their boundaries may be planar or irregular. Like
laminae, the term bed carries a thickness connotation, and so descriptions such as thinlybedded or thickly bedded have precise meanings. Thinly-bedded means beds 1-10cm thick,
medium-bedded describes beds 10-30cm thick, and thickly-bedded refers to all beds more
than 30cm thick. Boundaries between different beds can be recognised by changes in
mineralogical composition (e.g. a change from quartz-rich sandstone to clay-rich mudstone), by
changes in the pattern of layering, or from evidence that periods of non-deposition or erosion
have occurred. Beds with similar features may be grouped together to form a set. Finally,
stratum (plural: strata) is a general term used to describe any layer of rock. It does not imply
any particular thickness, so it can be used to describe layers of rock ranging from a few
millimetres to tens or even hundreds of metres in thickness.
105
METAMORPHIC PROCESSES AND ROCKS
Heating in the Earth's crust does not always result in melting. As rocks become warmer, their
constituent minerals become unstable and recrystallise to form different minerals that are
stable at higher temperatures. Deeper in the Earth's crust or upper mantle, the mass of the
overlying rocks results in higher pressures imposed on the rocks and new minerals form that
are also more stable at these pressures. Also the shapes and alignments of minerals change in
response to increasing pressures as rocks reorganise themselves to take up smaller volumes or
different shapes. The changes in the mineralogy and in the shape of minerals that make up a
rock when subjected to changes in temperature and/or pressure are known as metamorphism.
CLASSIFICATION OF METAMORPHIC ROCKS
Metamorphic rocks can be classified into two main groups – foliated and unfoliated rocks. A
foliated metamorphic rock is one which made up of layers due to the effects of pressure,
whereas an unfoliated rock is one which is not made of layers. The table below shows this
classification.
Rock Name
Colour
Spotted rock
Grey with
white spots
Grey
Unfoliated
Hornfels
Metaquartzite
Marble
Slate
Foliated
Phyllite
Schist
Gneiss
Grain
size
Fine
Medium
Red to sandy
White to
black
Dark grey to
black
Greenish
Pale & shiny
Light & dark
bands
Medium
Medium
Very fine
Finemedium
Medium
Coarse
Main minerals
Original rock
Clay & mica
Andalusite
Quartz, feldspar,
mica
Quartz
Calcite
Mudstone or shale
Clay & muscovite mica
Pyrite
Clay & chlorite mica
Pyrite
Mica, feldspar, garnet
Quartz, feldspar &
hornblende
Mudstone or shale
Mudstone & shale
Sandstone
Limestone
Mudstone or shale
Mudstone, shale
Mudstone, shale,
or granite
The term metamorphic grade describes the amount of metamorphism/change that has taken
place in rock. Generally, the higher grades of metamorphism reflect conditions of higher
temperatures and/or pressures.
TYPES OF METAMORPHISM
Metamorphism can occur in three ways: in contact (or thermal) metamorphism the heat from an
igneous intrusion alters the adjacent country rock; regional metamorphism takes place on a
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large scale in the roots of developing mountain chains; and dynamic (or dislocation)
metamorphism results from localised deformation along zones such as fault planes.
1. Dynamic metamorphism takes place in areas of intense local deformation, such as at fault
zones. In a major fault zone, there is mechanical movement of one rock over another so that
any rocks within the fault zone may be ground down in a process known as cataclasis. The
metamorphic rock formed is known as fault breccia.
More ductile deformation causes a slippage between layers within the rocks, and between
planes of atoms within minerals. Basically the minerals become elongated due to stretching.
Recrystallisation under these conditions results in a fine-grained foliated rock with a streakedout texture indicating the direction of shearing, which is known as a mylonite.
2. Contact Metamorphism (also known as thermal metamorphism) refers simply to changes
taking place in response to heat associated with igneous bodies. It is most obvious around
intrusive rocks, because, unlike their extrusive equivalents, most of their heat is not lost to the
atmosphere but is dissipated into the surrounding country rocks. Naturally, the temperatures
are highest close to the igneous body itself, and there is therefore a very marked increase in
metamorphic grade near the contact. The zone of metamorphic rocks around the intrusion is
termed a metamorphic aureole. Since the emplacement of magma rarely involves significant
deformation, the main characteristic of rocks that have undergone contact metamorphism is
the formation of metamorphic minerals in the absence of mineral alignment (foliation or
banding) forming a hard, compact and splintery rock called hornfels.
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If the original rock was a mudstone the hornfels will be made up of the metamorphic mineral
known as sillimanite. This is an aluminium silicate mineral only formed during metamorphism of
rocks containing aluminium and silica. Sillimanite is stable at high temperatures (500 - 800ºC)
so is commonly found in metamorphosed rocks close to an igneous intrusion. Further away from
the igneous intrusion the rock becomes less splintery and the sillimanite disappears, and instead
is replaced by a different metamorphic mineral occurring as isolated stubby crystals up to a cm
long. These crystals are the metamorphic mineral andalusite, which is another aluminium
silicate, but which grows at lower temperatures than sillimanite (200 - 700ºC). These relatively
large crystals of andalusite and sillimanite set in a much finer-grained matrix are known as
porphyroblasts.
Frequently, the first sign of
contact metamorphism in the
metamorphic aureole is not
distinct new individual minerals,
but a kind of "spotting" caused
by the growth of clusters of
new metamorphic minerals. Such
rocks are just called spotted
rock or spotted hornfels. As
you can see from the diagram
below, in a metamorphic aureole,
high temperature minerals will
grow nearest the contact
(sillimanite)
and
low
temperature
minerals
(andalusite) are found further away. These metamorphic minerals can be plotted on a map to
produce a map of metamorphic zones. Each zone is known by the name of particular
characteristic mineral present in that zone known as an index mineral, for example, an
andalusite zone or a sillimanite zone. Provided you are looking at rocks of similar composition,
the index mineral is a good indication of the metamorphic grade. Obviously the mineralogical
changes that occur during contact metamorphism depend on the composition of the original
parent rock. So far we have been looking at the changes on mudstones and shales in the
examples above (i.e. turning into hornfels and various types of spotted rocks). Usually during
metamorphism no new elements or chemical compounds are added to the rock. So the mineral
content of the new metamorphic rock is controlled by the chemical composition of the parent
rock. For example, limestone is made up of calcite and cannot metamorphose into anything else
but marble (which is also made up of calcite). In this instance when limestone is heated CO 2 is
temporarily released, but is confined within the rock and recrystallises around existing grains
of calcite to tightly fuse them together, as shown in the diagram over the page.
When sandstones are heated the edges of the quartz grains melt and recrystallise to fuse the
grains together. This gives a granoblastic (non-foliated) mosaic of interlocking grains, and
forming the rock metaquartzite.
108
3. Regional Metamorphism is a very much larger-scale phenomenon than contact
metamorphism. Regional metamorphism is usually associated with continental collision or
subduction, and therefore with deformation. At the highest grades, partial melting of rocks
occurs and ultimately granites are formed. The patterns of changing metamorphic grade differ
between contact and regional metamorphism. Firstly, in contact metamorphism, the zones tend
to be concentric around a particular intrusion due to successively higher temperatures.
Secondly, in areas of regional metamorphism, the zones are often linear due to the fact that
they represent deeply eroded mountain belts, perhaps formed along some ancient destructive
plate boundary or collision zone. Thirdly, and perhaps most important distinction between
contact and regional metamorphism is that the sources of heat are different, which has
important implications for the rate at which metamorphism occurs.
During contact metamorphism, the heat source is a cooling igneous body. This is an effective
heat source only for as long as the magma is crystallising. Even for a large granitic body, this
will be less than 100 000 years. For a small dyke, heat may be exhausted after 100 years or
less. During regional metamorphism, the heat source comes largely from radiogenic heat from
the decay of heat-producing elements. This is a much slower process. For example, garnets
from metamorphosed sedimentary rock in the Himalayas have grown during regional
metamorphism following continental collision. Highly precise dating of these garnets has shown
that although the core of a crystal formed 30 Ma ago, the outer parts of the garnet are only
25 Ma old. This result indicates that during regional metamorphism rocks can be heated for 5
million years at least! Our present understanding of the metamorphism at continental collision
zones is possible only because of the early work of a few 19th Century geologists. For example
in 1893, George Barrow made a classic study of the regionally metamorphosed rocks of the
Scottish Highlands where a sequence of country rocks had been metamorphosed during a period
109
of subduction and collision about 500 Ma ago (Caledonian Orogenesis). Barrow mapped the rocks
in detail and was the first person to recognise metamorphic zones and to use index minerals to
identify successive grades of metamorphism from the appearance of distinctive new minerals in
a traverse across rocks of successively higher grade. Each new mineral represents a further
step towards higher temperatures and/or pressures, and so becomes an index mineral indicative
of metamorphic grade. He recognised in the metamorphic rocks of the south-eastern Highlands
six metamorphic zones and the index minerals he based them on were (in order of increasing
metamorphic grade) chlorite, biotite, garnet, staurolite, kyanite and sillimanite.
The points of first appearance of an index mineral can be mapped and linked by a line. Provided
such minerals are in rocks of similar composition, this line marks the position of rocks with
"equal metamorphic grade" and so is known as an isograd. Conventionally, an isograd is named
after the index mineral on the side of higher metamorphic grade.
METAMORPHIC TEXTURES
Simplified map of Barrow's original metamorphic zones in the
rocks of the south-east Highlands of Scotland.
110
1. Mineral Alignment
When a mudstone is deformed under pressure it will shorten and thicken. How exactly does this
happen? Mudstones are composed of many tiny grains of platy minerals such as clay, and minute
amounts of quartz. Each of these individual grains deform by rotation and by dissolution and
regrowth. In shales these clay minerals tend to be mostly randomly orientated, so as the rock
shortens
under
pressure, almost all of
these clay grains will
physically
rotate.
Exceptions will be grains
already orientated with
their long axis parallel
to the shortening. As
shortening of the rock
progresses, platy grains
like clay will rotate first
into a crude alignment
and onwards into almost
perfect alignment. This
mineral alignment is
known as slaty cleavage in slates or foliation in schists (obviously if a rock does not consist of
platy clay minerals then there will be no mineral alignment, just recrystallisation of existing
minerals and possibly growth of new ones).
2. Recrystallisation of Existing Minerals
At the same time as the physical rotation of the clay minerals, increasing temperatures and
pressures during deformation cause minerals to dissolve and recrystallise by adding or
subtracting from the ends. Minerals at right angles to the pressure dissolve at their ends by a
process known as pressure solution. Conversely, minerals parallel to the pressure grow at their
ends and therefore grow at the expense of those at right angles to the pressure direction. The
effects of pressure solution on small quartz grains, which are relatively equidimensional, is
slightly different to that of clay minerals. Quartz grains tend to dissolve away at points where
they are in contact with other minerals. This process removes material from the quartz grain
into solution, reducing the size of the grain (which also allows nearby clay minerals to rotate
still further), which is then precipitated from solution onto the low-stress areas of existing
quartz grains. The diagram opposite summarises the effects of physical rotation and pressure
solution of minerals under pressure.
111
Stages in the development of mineral alignment and recrystallisation
3. Nucleation and Growth of New Minerals (Metamorphic Reactions)
Rocks are clearly solid objects and do not appear to react. However, minerals in a rock will
react if temperatures and pressures are changed sufficiently. So, metamorphic reactions occur
due to changes in temperature and/or pressure. These reactions typically result in the
formation of new metamorphic minerals such as garnet, cordierite, andalusite, sillimanite or
kyanite. Most metamorphic reactions take place as the metamorphic grade increases (i.e.
temperature and/or pressure increases), and are called prograde reactions. In theory, most
chemical reactions are reversible. If the temperature is increased, the reaction goes in one
direction, and if the temperature is reduced, the reaction goes back in the other direction.
For example look at the following metamorphic reaction;
muscovite + quartz = feldspar + sillimanite + water vapour
112
An increase in temperature of more than 550ºC will cause the reactants (muscovite and quartz)
to react with each other to form the mineral products feldspar and sillimanite, and a gas
product water vapour. This is a prograde reaction because there was an increase in
temperature. If the feldspar, sillimanite and water vapour were then taken back below 550ºC
again, they would react together and revert to muscovite and quartz.
However, in reality most metamorphic reactions do not return to their original starting point
for a number of reasons. One of the products of the prograde reaction (increasing
metamorphic grade) are often gases (such as water vapour or carbon dioxide) which are mobile
and tend to move away from the rocks in which they were generated. So, if feldspar and
sillimanite are unaccompanied by water vapour below 550ºC, they obviously cannot revert to
muscovite and quartz because they are unavailable for the return reaction. The rates at which
reactions take place at lower temperatures is very slow indeed. So, if a group of minerals (e.g.
feldspar and sillimanite) is cooled rapidly, it may not have enough time to react and form lowtemperature minerals (e.g. muscovite and quartz).
If conditions are such that reactions do take place during cooling (decreasing metamorphic
grade), they are called retrograde reactions; these occur as rocks attempt to readjust to the
lower grades of metamorphism. Rates of reaction during metamorphism tend to be slow, but
they increase exponentially with increasing temperature and are also speeded up by the
presence of water. Water acts as a catalyst and is an important method of speeding up
metamorphic reactions. In completely dry rocks, chemical changes can only occur by the slow
process of ionic diffusion (the passage of ions through a solid crystal lattice). When water is
present, it tends to spread itself along grain boundaries as an intergranular film, providing a
network of chemical "arteries" in which ions may move rapidly in solution and so speed up the
metamorphic reactions.
Since many prograde metamorphic reactions form H2O and CO2, both tend to be lost from the
rock. Thus by the time the highest grades of metamorphism are reached, almost all the fluid
will have left the system, and recrystallisation and the growth of new minerals will have sealed
off all the intergranular spaces, making the rock almost impermeable. Not only are H2O and CO2
not present to take part in reverse reactions, they are also unavailable as catalysts. Thus, such
"dry" rocks will then survive cooling more or less unaltered, because the rates of reaction are
limited by the very slow rates of ionic diffusion. So high-grade mineral groupings become
"frozen" in and preserved at low temperatures.
We know that metamorphic rocks are formed at high temperatures and pressures from
experiments designed to reproduce naturally occurring mineral groups under conditions of
known pressure and temperature. The results of such experiments are most simply illustrated
using a phase diagram. A single substance can exist as a gas, liquid or solid. These different
states are referred to as phases of matter. A change of temperature or pressure may result in
a phase transformation, e.g. from a liquid to a solid. At any particular temperature or pressure
two phases may coexist, but in general, only one phase will be found at a particular temperature
and pressure.
113
114
The ranges of pressures and temperatures over which a particular phase is stable (i.e. its
stability field) can be represented using a phase diagram. A phase diagram is a plot of
pressure against temperature (P-T) showing the lines bounding the stability fields of each
phase. Below is an example of a P-T phase diagram for the metamorphic minerals of andalusite,
kyanite and sillimanite. These minerals are polymorphs; they have the same chemical formula
(Al2SiO5) but different crystal structures and so different physical properties such as crystal
shape.
High temperature minerals such as kyanite and sillimanite can be found in rocks at the Earth's
surface today (e.g. hornfels) but they cannot be in equilibrium. In other words, they do not lie
within their stability field when plotted on a phase diagram. Minerals that exist outside their
stability fields are described as being metastable. This means that while they are not in a state
of change, they can be made to change by some sort of impetus. This impetus could be further
metamorphism (rock is reheated and/or squeezed under a new set of P-T conditions) or from
weathering (when the rock is attacked by atmospheric agencies and its minerals are
transformed into those that are in equilibrium at the temperatures and pressures present at
the surface of the earth - like clay minerals in sediments).
METAMORPHISM AND PLATE TECTONICS
Plate tectonics provides a framework for understanding metamorphic rocks. Different types of
metamorphism are likely to occur in different tectonic settings. Destructive plate margins
contain a zone of high heat flow where magmas are generated adjacent to one where the
temperatures have been reduced by subduction. This results in a parallel suite of metamorphic
rocks, one containing relatively high-temperature and low-temperature mineral groupings and
one of relatively high pressures and low temperatures. This is called a paired metamorphic belt
and is characteristic of destructive plate margins. At conservative plate margins extensive
shearing along the plate boundary produces cataclastic deformation textures at shallow depths
115
and Mylonite deformation textures at deep levels in the crust.
THE FOSSIL RECORD
The study of fossils is known as palaeontology. Fossils are the remains or traces of organisms
preserved in rocks. Most fossils are found in sedimentary rocks though a few are found in
extrusive igneous rocks and in weakly metamorphosed rocks.
A fossil is naturally preserved evidence of ancient life. Most palaeontologists (people who
study fossils) consider any evidence of life over 10,000 years old to be a fossil. Most fossils
are millions of years old. Evidence of ancient life can take a number of different forms. A fossil
can be part of the body of an ancient organism such as the bones, shells or leaves. It does not
matter whether or not the parts of the body have been altered in chemical composition and
physical structure. These are known as body fossils. Others are the signs of an organism's
activities, such as the tracks, trails, burrows and borings. These are known as trace fossils.
Trace fossils are often the only evidence we have of extinct organisms whose bodies lacked any
hard parts (like many types of worms). Even if the organism that made a particular trace fossil
had hard parts, the culprit is very rarely found at the scene - at the end of the track or trail.
Trace fossils do, however, have one major advantage over body fossils. Unlike body fossils, in
which the body may be transported after death a long way from where the original organism
lived, most trace fossils are direct, in situ evidence of the environment at the time and place
the organism was making its living.
PRESERVATION POTENTIAL OF AN ORGANISM
Although the fossil record as a whole represents a very small proportion of past life, some
types of organism leave a pretty good record, with abundant fossils. An organism's
preservation potential (the chance that it has of getting into the fossil record) varies
according to a range of factors:


its morphology; i.e. whether its body has any durable parts.

how abundant it is; the more abundant the species is, the higher its preservation potential.
Species with a small average size tend to much more abundant than very large species, and
so are more likely to be preserved somewhere in the fossil record.

how geographically widespread it is; a wide geographical distribution increases the chances
of a species being found in the fossil record.
where it lives and the circumstances of its death, and especially whether or not it is buried
in a marine environment where sediments tend to accumulate. The land tends to be a site
of net erosion, and the sediment that starts off in rivers and lakes mostly ends up in the
sea, especially the shallow seas on the edges of continents. This opportunity for long-term
burial is one of the main reasons why animals from shallow marine environments dominate
the fossil record, and why fossils of land-based organisms are scarce.
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
presence of scavengers and bacteria; anoxic environments (lacking in oxygen) such as deep
marine environments will have far fewer scavengers and bacteria and hence increasing the
preservation potential on any organism deposited in this environment.

whether that sediment subsequently becomes part of the rock record and not destroyed
by tectonic activity and metamorphism.
MODES OF FOSSIL PRESERVATION?
Hard biological materials such as bones, shells and wood often contain tiny pores. When hard
parts are lying in sediment, any such pores tend to be filled up with minerals that crystallise
out from the water seeping through the sediment. If the original spaces in a porous material
are impregnated with extra minerals in this way, the material is said to be permineralised. The
growth of new minerals at the expense of any original biological material (such as the cell walls
of bone or wood) is called mineral replacement. The minerals most commonly involved in
permineralisation and replacement include: silica (quartz); calcium carbonate (calcite) and iron
sulphide (iron pyrites). Both the infilling of pores (permineralisation) and the replacement of
biological materials (mineral replacement) may occur in a single fossil. Taken together these
processes are called petrification (literally "to turn to stone"). However, neither of these
processes has to occur for something to be called a fossil; sometimes the fossil can still be
composed of the original, barely altered shell or bone.
The diagram below illustrates some terms that are useful when describing some types of
preservation commonly found in fossils. The surface of the sedimentary rock lying against the
inside or outside of the shell is called a mould, i.e. the impression of the shell's surfaces on the
adjacent rock. Usually both internal and external moulds are formed, on the inside and outside
of the shell, respectively. If at some stage the shell is completely dissolved away, a space is
then left between the internal mould and the external mould, forming a crude cast of the shell,
which lacks details of the original structure. In general, casts are rarer than moulds.
So, a single fossil may show several modes of preservation; for example most fossil shells show
some degree of replacement or permineralisation, and the shells are normally associated with
their internal and external moulds in the adjacent rock.
CLASSIFICATION OF FOSSILS
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Taxonomy is the study of classifying organisms. There are three major groupings, or kingdoms,
into which life can be classified; animals, plants and fungi. Within the kingdoms, organisms
naturally fall into broad groups, each of which is characterised by a distinct body plan which is
fundamentally different from that of other groups. Such a group is termed a phylum ("fileum"). Humans belong to the Phylum Chordata (vertebrates). Within each phylum there are
subgroups called classes. Each class has its own characteristic modification of the phylum body
plan. There are for example, 5 vertebrate classes living today: fish, amphibians, reptiles, birds
and mammals each recognised by their distinctive morphology. Within classes a series of
smaller subgroups can be recognised, these groups are order, family, genus ("jen-us"), and
finally species. The groupings, generally known as taxa, thus form a series known as a taxonomic
hierarchy. Below is the taxonomic hierarchy for the domestic cat!
Fossils organisms are placed in the same taxonomic groups as living organisms, but since their
breeding behaviour can hardly be tested, they have to be classified in terms of their
appearance and structure.
Organisms can also be classified by the type of environment they live in. The main types of
environments in which organisms live are aquatic (in water) or terrestrial (on land). In the
marine (sea), environments are classified in terms of water depth into littoral (between high
and low tide), neritic (down to about 200m), bathyal (between 200 and 400m), abyssal
(between 4000 and 5000m), and hadal (below 5000m). Freshwater environments include lakes
and ponds (lacustrine), rivers (fluvial) and marshes and swamps (paludal). Terrestrial
environments include those of deserts, forests, grasslands and tundra. They tend to be more
variable and to have more extreme conditions than aquatic environments.
Organisms are often grouped according to their lifestyle. For example marine organisms can be
grouped into two categories; pelagic or benthonic. Pelagic organisms live within the body of
water above the sea floor, whereas benthonic organisms live on the sea floor. Among benthonic
animals, those which live within the sediment are infaunal (i.e. burrowing animals) while those
which live on the surface of the sediment or rocks are epifaunal. Among the pelagic group,
organisms which live floating or suspended in water are described as planktonic (plant plankton
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is phytoplankton and animal plankton is zooplankton), whereas organisms which swim rather than
float are known as nektonic.
USING FOSSILS TO RECONSTRUCT PAST ENVIRONMENTS AND INTERPRET MODES OF LIFE
When reconstructing past environments, the geologist relies to a great extent on comparing
them with what he/she believes are similar environments today. This principle of comparison is
known as Uniformitarianism. It is a basic concept in geology and it dates back to the time of
the Scottish geologist James Hutton (1726-97) who was one of the first to realise that what
was happening on the earth‘s surface today, such as erosion and deposition, must have been
going on in the past. He came to this conclusion after studying sedimentary rocks near his home
in Edinburgh.
The principle of Uniformitarianism has been encapsulated in the phrase. ―The present is the key
to the past‖. But to what extent is it true that processes on the Earth‘s surface today are
typical of what has happened in the past? Generally, it seems as if there is close agreement and
the study of modern environments provides a good basis for the reconstruction of the past,
although the level of comparability between modern and ancient decreases with time.
With more recent geological periods it is relatively easy to find fossils that are closely related
to organisms living today. It is possible to state with some certainty that a fossil of Ostrea
bellovacina found in the Tertiary would have had a very similar mode of life to a modern Ostrea
(oyster). However, what happens if the fossil is from an extinct group? By definition there are
no living members with which comparison can be made. Here, there is plenty of scope for the
geological detective.
The trilobites are an extinct group but we have quite an accurate idea of their mode of life.
The type of rocks in which they are found and the other fossils associated with them provide
valuable clues. Trilobites are found with brachiopods and corals, which we know by comparison
with modern groups live in the sea; they are not found with land dwelling organisms. We can
therefore say trilobites were marine creatures.
Further evidence comes from studies of the morphology of trilobites. Trilobites have hard
exoskeletons on the upper sides of their bodies. We know from exceptionally well preserved
ones (e.g. in the Burgess Shales in Canada) that they had legs on their undersides. Where eyes
are present, they are usually on the tops of their heads and would have given the trilobite a
clear field of vision into the surrounding water. Taken together, these observations indicate a
mode of life on the sea bed, where the animal probably crawled around in search of food. Some
trilobites lack eyes and so it seems that they must have inhabited areas where eyes were not
needed. Perhaps they lived in dark burrows or in deep water through which the light did not
penetrate.
Another line of approach with extinct groups is to compare them with their closest living
relatives. A very accurate idea of the mode of life of the extinct cephalopod groups, the
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ceratites, goniatites and ammonites, can be obtained by comparing them with a living
cephalopod group, the nautiloids.
With other groups comparisons are not quite so easy. The exact mode of life of dendroid
graptolites is still not quite clear as they do not have any close living relatives and their
morphologies do not readily suggest a particular mode of life.
Even when living relatives of fossil organisms are known, we have to be careful not to misapply
the principle of Uniformitarianism. For example, stromatolites are some of the earliest fossils
found in the Precambrian. They have been recorded from rocks such as the 1,900 million year
old Gunflint Chert on the shores of Lake Superior, or the even older Bulawayo Limestone of
Zimbabwe which may be more than 2,700 million years old. It is believed that they have been
formed by primitive organisms related to the algae known as cyanophytes (blue-green algae).
Further examples of stromatolites are found throughout the fossil record and the layered
structure that they all show closely resembles modern forms.
Modern stromatolites are found in rather restricted intertidal areas in the Bahamas and
Western Australia. They occur as mounds of around a metre in height built up of alternating
layers of sand impregnating the algal mat and carbonates precipitated by algae. The algae are
mainly the very primitive blue-green algae together with some green algae.
In the parts of Shark Bay, Western Australia where the stromatolite mounds are found
growing today, a sandbar has largely cut off direct access to the sea. The hot climate causes
evaporation of the water behind the sandbar making it very saline. This does not hamper the
growth of the stromatolites but molluscs and other creatures which would normally feed on the
blue-green algae cannot survive. Stromatolites are only found today in areas such as Shark Bay
where they can grow unchecked because predators are kept out by hostile conditions.
A strict application of the theory of Uniformitarianism would have it that stromatolites would
always have been found in these rather unpleasant, hot, saline conditions. Back in the
Precambrian, however, when the stromatolites were the most advanced life form, they would
have had a much wider distribution, as they could have flourished uncropped in a variety of
different environments.
This case illustrates that there has to be a certain amount of caution in relating modern
environments and conditions to the past. The changing variety and abundance of life forms
causes different pressures. In the past, stromatolites survived unchecked in many different
environments, but today competition and predators have forced them into the rather
unfavourable habitats in which we now find them.
BRACHIOPODS
Brachiopods are shelled, benthonic marine organisms. The shell consists of two valves usually of
different size which are symmetrical about a median plane running from one end of the shell to
the other. Brachiopods are found mostly in sedimentary rocks deposited in shallow water.
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About 70 genera exist today mostly in warm, shallow eastern seas such as around Japan,
Australia and New Zealand, but some live in colder water to depths of about 5000m. Adult
shells are usually 2 – 7cm in length.
The phylum Brachiopoda is divided into two classes; the Articulata and the Inarticulata. The
Articulata have calcareous shells (i.e. made of calcite) hinged by teeth and sockets. The
Inarticulata have horn-like shells made from a protein called chitin (similar to beetle wings).
The valves are not hinged by teeth and sockets.
In articulate brachiopods the valves are not equal in size. The larger valve is called the pedicle
valve and the smaller valve is the brachial valve. The shell is hinged along a hinge line at its
posterior end. The posterior end of each valve is pointed to form a projecting region called an
umbo (plural umbones). The valves may be marked by radial ribs and by concentric growth lines.
The umbo of the pedicle valve usually has a round opening called the pedicle opening. From this
pedicle opening comes a stalk called the pedicle which attaches the brachiopod to the sea floor.
Some forms have no pedicles and they have their pedicle vales cemented to the sea floor or
they may be attached by means of spines. The hinge line of the pedicle valve has two
projections called teeth which fit into sockets on the hinge line of the brachial valve.
The brachial valve holds the lophophore. The lophophore is used for gas exchange and for
feeding. Also inside the shell are three types of muscle. Adjustor muscles from the pedicle
valve to the pedicle turn the shell into or out of water currents to help feeding. The other two
muscle sets run between the valves. The adductor muscles close the valves and the diductor
muscles open the valves. The points of attachment of the muscles to the insides of the valves
leaves marks called muscle scars.
Inarticulate brachiopods are often chitinous but some are calcareous. The lack of hinge
structures means that the valves cannot be opened in the same way as those of the articulate
brachiopods. Another difference from articulate brachiopods is that the lophophore has no
supporting structures. Where a pedicle is present it comes out between the two valves though a
few have a pedicle opening.
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All brachiopods are benthonic marine animals, and most are unable to tolerate the influence of
fresh water. The brachiopod Lingula is an exception in that it can withstand brackish
(freshwater mixed with salt water) coastal waters (e.g. estuaries). A major consideration in a
brachiopod‘s mode of life is its type of attachment. Many articulate brachiopods live attached
by a pedicle. This is a short cord of connective tissue clothed in a tough cuticle. It usually leads
to a rock or to another shell, but some are known which are root-like and are able to anchor the
animal in soft sediment. The brachiopod may be able to swing round by using its adjustor
muscles so as to face different ways according to the current, but in some species the pedicle
becomes atrophied (wasted away) and can only function as a tether. In fossil forms it probably
withered away entirely, so that the mature animal was free-lying.
Free-lying forms were generally of a shape which ensured their stability. Rather in the way a
self-righting toy always comes to rest in the correct orientation, their shape contrived to be
unstable in all positions except the one desired. A few free-lying forms may have been mobile.
The design of the shell in these forms suggests that it could jet-propel itself backwards by
snapping its shells closed and expelling the water from its mantle cavity.
Both articulate and inarticulate brachiopods are known to have taken to cementation as a mode
of fixing themselves. In this life style there is a strong chemical bond between the fixed valve
and the rock. Cemented brachiopods were unable to move once they had the settled, nor could
they change their orientation. It is therefore likely that their larvae were programmed to
select sites where there was consistent water flow, and to begin growth in precisely the
correct alignment relative to it. Cemented brachiopods tend to be found in habitats beneath
things, such as beneath boulders, coral colonies and in reef cavities.
No brachiopod ever burrowed very deeply into sediment because they needed to maintain
contact with the water and were not able to form tubes or siphons. Lingula, perhaps, comes
closest to achieving this, in the way that its setae (sensory organs/spines) are able to form
short siphons. There is a difference, however, between active burrowers like Lingula and the
forms which merely allowed themselves to become half-buried. Such types were typically
concave in shape on the top valve and concave in shape on the bottom valve. They lived with
their narrow, slit-like apertures projecting up through the sediment. They may have had spines
to help with anchorage.
All filter feeders need to ensure as little water as possible is recycled between the inhalant
and exhalent currents. In brachiopods this has led to a very common feature, namely the
folding of the commissure (shell opening) into usually a median upfold and a pair of flanking
downfolds. This pattern of folding separates the two inhalant streams from the exhaust
current rising medially. It is a feature which may be most significant to brachiopods living in
very quiet water conditions where recirculation would be an even greater danger.
Another common feature of the commissure, seen in many brachiopods, is the zig-zag. This is
often combined with the larger scale of folding just described and is principally a protective
measure to prevent particles which might damage the lophophore from entering the shell. By
zig-zagging the commissure, its length is much increased, and so a relatively large amount of
water can flow even with quite a small angle of gape. This has the additional advantage of
keeping the mantle edges close together ensuring a closer monitoring of the inflowing streams.
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BIVALVES
Molluscs are soft-bodied invertebrates which often have calcareous shells. The phylum
Mollusca is divided into various classes of which one is the Bivalvia (mussels, cockles, scallops
and razor shells). Bivalves are marine and freshwater organisms whose shells consist of a pair
of calcareous valves. The two valves may be ornamented by ribs radiating from the umbones or
by concentric growth lines. The two valves are usually similar or equivalve; if dissimilar, they
are inequivalve. The valves are joined in the region of the umbones. Each valve has a flattened
area called a hinge plate which bears an internal or external ligament, protruding teeth and
sunken sockets. The ligament acts as a spring which tends to open the valves. The teeth fit into
sockets in the opposite hinge plate. Away from the hinge the valves meet along the commissure.
In some burrowing bivalves the shell does not close completely and there may be an opening or
gap. The valves are held together by adductor muscles. Sometimes, as in the scallop, there may
be one central muscle but, more often, there are two muscles. The points of attachment of the
muscles to the insides of the valves leaves marks called muscle scars. Between the muscle
scars and close to the margin of the valve is the pallial line. The pallial line marks the edge of
the inner shell layer secreted by the organism living in the shell. In some bivalves the pallial line
is indented into the shell to form the pallial sinus. Such a sinus is found in most burrowing
bivalves since it marks the edge of the pocket into which the large siphons or water-breathing
tubes can be withdrawn.
Four basic modes of life are exhibited by bivalves: burrowing and boring, byssal tethering,
cementation, and free-lying.
Burrowing may be very shallow (e.g. the cockle), or deep (e.g. Mya and the razor clams). The
cockle has a thick shell for its small size. It is equivalve, has a rounded outline, with radial ribs.
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These ribs are created by a folded commissure (shell opening) whose crenulations interlock
when the shell is closed. The shell has no permanent gape and internally there are two equal
sized muscle scars and a pallial line
which has no sinus. The thickness of
the shell and the strengthening
provided by the ribs give the cockle
a high level of protection. It can
close completely on the soft parts
and its meshed commissure resists
the battering of waves and the
attacks of predators by preventing
easy dislocation of the valves. The
cockle lives just below the sediments
surface in high energy tidal
environments. It therefore needs
only short siphons – hence the lack
of a pallial sinus – and is capable of reburying itself rapidly if uncovered. The outline of the
cockle shell is rounded. Generally, the deeper a bivalve buries itself, the more elongated its
shell becomes. The shell of Mya is a good example of this. The protection afforded by the
burrow (which can be 50cm deep) means that the shell can be much thinner than that of the
cockle. Its valves do not mesh together, and indeed they do not even close properly forming a
gape. The siphons of the deep-burrowing bivalves were often very bulky. Some, like Mya, are so
large they need to be permanently extended and are far too large to be fully withdrawn into
the protection of the shell. Therefore, bivalves like Mya have shells which gape permanently at
the posterior end, even when fully closed. A similar gape may also exist at the anterior end of
deep burrowing bivalves to accommodate the muscular foot. The long siphons have created a
deep pallial sinus, which is where the siphons emerge from the body of the organism.
Shell elongation is carried out to extremes in the razor clams, whose valves form an open-ended
tube. Although they are deep burrowers, their siphons are fairly short, so that the animal must
rise in its long burrow to feed. It can, however, retreat downwards very quickly if disturbed.
Once again the dentition is weak and the shell thin. Bivalves burrow by firstly opening their
valves. This forces them against the burrow walls, locking the shell in position and providing a
firm anchor against which the foot can push as it forces downwards. When fully extended, the
foot swells at its end; the valves close and water in the mantle cavity is expelled, liquefying the
sediment around the shell. The foot then contracts; pulling the shell down through the liquefied
sand. The cycle is repeated until the required depth is attained.
Bivalves which bore into rock or wood carry the reduction of the shell‘s protective role even
further. Rock-borers (e.g. Pholas) have permanently gaping shells and rasp-like ornamentation
which enhances the shell‘s abrasive properties. The use of the shell as a borer requires
powerful and unconventional musculature. Extra attachment is provided by the projecting
myophore and, since a ligament would not provide enough opening force for the movement to
have any abrasive power, one of the adductor muscles has come to attach itself to the exterior
of the shell at one end, so becoming a diductor muscle. Its emplacement is protected by the
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secretion of accessory shell in the umbonal region. Another rock-borer, Lithophaga (literally
―rock-eater‖), in addition to mechanical abrasion, can produce acidic mucus to soften the rock
and ease penetration.
Byssal tethering is achieved by use of a byssus (a group of tough threads made of conchiolin),
which is secreted by a gland near the base of the foot. Each thread begins as a liquid extrusion
of this gland, but soon becomes tough and fibrous in contact with water. The byssus may serve
to attach the bivalve to a rock or to seaweed or may anchor the shell in sediment, rather like a
root system. In most byssate forms the shell opening (commissure) is held vertically. In the
mussel Mytilus, the shell‘s anterior portion is much reduced, and the posterior projects upwards
from the rock, so allowing for very close packing of individuals in clusters. This is a protective
device, making mechanical dislodgement by predators (such as seagulls) more difficult. The
byssus emerges from a shallow notch known as the byssal notch. These bivalves have a
streamlined air-tight shell (no gape) to tolerate the pressures induced by wave action.
The free-lying scallop Pecten can swim to escape danger by clapping its valves together. Water
can be expelled at various points around the edge. This means that considerable
manoeuvrability is possible. Usually water is expelled in the hinge areas, so that the scallop
appears to swim by taking bites out of the water ahead of it. A number of devices are found to
convey extreme stability in one position in free-lying bivalves. Weight is an obvious stabilizing
influence, and the enlargement of the lower valve coupled with thickening of the shell is
frequently seen in free-lying species. Others may extend one or both ends of the hinge line into
long fingers. These are thought to act, like outriggers of a boat, to prevent overturning.
Gryphaea has an enlarged and incurved lower valve. This gives the animal a horn-like shape. This
is a shape seen in many soft-bottom recliners. The reason is its efficiency as a self-righting
form. It is a shape which lies naturally with the opening pointing upwards.
An example of a cementing bivalve is the oyster. Oysters are among the most successful
bivalves. They are good indicators of shallow seas. Like free-lying forms, the shells are large,
but for not entirely similar reasons. For, although stability is no handicap, rapid growth is
important when competition for space is high. Also, a thick shell helps ensure that parasitic
shell-borers are not able to pierce the protective armour. The larva of the oyster settles on a
rock and is fixed; in the first instance, by a sticky secretion of the byssal gland. Thereafter, it
is the mantle edge of the lower valve which cements the shell. In cemented forms the need to
conform to the rock and the close pressing of other individuals makes shell form very variable.
Competition for space is usually a prominent factor of life for many organisms. The rudists
were very successful bivalves, whose fixed valves grew upwards like organ pipes, often to
height of half a metre.
TRILOBITES
Trilobites are hard-shelled, segmented creatures that existed over 300 million years ago in the
Earth's ancient seas. They went extinct before dinosaurs even existed, and are one of the key
signature creatures of the Palaeozoic Era, the first era to exhibit a proliferation of the
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complex life-forms that established the foundation of life as it is today. Trilobites first
evolved in the Cambrian and became common in the Ordovician and Silurian Periods. By the end
of the Permian they had become extinct. Therefore they have no modern equivalents and an
understanding of their soft parts has to be based on modern day arthropods that show some
similarity i.e. crustaceans.
Trilobites were marine animals and therefore were found in water based sediments. A few were
found in delta sediments but none in fresh water. Mostly they were benthonic and crawled
around on the sea bed. Some may have spread their weight on soft mud by using spines; others
may have ploughed through the soft sediment on the sea floor. Some may have even buried
themselves. A few trilobites were pelagic and were streamlined enough to be able to swim close
to the sea floor.
They usually lived on the continental shelf where there was abundant life as here the water was
shallow enough to have some light to stimulate the food chain. Trilobites are commonly found in
shales, mudstones and limestones; therefore they are most common in low energy environments
in which these rocks occur. Shales and mudstones indicate that some species could live in
muddy conditions. Limestone indicates some species preferred clear conditions. Some species
are found in both. Delta trilobites lived in a high energy environment and can be distinguished
by their heavy shell which gave protection from the waves. Marine trilobites lived in a low
energy environment away from the coast so the exoskeleton was thinner except those at great
depths (because of the pressure).
Trilobites tend to be fairly small being 5 - 8 cm long on average although extremes do occur
from 5 mm to 70 cm. Because they had a rigid exoskeleton growth caused problems. How did
they grow? They moulted i.e. they shed their exoskeletons for a larger by a process called
"ecdysis". They are segmented animals and have a chitinous exoskeleton. They have a bilateral
symmetry i.e. either side is symmetrical and consist of jointed limbs which have an identical
pair on the other side of the body. The body can be divided into segments:
1. Laterally:
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2. Transversely into three regions:
A central or axial segment bounded by two
lateral segments. This is where they get their
name from "three lobes".
A "head" area known as the cephalon, "body"
with hinged segments known as the thorax and
a "tail" with fused segments known as the
pygidium.
The cephalon is the head shield which consists of a central region glabella or axial region. The
glabella is usually convex but does vary in size and shape in different species. The glabella is
separated on either side by AXIAL FURROWS with cheeks at either side. Sometimes
transverse furrows exist which are called GLABBELA FURROWS. The OCCIPITAL FURROW
occurs at the base of the cephalon just before the thorax. In some species the glabella
reaches right to the front of the cephalon truncating the cheeks on either side. In some
species the glabella does not reach as far and the CHEEKS continue round in front of the
glabella. Normally the cheeks are divided into 2 areas by the FACIAL SUTURE the line along
which the skeleton is cracked to allow ecdysis: Eyes are usually found on the free cheek,
situated close up to the facial suture. Most trilobites had eyes although some had none and
were blind. The eyes are usually compound (like flies) and were made of calcite! Some species
had eyes on stalks. Eyes in early trilobites were elongate and crescent shaped. Eyes in later
trilobites were more compact and kidney shaped. The genal angle occurs at the corners of the
cephalon closest to the thorax, usually best shown by spines which extend backwards from the
cephalon: genal spines. Plates are thought to protect the soft parts.
The thorax is made up of the exoskeleton. The exoskeleton has a curved, flexible shield which
is usually convex. It has a central axis separated from the 2 lateral regions by the AXIAL
FURROW. Each segment of the lateral area has a PLEURA(E). The thorax contains segments
(which are not fused together like the pygidium) which are jointed and able to move
independently. This flexibility allows some trilobites to roll up to give protection to the softer
under part. The number of these segments varies e.g. 2 - 40. However, in any Genera the
number is constant e.g. 13 in Calymene. The segments can vary in size, working down the body
although they would have the same basic features. Trilobite segments are usually wider and
longer towards the cephalon and narrower and shorter towards the pygidium. The walking limb
extended out under the pleura. Each pair of pleura protected a pair of legs and gills.
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The pygidium is a semi-circular or triangular shield at the ―tail‖ of the animal. These have a
number of fused segments which varies from 2 - 30 but cannot move independently. However, in
some genera the furrows between the pleura can be lost giving a smooth pygidium.
The many forms seen in trilobites probably reflect different modes of life, and
palaeontologists have long been tempted to make speculative suggestions as to what features
correspond to which life style. With increasing knowledge it is now becoming possible to support
or disprove many these interpretations. Let us look as the eyes first of all. Trilobite eyes can
show various degrees of development or suppression. Trilobites which were blind may have been
nocturnal in habit, or may have lived in cavities such as occur in reefs. Alternatively, they might
have been burrowers. The absence of eyes is a difficult feature to which to ascribe any
definite single cause. It is best weighed with other features before deciding.
For example, those forms with small eyes combined with a wide axial region and smooth
exoskeleton, whose thoracic segments appear to have fitted very closely, may have been
burrowers. The wide axis possibly reflects the presence of strong muscles, and the smoothness
of the exterior aided passage through sediment. There would have been some need for the
exclusion of grit from the gaps between the segments, hence their close fit.
Forms with very large eyes could not have been benthonic, since this life style would have
buried their enormous compound eyes in the sediment. Such a degree of all-round vision could
only have been of use to a free-swimmer. They may have swum upside down near the surface, so
keeping a sharp look-out for prey or predators. They are often found in deep-water oceanic
deposits.
The wide cephalic fringes seen in some trilobites (e.g. trinucleids) have been subject for
speculation for a long time. The trinucleids brim is made up of two sheets of cuticle joined by
roughly cylindrical tubes which look, on the surface, like pits, but which really act as pillars
separating the two cuticular sheets. Halfway down these cylinders there is a delicate wall which
closes off the tube but for a microscopic central perforation. Evidence from the traces left by
trinucleids resting upon the sediment surface suggests that the animal used its legs to dig a
small hole or depression, and that it then positioned itself in the hole with the cephalon angled
downwards. Since these traces appear to point consistently in one direction when found
together, it is assumed that the animals were aligned towards some prevailing current. The
implication from this is that trinucleids were filter feeders. The deep arch of the cephalic
fringe may have allowed currents to pass under the creature and towards its long gill branches.
Spines on trilobites are common, especially in certain types of trilobites, and it has been
suggested that they acted as baffles to increase water resistance and so help prevent sinking.
In most cases this seems rather unlikely; the spines would probably have had only a marginal
effect upon such a relatively large creature. The spines are probably more likely to be used as a
means of defending itself. Spines on the underside of a trilobite probably acted as supports,
but spines on the backs of trilobites probably had a protective function when the trilobite was
rolled up.
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The following 7 trilobites show how morphology relates to mode of life.
CALYMENE:
Clear shallow water. It crawled on the sea bed and may have
swum just above it. Occurred from the Silurian to Devonian and
it‘s long duration and evolution suggests that it was well
adapted to its environment.
The eyes were kidney shaped (but not as fully developed as in
other species). Eyes on top of cephalon suggesting that it lived
on the sea floor. The fact that the mouth faced backwards
suggests that it was not a predator.
Some could roll up - for protection? (Fexycalymene).
Age: Lower Silurian - Mid Devonian.
DALMANITES:
Well-developed eyes could see all around: the eyes have
evolved. Give better protection. Likely to have walked on the
sea floor.
The long spine on the pygidium was of an uncertain use but
possibly gave support on the soft sediment as did the genal
spine and small spine on the front of the cephalon.
Age: Silurian - Lower Devonian.
DEIPHON:
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Pelagic and may have been a free swimmer. The large
surface area was created by the well-spaced out
plurae with spines on the cephalon, thorax and
pygidium. The glabella was also inflated which could
have aided floatation.
It was prone to swimming predators and the spikiness
and tubercles would give some protection. The eyes
were slightly forward facing suggesting life above the
sea floor.
Age: Silurian.
TRINUCLEUS:
Has a pitted fringe which was of uncertain use but may have
allowed the trilobite to burrow through mud and the pits were
sensory and allowed it to feel its way around. The eyes would be
lost if:
1. The trilobite lived in deep water where there was no light and
therefore no need for eyes or:
2. The trilobite burrowed and there was therefore no light and
so the eyes were lost. The cephalon was also roughly shovel
shaped which may have facilitated movement through mud. The
trilobite would then eat the sediment extracting any organic
matter from it. Age: Ordovician.
AGNOSTUS:
This trilobite was blind and was a very small and simple trilobite.
It was a primitive early form. Had no need to see because it
lived either buried in mud or deep water where there was no
light. Latter most likely because usually occurs within deeper
water fauna. May swim in a partly enrolled form. The legs are
too small and unsuited to walking - partly enrolled. Fed on minute
particles wafted in by antennae.
Age: Upper Cambrian.
PARADOXIDES:
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Early form, eyes elongated and not highly developed (compare
with Dalmanites). Long spines are common in many trilobites.
Possibly used to support a relatively heavy animal on very soft
sediment. Therefore crawled along the bottom.
Age: Mid. Cambrian.
ENCRINURUS:
Eyes are on short stalks which suggest that it lived buried in
sediment. The tubercles (little round lumps) also gave it added
strength and protection.
Age: Mid. Ordovician - Silurian.
DINOSAURS
The term 'dinosaur' originates from the Greek words deino and sauros, meaning 'terrible' and
'lizard'. Sir Richard Owen invented the name 'dinosaur' in 1842, to describe an extinct group
of terrestrial reptiles that lived during the Mesozoic. At that time, he had only three or four
dinosaur skeletons to consider, but many hundreds more have been found since 1842.
Dinosaurs show a range of adaptations to a wide range of environments and conditions, and as a
result they became the most successful land animals of all time. There are many misconceptions
about what the dinosaurs were. Dinosaurs were reptiles (horny or scaly skinned tetrapods that
lay eggs) which lived on land during the Mesozoic Era (250 - 65 million years ago). Dinosaurs
arose in the Middle or Late Triassic, some 235-230 million years ago. The large marine reptiles
and the flying reptiles of the Mesozoic were not dinosaurs. Dinosaurs are found the period of
Earth's history known as the Mesozoic Era, which began 250 million years ago and ended 65
million years ago. The Mesozoic is further subdivided into three periods, the Triassic, Jurassic,
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and Cretaceous. Fossils show that the dinosaurs originated during the Mid to Late Triassic,
perhaps 230-235 million years ago, and they remained the dominant land animals for another
160 million years, until they died out at the end of the Cretaceous.
A typical dinosaur had many more bones than a human does – about 300 in all. The extra bones
are the 40 to 50 in the tail. Otherwise, a dinosaur had a skull, a backbone, ribs, arms and legs
just like any other backboned animal (vertebrate). The vertebrates, as a group, include the
fishes and the tetrapods (literally ―four-footers‖). Tetrapods comprise amphibians (frogs etc.),
reptiles (lizards, snakes, crocodiles and dinosaurs), birds and mammals. All vertebrates have a
skull and backbone, and all tetrapods have arms and legs, and the other necessary bits and
pieces to tie all these parts of the skeleton together. A typical dinosaur skeleton is shown
below.
Dinosaurs are divided into two major groups based on the structure of the hip; they are known
as the bird-hipped (ornithischian) and lizard hipped (saurischian) dinosaurs. The names are
rather misleading as it is from the saurischians that the birds actually evolved; these names
were given before this was fully understood. The pelvic girdle consists of three bones: the ilium
attached to the backbone and the ischium and pubis which may crudely be described as holdingin the rear of the animal. In humans the position of the ischium means that it takes the weight
when we sit, while the pubis is in the front. Together these three bones enclose the hip socket
of the hind leg. In the saurischians the pubis points forward, but in the ornithischians it points
backwards alongside the ischium. This difference can be understood if the dinosaurs‘ feeding
habits and stance when walking are considered. All ornithischians were herbivores and many
were bipedal, walking for at least part of the time erect on the hind legs. As vegetarians they
would have a large gut to allow the food to pass through sufficiently slowly to allow it to be
digested, a process involving symbiotic bacteria. We can imagine an erect ornithischian as
having a ―beer belly‖ and if this had rested on a forward pointing pubis it would have
overbalanced the animal. The backward pointing pubis permitted the gut to hang between the
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legs. The bipedal saurischians (e.g. Tyrannosaurus) were all carnivores so the gut would have
been much smaller as meat is quickly digested.
As recording the shape of an animal, a fossil skeleton can contain clues to the health of its
owners. Only about 1% of diseases leave their mark on the bones, but these include diseases
like arthritis, which has been found in two Iguanodon ankles! It used to be thought that
dinosaurs suffered a lot of osteoarthritis because their back vertebrae often appeared fused.
This has since been revealed as a condition known as diffuse idiopathic skeletal hyperostosis or
DISH. This occurs when the ligaments running next to the vertebrae become calcified and
gradually lock the backbone into a rigid shape, but it is not associated with any disease. Indeed,
it looks as if it might have been an adaptation to help dinosaurs support their weight over their
back legs. Duck-billed hadrosaurs, horned dinosaurs, sauropods and Iguanodon all show this
condition in their back vertebrae, which could have helped them hold their tails high.
Stegosaurus needed a highly flexible tail for defence, so did not suffer from DISH. Horned
dinosaurs, such as Triceratops, show fusion of their neck vertebrae as well. This was probably
to help them support such a large head.
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Dinosaurs ate to live. Mostly, we have to guess what they would have eaten, but as we have seen
already the skeleton is full of clues. For example, tooth shape is a good guide to diet. The teeth
of carnivorous dinosaurs are instantly recognisable. They are curved and pointed in side view,
and rather flattened. In close-up the stabbing teeth of meat eating dinosaurs such as
Allosaurus were fiendish! Most of them had finely serrated front and back edges, maybe two or
three zigzags per millimetre. As the tooth came into use, these serrations wore down gradually.
Eventually they would be blunted, but by then the tooth had had its day and would be shed, and
a new incredibly sharp specimen would move into place. An ever-ready, self-sharpening set of
steak knives! Also with curved, backwards-pointing teeth any prey that struggled in the
dinosaur‘s mouth would end up moving further down its throat. Plant-eating dinosaurs never
have these kinds of teeth (e.g. Diplodocus). Theirs may be rather feeble little pencil-like
structures, or sort of leaf-shaped in side view, with a narrow ridge fore and aft. In some forms
that specialized on especially tough vegetation, the tooth has extra ridges and grooves to give
it strength.
A number of plant-eating dinosaur groups had crests, horns and spines. Certainly these
structures were probably used in fighting and defence, but structures can have more than one
function. These head crests, horns and spines could have also made good signalling devices,
especially to attract a mate. However, it wasn‘t just the plant-eating dinosaurs that had crests
and horns for signalling and fighting. Some flesh-eating dinosaurs did, too. Some of these
predators had crests and knobs on the skull. It would hardly seem necessary to use these to
scare their prey: the sight of the sharp teeth and claws would be quite enough. So these
structures might have been used when males postured or blustered, trying to make their rivals
back down. Some of the knobs and crests contained air spaces, and they might also have acted
as resonating chambers to help create a scary roar.
It is unclear whether the huge sauropods dinosaurs, such as diplodocus, were entirely immune
from predation because of their large size or not. Despite the finds of theropod (flesh-eating
dinosaur) tooth marks on some skeletons (only very few), most palaeontologist accept that, like
elephants, the sauropods were pretty well left alone. Only babies might be snatched. But some
sauropods did have some defensive equipment. The titanosaurs, for example, had armour. The
armour plates were roughly round discs of bone set in the skin in regular close-packed patterns.
It wasn‘t a solid armour like a turtle, but it formed something like chain mail, probably an
effective defence against a passing attack. In a full-scale onslaught, however, a large theropod
could bite through to the vital organs, but would lose a few teeth in the process. Some
herbivorous dinosaurs took this form of defence to the extreme, such as the armoured
dinosaurs. The armour suit was perfected by the Anklyosaurs. The rows of bony plates over the
back and sides were fused into a single armour coat. Some had long spines along the sides; some
had tail clubs. Even their heads were completely encased in extra layers of bone.
Stegosaurus had a row of tall plates down its back. These might have had a function in defence
or in scaring off enemies. But they were almost certainly used in heat exchange as well. The
bony structure of the plates is roughened, showing that it was covered with skin in life. And
there are grooves around the base, where blood vessels ran over the bone. If Stegosaurus was
hot, or if it wanted to look scary, it could pump blood into these vessels and flush its plates red.
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Heat would be radiated. Likewise, on a cold day it could cut down the blood flow to the plates
and conserve heat.
The diagram overleaf shows the variety in size and shape that existed with dinosaurs. The small
plant-eating dinosaurs were mostly bipedal. They were fast movers, with long slender legs and
stiff balancing tails, because at their size, form 1 to 5 metres long, they couldn‘t afford to be
around if a slavering Allosaurus came on the scene! Higher plants were consumed by the
sauropods, the giants of all time. At 50 tonnes or more they were huge (e.g. Apatosaurus,
Brachiosaurus and Diplodocus). Palaeontologists have often wondered why sauropods didn‘t seem
to have any way of defending themselves. They were the original gentle giants, 50 tonnes of
floppy flesh and not a spine, horn or a serious armour plate to be seen. But maybe, like modern
elephants, they were immune from attack. It would have taken a dozen Allosaurus to gang up
against one Apatosaurus adult and bring it down. Probably they wouldn‘t have been able to pull it
off. Some of the smallest dinosaurs were carnivores. Ornitholestes, for example, was 2 metres
long. A small bird-like biped it scuttled rapidly from bush to bush snatching up lizards and
mammals when it could. The big flesh-eaters (theropod) included Allosaurus, which was a 12
metre-long beast. It was a classic carnivorous dinosaur: long powerful legs, a heavy tail for
counterbalancing, small arms but with powerful hands, and a deep skull with powerful jaws.
USING THE FOSSIL RECORD TO PROVIDE EVIDENCE OF CHANGES IN ORGANISMS THROUGH
GEOLOGICAL TIME
The fossil record provides evidence of changes in floras (plants) and faunas (animals) through
geological time and of the progressive development of higher life forms.
Early life on Earth: Sedimentary rocks from Greenland, 3850 Ma old, contain carbon in a form
interpreted as evidence of biological activity. These chemical fossils are thought to have
originated from marine bacteria or bacteria-like organisms. The implications are that life must
have originated some time before this, perhaps as long ago as 4000 Ma. From about 3500 Ma,
Stromatolites (mound-like structures produced mainly by cyanobacteria) became increasingly
abundant in warm shallow sea. It was photosynthetic bacteria like these that produced most of
the oxygen that eventually formed a permanent accumulation in the atmosphere, reaching about
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15% of its present level by 2000 Ma. Life probably consisted only of simple bacteria in which
the genetic material was not enclosed in a nucleus (i.e. prokaryotes) until about 2100 Ma ago,
when more complex cells with nucleus (eukaryotes) first appear in the fossil record.
Multicellular algae had evolved by about 1200 Ma ago. Late Precambrian times, from around 610
Ma ago, a variety of large, soft-bodied, marine organisms appear in the fossil record around the
world. Collectively called the Ediacaran fauna, they are generally regarded as the first
multicellular animals, and they include worm-like, jellyfish-like, sponge-like and other creatures
of unknown relationships. By the end of the Precambrian, 545 Ma ago, all life was still confined
to the sea, and almost no animals with hard parts (such as shells, skeletons or teeth) had yet
appeared.
Life in the Palaeozoic: During the Cambrian explosion, starting about 545 Ma ago, animals with
hard parts, and animals capable of making complex trace fossils, proliferated in a sudden burst
of evolution. From the fossil record, almost all animal groups seem to appear at this time. From
the early Cambrian into the Ordovician there was a massive increase in the abundance and
diversity of readily fossilisable shallow marine organisms. Some such groups were confined to
the Palaeozoic, including for example, trilobites, graptolites, and jawless fish. The main invasion
of the land in Silurian times, initially by small plants, and by detritus-eating arthropods, opened
up a vast new range of ecological niches. Jawless fish swarmed into freshwater rivers and lakes,
to be followed by jawed fish, which were much more efficient predators. In plants, competition
for light soon forced structural improvements for increased height, stem strength and better
reproduction. The forests of the late Palaeozoic became populated with distinctive plant groups
that provided the basis for the terrestrial food chain. After emerging in late Devonian times,
tetrapod (four-legged animals) initially dominated by large amphibians, especially in
Carboniferous; their key innovation was shelled amniotic egg, freeing them from dependence on
nearby bodies of water for reproduction. During the early Carboniferous times, the existence,
shallow tropical seas again promoted the development of coral reefs, which in turn supported a
huge diversity of reef-dwelling organisms such as bryozoans, trilobites, molluscs, crinoids,
brachiopods, and a great variety of fish. In the late Carboniferous equatorial forests and lowlying swamps, massive amounts of plant debris accumulated, later becoming coal. By the end of
Palaeozoic, many of the major and subsequently important groups of animals and plants had
evolved. Important exceptions include the modern amphibians (frogs, toads and newts),
dinosaurs, crocodiles, lizards, snakes, turtles, tortoises, mammals, birds and flowering plants, all
of which were to appear in the Mesozoic. The end of Permian times saw the final assembly of
the supercontinent Pangaea, a lowering of sea level, a huge reduction in the area of shallow
seas, the eruption of enormous thicknesses of basalt lavas in Siberia and the release of
methane trapped in sediments below the oceans, all of which must have led to climate change.
Although the chain of cause and effect is not certain, the end of the Palaeozoic Era was
marked by the biggest extinction event in history of life, with 57% of marine families and 70%
of land vertebrate genera becoming extinct.
Life in the Mesozoic: The Mesozoic Era began with many forms of life struggling to recover
from the devastating extinction at the end of the Permian. This period saw very important
changes in land-dwelling vertebrates. They developed increasingly mammal-like features, such
as being warm-blooded and having hair. The first mammals appeared in the late Triassic, and
for much of the Mesozoic mammals remained small, shrew-like animals living in nooks and
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crannies of the dinosaur world. Supporting all Mesozoic animal life on land was a distinctive
plant-based food chain, dominated by conifers, cycads and ferns. In the early Cretaceous, one
group of plants gradually acquired flowers. Dinosaurs first appeared in the late Triassic, and
reached their greatest diversity in the late Cretaceous. There were two groups, distinguished
mainly by their pelvic structure. The saurischians included bipedal carnivores and quadrupedal
herbivores. The ornithischians were all herbivores, and were either bipedal or quadrupedal.
Birds arose in the late Jurassic from a group of small bipedal saurischian dinosaurs. The seas of
the Jurassic and Cretaceous abounded with life, including ammonoids. In the late Cretaceous,
countless billions of tiny coccolith plates from phytoplankton carpeted the sea floor, and they
form the bulk of the sedimentary rock called chalk. The K-T extinction at the end of the
Mesozoic re-set the evolutionary stage, providing opportunities for groups that had been more
or less waiting in the wings during the Mesozoic Era, such as flowering plants, insects, birds and
mammals.
Life in the Cenozoic: The Cenozoic Era began with parts of the biosphere recovering from the
end Cretaceous mass extinction. Gone forever were groups such as ammonoids, belemnites,
ichthyosaurs, and mosasaurs from the seas; dinosaurs from the land, and pterosaurs from the
skies. The evolutionary radiation of mammals in the early Cenozoic led to groups adapted to a
huge range of modes of life and environments. Some, for example, took to the sea (whales)
while others took to the air (bats). As landscapes became increasingly coloured by flowering
plants, many insects, birds and other animals evolved adaptations to match those of specific
plants, whilst others were more generalist feeders. The payoff to the flowering plants for
snacks of nectar they offered to insects was assistance in pollination. Larger animals, especially
birds and the newly evolved mammals played their part in distributing seeds from fruit and
nuts. The evolution of grasses (a group of flowering plants) in the mid Tertiary had farreaching effects on animal life. Forests were replaced on many continents by open grasslands,
favouring grazers over browsers. The evolution of horses, for example, was strongly influenced
by these changes in vegetation.
EDIACARAN FAUNA
Fossils widely interpreted as ―animals‖ appear in the stratigraphic record of the Ediacaran
Period (635-542 Ma), representing the last 90 Ma or so of the Proterozoic Era. From many
parts of the world, assemblages of characteristic types of body fossils have been recovered,
which are collectively known as Ediacaran faunas after the Ediacara Hills in South Australia,
from where many such fossils were first described in detail.
What were these Ediacaran organisms like? Perhaps their most striking feature is their size,
with some of the later forms reaching up to a metre or so in length. Yet there is no evidence
for them possessing any supporting skeletal hard parts. The fossils are found in sedimentary
rocks originating from a variety of environments, ranging from deep marine areas to inshore,
even tidal flat, settings. The fossils are preserved only as flattened impressions at the base of
sandy to silty sedimentary layers. Palaeontologists, therefore, have to interpret the Ediacaran
fossils on the basis of their gross anatomical features. Needless to say, opinions have differed
quite sharply.
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The Ediacarans are named after the fossil locality in Australia where they were first
discovered in 1946 by Reginald Sprigg, an assistant geologist to the government of South
Australia, who was examining old silver and lead mines in the Ediacaran Hills about 600km north
of Adelaide with a view to prospecting them for uranium.
Apart from his professional interest in uranium, Sprigg was also interested in the possibility
that there might be fossils in these Precambrian sequences – an idea that at the time flew in
the face of received palaeontological reason. As he approached the old mines he noticed
numerous impressions of what he took to be jellyfish outcropping in the sandstones of the area.
He took the fossils home, wrote them up in a scientific paper and then spent years trying to
convince the palaeontological community that they were indeed evidence of Precambrian life.
By the late 1950s, additional material had been found in the Flinders Range (also in Australia)
and, crucially, from another locality on the other side of the world. In Charnwood Forest,
Leicestershire, England, a schoolboy named Roger Mason came across a large fern-like
impression embedded in the sandstone. Mason contacted a palaeontologist at the University of
Leicester, Trevor Ford, who returned with him to the quarry. Over the next few months they
discovered several other fossils, and the Ediacarans were firmly established as a real page in
the book of life.
But what are the Ediacarans? The crucial pointer is something that Sprigg noticed instantly:
they looked like jellyfish. Jellyfish are a member of a group of animals known as cnidarians
(pronounced with a silent ―c‖: ―nigh-dare-re-uns‖). These animals have a central mouth around
which are stinging tentacles for catching prey and possess a radial symmetry and only two
layers of body tissue, which are separated by a layer of jelly. Such creatures are quite
different from the rest of the animal world, the bilateria, which are characterised by having
bilateral symmetry (they have front and a back end, as well as an upside and a downside) and
three layers of body tissue. Thus, it seemed quite reasonable to these palaeontologists who
carried out the first in-depth study of the Ediacarans to consider them as an evolutionary step
on the road to the more complex bilaterian animals of the Burgess Shales (Cambrian period).
This view was, however, rejected as some of the strangest Ediacarans that do not look like
jellyfish appear to have no living relatives and are actually bilaterally symmetrical themselves.
It was only in 1983 that Adolf Seilacher of the University of Tübingen, Germany came up with a
different and radical notion. Seilacher noted that the animals of Ediacaran times do not fit into
the same basic body plan categories as the animals around today (or indeed those of the
Burgess Shale fauna). In his view, the Ediacara consisted of separate segments, similar
superficially to a quilted airbed, a plan that is not used in oceanic organisms today. To him,
these creatures seemed to show continuous gradations in form between animals that, if they
were alive today, would be placed in separate phyla.
This view of the Ediacarans as being morphologically unique has recently been acknowledged and
amplified by palaeontologists working at Oxford University. They have suggested that many of
the Ediacaran fossils are different life stages of the same creature. This, of course, does not
help with understanding what the Ediacarans were exactly. There are some who say that they
were not multicellular animals at all: recent speculations have suggested that they may have
been related to the group that gave rise to the fungi or were perhaps even lichens.
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Whatever the truth, the consensus is that the Ediacarans area group of wholly extinct
multicellular organisms that were constructed around a different type of tissue organisation
from the rest of the known animal world. To many, these creatures, lacking head and gut, are a
failed experiment in the history of life.
CAMBRIAN EXPLOSION
At the same time as the Ediacarans were alive, the ancestors of modern bilaterian animals were
diversifying with unprecedented rapidity; this is called the Cambrian Explosion. The Cambrian
Explosion represents a time when a large range of different bilaterian body plans appeared in
the fossil record in a relatively short period of time. So major were these body plans that they
equate to different phyla – one of the highest taxonomic groupings.
From the palaeontological perspective, there are only a few localities where the world of the
earliest Cambrian fossils may be observed: the famous Burgess Shales fauna of British
Columbia (~505 Ma); the Chengjiang fauna of China (~520 Ma); and Sirius Passet fauna of
Greenland (~530 Ma).
The reason why the evolutionary radiation (phase of significant increase in numbers of species
within groups of organisms) of bilaterians in the early Cambrian is considered an ―explosion‖ is
because no examples of these bilaterian animals have been found at the time of the Ediacarans
while by ~530 Ma (Sirius Passet fauna) there were abundant examples of animals that can be
recognised as ancestral to those alive today. These animals, therefore, appear in the fossil
record within a relatively short period of approximately 12 million year duration – hence the
term ―explosion‖.
The Cambrian Explosion marks the start of the Phanerozoic (and, therefore, the end of the
Precambrian and the start of the Cambrian period) and is set at 542 Ma. The exact timing of
this event has been narrowed down by examining fossils from Siberia where, on the banks of
the Kotuikan River, one of the most spectacular Cambrian outcrops in the world can be
observed. At river level (where there are only a few unspectacular trace fossils), the rocks
have been dated at about 545 Ma. A little higher in the section is a sandstone unit that has
been dated as 544 Ma – a difference of only a few million years. Starting immediately above
this unit, small shelly fossils begin to increase in frequency up the rock section. It is the first
evidence of the shelly part of the Cambrian Explosion. Further up the cliff, sediments dated at
530 Ma contain more than 80 recognisable groups, including small shelly fossils, tracks, burrows
and trails.
The first 10 Ma or so of the Cambrian was a time when small shelly fossils developed in
diversity and complexity. It seems likely that the ancestors of molluscs evolved at this time.
After this 10 Ma period and close to the end of the early Cambrian, crown groups (a group of
closely related organisms that includes the common ancestor plus all its descendants) of all the
major phyla alive today (e.g. annelids, arthropods and brachiopods, and the chordates that
ultimately gave rise to the vertebrates and humans) suddenly appeared on the scene.
Obviously, these fossils cannot have appeared fully formed out of nowhere. The latest studies
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using molecular clock evidence indicate that the bilaterians arose between 630 and 600 million
years, i.e. between 88 and 58 million years before the Precambrian-Cambrian boundary. This
clearly implies that the bilaterians were contemporary with the Ediacarans; but there is
virtually no evidence for this from the fossil record. The reason for this lack of evidence is
believed to be because all the late Proterozoic bilaterians were microscopic and, hence, were
not preserved in the fossil record. Whether or not the bilaterian animals did originate in the
Ediacaran (as seems overwhelmingly likely), it is clear that they evolved very rapidly in the early
Cambrian, possibly due to an extinction event that decimated the Ediacarans.
The causes of the Cambrian Explosion fit into two categories: internal (i.e. biological) and
external (i.e. environmental) causes. The late Proterozoic and the early Phanerozoic were times
of enormous environmental change. One of the more remarkable ideas to emerge from the
study of ancient climates in recent years has been the suggestion by Harvard geologists Paul
Hoffman and Dan Schrag that the world was covered from the Equator to the poles by a sheet
of ice several kilometres thick in the late Proterozoic termed snowball Earth. They base this
extraordinary conclusion on a combination of geochemistry and classical geology. Evidence for
these extreme ice ages has been found in rocks of late Early Proterozoic age from every
continent. Evidence from fossil magnetism in the rocks shows that some of these glacial
sediments were deposited at low latitudes; so, it is clear that these ice ages were global in
extent. So prevalent were these extreme climatic conditions that the penultimate period of
Precambrian time has been named the Cryogenian. Geological and geochemical evidence is
unclear, but there may have been as few as two or as many as five Early Proterozoic ice ages;
the important point to remember is that the last global glaciations (the Marinoan) occurred at
about 635 Ma and marked the end of the Cryogenian.
Not all geologists agree that the earth was entirely covered in ice during the late Precambrian.
Some suggest that the ice sheets extended to lower latitudes than they did at the height of
the late Quaternary glaciations, into temperate zones and, as this idea is not as radical as the
idea of a snowball Earth, it is known as the ―slushball Earth‖ compromise.
It is tempting to think that the rapid diversification of the Ediacarans and the bilaterian
animals of the Cambrian Explosion were direct responses to the environmental changes at the
time. However, it seems likely that, in addition to the environmental instability, there may have
been an internal trigger related to changes in the genetics of these organisms.
Many of the Cambrian animal forms exhibit greater differentiation of body parts, including the
concentration of food-trapping organs around a ―head‖ end, and the appearance of limbs and of
discrete tabular, two-ended guts; these innovations may have been fuelled by Hox gene
duplication. Hox genes are responsible for pattern formation, i.e. the overall arrangement of
appendages on the body of an organism. Comparison of the distribution of Hox genes in
diploblastic animals (animals having a body made of two cellular layers) and their distribution in
the bilateria shows that there are many more varieties of Hox gene in the latter group than
the former. The much higher number of Hox genes in the bilaterians was probably caused by a
process known as gene duplication. These duplicated Hox genes were subsequently assigned to
control the development of other body parts. Hence, it seems likely that Hox genes are
implicated in the rapid diversification of the bilaterians in the early Cambrian. Hox genes may
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be thought of as an internal trigger that drove the Cambrian Explosion.
Some scientists have suggested that the regulation of many new body patterns by Hox gene
control allowed the evolution of the ability to ingest and digest food extracellularly by allowing
the development of the gut. By allowing the bilaterians to eat and digest food particles that are
larger than single cells, organisms were able to grow bigger. This might explain the sudden
appearance of the small shelly fossils in the early Cambrian. Thus, the bilaterian increase in
body size may be ascribed to the development of a novel method of feeding that developed due
to the sudden evolutionary deployment of a powerful set of genes. The evolution of the ability
to ingest and digest food extracellularly in turn set up a whole new world of competition, as
predators and prey suddenly indulged in the biggest evolutionary arms race of all time. It would
also support the notion that the stem-group bilaterians were contemporary with the
Ediacarans. From the evidence of molecular genetics, it seems likely that the Hox gene
revolution occurred sometime between 650 Ma and 550 Ma, precisely coincident with the
Ediacarans and the implied hidden beginnings of the Cambrian Explosion.
The terminal Proterozoic snowball event (the Marinoan glaciation) finished at about 635 Ma,
the same time as when the Hox gene revolution may have been occurring. A heavily glaciated
Earth would have forced the early biota into refugia (i.e. geographically isolated populations)
with limited gene flow between them that could have stimulated rapid evolution and easy
dissemination of novel gene plans like the Hox complex. In addition, it has been suggested that
there may have been an increase in deep-ocean oxygen concentration at this time, which would
have assisted organisms increasing in size in the rapidly evolving bilateria.
BURGESS SHALES
High in the Canadian Rockies is exposed a deposit of middle Cambrian age, about 530 Ma old,
called the Burgess Shales. It contains the fossils of animals that lived on a muddy sea floor,
and which were suddenly transported into deeper, oxygen-poor water by submarine landslides.
Their catastrophic burial has given us an exceptional view of Cambrian life. Not only have
animals with hard shelly parts been preserved but entirely soft-bodied forms are also
preserved as thin films on the sediment surface.
Some of the most common Cambrian fossils, which appear immediately after the first shelly
fossils, are trilobites. These were a group of exclusively marine arthropods, members of the
enormously diverse phylum of animals with jointed, external skeletons that today include forms
such as crabs, lobsters, insects and spiders. The trilobite fossils of the Burgess Shale are like
many trilobites found elsewhere but exceptional in that not only is the main part of their outer
skeleton (or exoskeleton) preserved, but so too are their appendages such as antennae and legs.
Elsewhere, trilobite appendages are extremely rare as they were poorly mineralised.
Only about 15% of the 120 genera present in the Burgess Shales are shelly organisms such as
trilobites and brachiopods that dominate typical Cambrian fossils assemblages (fossils that
occur together) elsewhere. The shelly component was therefore in a minority, and organisms
with hard parts probably formed less than 5% of individuals in the living community. So, if the
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soft-bodied fossils of the Burgess Shales are taken away, all that remains is a typical Cambrian
assemblage of hard-bodied organisms. This is important to bear in mind when geologists try to
interpret other Cambrian fossil assemblages because they also have been dominated by softbodied animals, even if the only fossils they now contain are of hard-bodied ones.
Another important revelation of the Burgess Shale lies in the wide diversity of animal types
that were around in the middle Cambrian time, about 530 Ma ago. There are representatives of
about a dozen of the phyla that persist to the present day. One form closely related to early
arthropods was Anomalocaris, the largest Cambrian animal, some individuals of which may have
reached two metres in length. Its extraordinary jaw consisted of spiny plates encircling the
mouth.
About a dozen types of Burgess Shale fossils have been said to be so unlike anything living
today an so different from each other that, had they been living now, each would have been
placed in a separate phylum. With further study, however, the relationships of these puzzling
animals are becoming clearer. It seems that some Burgess Shale forms are hard to classify
simply because the boundaries between major categories of animal life were still blurred
shortly after the Cambrian Explosion. In other words, by mid Cambrian time, there still had not
been enough time for some groups to have diverged sufficiently from their recent common
ancestors to be distinctly different.
Burgess Shale-type faunas have been found in about 30 sites ranging from North America and
Greenland, to China and Australia. The wide range of animals they contain seems to reflect an
unpruned ―bush of diversity‖ resulting from the Cambrian explosion. Not long after, though,
extinction lopped off some of the branches, leaving phyla with the relatively distinct features
that have remained to this day.
RADIATIONS: THE DIVERSIFICATION OF ORGANISMS
The Cambrian explosion is the first of the major evolutionary radiations that can be charted in
some detail from the fossil record. The diagram over the page shows the total numbers of
marine animal families based on fossil evidence and estimated to have coexisted at different
times through the Phanerozoic. This graph is based on the work of Dr. Jack Sepkoski (an
eminent palaeontologist at the University of Chicago). In the 1980s, Sepkoski scrounged
together as much fossil data on marine invertebrates as he could get his hands on and plotted a
graph of global diversity over the past 500 million years or so. His graph shows a general
increase in diversity over time through the Phanerozoic but with peaks, troughs and plateaux.
The steep rise in families throughout the early Cambrian represents the Cambrian Explosion.
Throughout this episode, the evolution of new organisms clearly exceeded the number of
extinctions over this time. This illustrates the importance of the evolutionary innovations that
occurred in the bilateria during the Ediacaran period and how they fuelled rapid diversification
in the early Cambrian.
Another striking pattern from Sepkoski's curve is that global diversity of marine invertebrates
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appeared to plateau for around 200 million years (from the end of the Cambrian (Cm) to the
end of the Permian (P) and then increases dramatically after the big mass extinction at the end
of the Permian.
The diagram above shows the three main evolutionary faunas found in the Phanerozoic; the
Cambrian fauna; the Palaeozoic fauna; and the Modern fauna. The Cambrian fauna diversified
more rapidly during the initial radiations of the Cambrian than the others. However, it soon
experienced a decline in numbers as the next Palaeozoic fauna began to diversify. The
Palaeozoic fauna continued diversifying until the late Ordovician, after which it began a slow
decline in numbers, while the modern fauna diversified rapidly eventually dominating the scene
in post-Palaeozoic times.
Note that the names of these evolutionary faunas refer only to their times of dominance: all
three existed throughout the Phanerozoic (although very few members of the Cambrian
evolutionary fauna survived beyond the Palaeozoic). Moreover, the three faunas do not
represent discrete sets of animal groups, as in some cases different classes from a single
phylum have been allocated to all three faunas, e.g. the molluscs, which consist of bivalves
(clams) and gastropods (snails) in the modern fauna and cephalopods (including ammonoids) in
the Palaeozoic fauna.
The Cambrian fauna consisted of mud loving animals that lived on the ocean floor, such as
trilobites and other arthropods. This fauna are characterised by low diversity of species and no
tiering into the water column above. The Palaeozoic fauna was dominant in the lower Palaeozoic,
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although there are still remnants left today. Although benthonic animals still existed, the fauna
in the Palaeozoic was much more tiered, with filter feeding organisms that grew up into the
water column such as corals and crinoids, as well as the first pelagic (swimming) life. The main
predators were the Cephalopods who pressurised organisms to adapt and swim faster to escape
predation. The Modern fauna is very diverse, consisting of over 600 families. The fauna is
characterised by increased tiering, with more pelagic organisms, as well as infaunal ones too.
This last point is perhaps the most remarkable of the faunal changes and was probably due to
predation. Although predation was common in the Cambrian, with some predators even ingesting
entire small shelly prey, throughout the Palaeozoic there was an abundance of exposed shelly
animals that lived rooted to the sea floor with evidence for predation (e.g. in the form of
diagnostic damage preserved in the fossil shells, rare. Throughout the Mesozoic, however, the
fossil record shows several lines of evidence for a marked increase in the intensity of
predation on shelly prey. These include an increase in the frequency of predator damage, the
appearance of many new kinds of predators with specialised adaptations for tackling shelly
prey, and the emergence of new defensive adaptations by shelly organisms (prey).
Among the newly evolving predators in early Mesozoic seas were crabs and lobsters, along with
several groups of fish and marine reptiles, variously equipped to crush, smash or pierce shells.
New kinds of starfish evolved the ability to pull open bivalves, insert their stomachs into the
opened shell and digest the occupant. In the Cretaceous, these adapted predators were joined
by gastropods capable of drilling through shells to reach their prey, while in the Tertiary,
various shell-breaking birds and mammals evolved.
Although the major groups of shelly prey animals developed a wide variety of protective
adaptations, the one notable evolutionary trend was a boom in burrowing, especially by bivalves
and echinoids (sea urchins) to increasing depths in the sediment to gain refuge from predators.
The extent of bioturbation of the surface sediment was therefore both intensified and
deepened. This effect, coupled with an increase in the amount of disturbance at the surface by
grazers, detritus feeders, and the number if predators excavating the sediment, made the
larval settlement of shelly animals permanently anchored to the surface more hazardous. As a
result, these animals, so successful in the Palaeozoic, waned in relative diversity.
These linked ecological changes can be collectively referred to as the Mesozoic marine
revolution, transforming the character of marine life to that which can still be recognised
today. It could also be said that this revolution helped to provoke the unprecedented rise in
marine animal diversity observed in the fossil record from the end of the Palaeozoic onwards.
The most profound feedback to biogeochemical cycles, however, came from associated changes
in plankton. Starting in the late Triassic, groups of microscopic plankton with calcareous
skeletons began to appear, including both single-celled algae (e.g. coccoliths) and protists (e.g.
planktonic foraminifera). One potential interpretation of this simultaneous adoption of
calcareous skeletons (or toughened organic walls in some other planktonic groups) by both
groups is as a defensive adaptation against increased grazing pressure. The associated rise of
various planktonic groups with calcareous skeletons enhanced carbonate sedimentation in
deeper water, increasing the oceanic carbonate sink.
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MASS EXTINCTIONS
Diversity has not increased indefinitely, however, and the graph suggests a temporary lull in the
late Cambrian before diversification rapidly increased during the Ordovician. The reasons for
the late Cambrian lull are still unclear. However, more striking is the sharp drop in diversity
accompanying the close of the Ordovician, which marks the first of five large-scale mass
extinctions during the Phanerozoic as shown by the fossil record. These five mass extinctions
include:
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



Ordovician – 440 million years ago
Devonian – 365 million years ago
Permian – 245 million years ago
Triassic – 210 million years ago
Cretaceous – 65 million years ago
In fact there has been a significant extinction episode terminating each of the geological
periods, although not a mass extinction. Mass extinctions all share the same characteristics:



over 30% of species became extinct each time
extinct forms came from a broad range of habitats (terrestrial and marine) and sizes
extinctions happened within a short time span (which indicates a single cause or a
cluster of interlinked causes)
A wide range of different hypotheses has been put forward that might explain these
devastating changes to life on Earth. Studies of the five major periods of mass extinction
indicate that there was no single cause. In several, climatic change is involved, with associated
sea level and habitat changes. Extra-terrestrial impact has been implicated in the extinctions
at the Cretaceous-Tertiary (K-T) boundary along with eruptions of flood basalts.
The fact that humans are here today is a demonstration of the fact that mass extinctions do
not affect all living organisms, and that there are always survivors.
What decides who lives and who dies, and how do you win in the ultimate evolutionary challenge?
Of course, the causes of a mass extinction will be an important influence on which species
become extinct - for example a fall in sea level will have a devastating effect on the life in the
shallow seas, as their habitat simply ceases to exist, but will have little or no direct effect on
life on land. Indeed, the species which disappear in an extinction can tell us a lot about the
characteristics, or even the cause of the extinction - if only species adapted to live in warm
climates vanish, then it is reasonable to assume that global cooling is associated with an
extinction.
However, even when huge areas of habitat disappear or are changed drastically some species
will survive, though we don't know for sure what determines these 'winning' species. Luck
almost certainly plays an important part in deciding who makes it through a mass extinction, and
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there are some characteristics which seem to help tilt the odds in favour of surviving. It helps
to be an opportunist and a generalist, able to survive in a wide variety of conditions and to
quickly take advantage of any favourable changes, and it helps to be able to reproduce quickly
and prolifically. Cockroaches, rats and mice almost certainly have little to fear from mass
extinction!
K-T MASS EXTINCTION
The extinction at the boundary of the Cretaceous and the Tertiary periods, 65 million years
ago, is the most famous of all mass extinctions. Its fame comes not from its magnitude (the
Permian extinction was far larger) but from the victims of the extinction - the dinosaurs. The
Cretaceous-Tertiary (or 'K-T') extinction wiped out around 85% of all species.
The dinosaurs were not the only victims - pterosaurs (flying reptiles), mosasaurs and other
marine reptiles, fish, brachiopods, plankton and many plants either died out completely or
suffered heavy losses. Even the ammonites, who had survived 4 previous extinctions, finally
disappeared. For some reason though, some groups seem to have been almost entirely
unaffected by the K-T extinction - crocodiles, turtles and lizards, mammals and birds all made
it through relatively unscathed.
The cause of the K-T extinction event has been the subject of intensive research, with many
hundreds of research papers published since 1980, when Luis Alvarez reawakened interest in
the subject, almost overnight. He suggested that the death of the dinosaurs, and all the other
victims of the K-T extinction, was due to a giant meteorite crashing into the earth, severely
disrupting the earth's ecosystem. While other theories suggest volcanic activity, climate
change, environmental pollution or even cosmic radiation as causes, the asteroid impact theory
remains the most probable - at least for the moment.
In 1979 Luis Alvarez found high concentrations of iridium (an element of the platinum group) in
a thin geological deposit separating Cretaceous rocks from Tertiary rocks. This boundary
became known as the K-T boundary. Above this boundary the fossils shift from a prevalence of
dinosaurs and a few small mammals to no dinosaurs and a prevalence of mammals.
Iridium has a strong affinity for iron, so during the formation of the planet most of it had been
drawn into the iron-rich core of the Earth. The presence of this layer of iridium so close to the
Earth's surface became a mystery. Alvarez realised that iridium is also abundant in some
meteorites and led him to publish his Asteroid Impact Theory as a cause for the Cretaceous
mass extinction.
The theory suggests that 65 million years ago, a meteorite 10 km in diameter crashed into the
Earth at 72,000 kph (an impact force far greater than the detonation of all the nuclear
weapons in the world). This impact literally shook the Earth, ignited fires, formed tidal waves
and set off dust clouds that covered the planet, either cooling the atmosphere by blocking out
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the sun or warming it by trapping in the heat. As this dust settled it formed a clay layer 0.5cm
thick including iridium from the impacting meteorite.
Many of these K-T deposits also had tiny pieces of "shocked quartz" within them and was used
as further evidence for a meteorite impact. Quartz develops very curious properties when
impacted at high pressures, as had previously been recognised around several well-known
meteorite impact craters.
Eventually an impact crater site was found on the Yucatan Peninsula in Mexico. The Chicxulub
Crater, as it is called, is some 200 km across and deeply buried by sediments deposited over
the last 65 million years. Boreholes have revealed that high levels of sulphur concentrations are
found in the rocks in this area. The impact would have vaporised these sulphur-bearing
sediments sending clouds of acid rain to add further misery to the planets inhabitants! Various
pieces of evidence have been gathered from the surrounding areas of the Caribbean and North
America that point to a massive impact. The enormous pressures due to shock waves generated
at sites of meteorite impact, or where an atomic bomb has been exploded produce multiple
layers in grains of quartz. These pieces of shocked quartz have been found in the K-T boundary
layer in many sites in North America and more sparsely in Europe and the Pacific. There was a
gigantic impact and it was clearly in or near North America. When molten rock is splashed out
into the air small spherules of glass and larger bead-like objects (called tektites) are formed;
these have been found as far away as several hundred kilometres to the east in Haiti and to the
north-west in Colorado and New Mexico. Rocks thrown out of the crater, which has the
particular features of the underlying rock in Yucatan, have been found as far afield as Central
Canada. Evidence of a tsunami has been found in Texas and north-eastern Mexico. This giant
wave, having swept over the land and then drained back to the Gulf of Mexico, deposited leaves
in marine sediments. Studies on crater indicate that the asteroid had a diameter of ten
kilometres, which was the diameter calculated by the Alvarezes from the global quantity of
iridium. Material would have been blasted out to a depth of 14 kilometres.
The dinosaurs were killed by an asteroid, right? Well, maybe not! Despite the fact that there is
good evidence that the Earth was hit by a giant asteroid 65 million years ago, marking the end
of the Cretaceous period, the cause and effect relationship between the impact and the death
of the dinosaurs has never been proven. And, there is compelling fossil evidence that the
effect of the impact was largely limited to North America, in other words, close to the impact
site.
Professor Bob Spicer of the Open University's Earth Science Department has been examining
what happened to the plants and forests at the time of the impact. Unexpectedly, the only
ecological damage on land appears to have been within the North American continent. Elsewhere
in the world there were changes to vegetation but, in north-eastern Russia for instance, the
changes occurred several million years before the time of the impact, not at the impact level
itself. Clearly, something else was happening at the end of the Cretaceous period which was
profoundly affecting the global environment. So, if the effect of the asteroid impact was
smaller than previously thought - then what did kill the dinosaurs? Some people have suggested
the eruption of the Deccan Traps in India might have been responsible.
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The Deccan Traps are the remnants of a continental-scale volcanic event, one of only half a
dozen such cataclysmic events on Earth since the rapid radiation of multicellular life 600 million
years ago. Many researchers now believe these massive volcanic events were so rapid that they
may have literally blasted the ecosystem and Earth's atmosphere into a mixture of toxic gases
and enormous ash clouds. In case of the Deccan, up to one million cubic kilometres of lava
rapidly erupted through enormous fissures across the area that now lies inland of Bombay (that
is equivalent to covering most of Western Europe to a depth of 500 meters with lava flows!).
Calculations demonstrate that these eruptions would have created up to 10 billion tonnes of
sulphuric acid and released 3 billion tonnes of carbon dioxide which would have increased global
atmospheric carbon dioxide levels by 75 parts per million. This, and associated climatic effects,
would have been enough to make more than one Tyrannosaurus retch, and might have had a
similarly devastating effect on the world's living organisms.
Most crucially, new evidence from the Deccan indicates this eruption occurred within the same
time frame when the asteroid hit. Therefore, we should be looking instead for evidence for two
synchronous catastrophes as a mechanism for destroying the dinosaurs. If this is the case,
then we are looking at one of the most amazing periods of Earth history. A double cataclysm
that would have taken the planets' atmosphere and biosphere to the very limits of recovery.
To test what effect the Deccan eruptions had on the Indian environment Professor Spicer and
his colleagues at the OU are examining the fossils from the ancient soils trapped between the
lava flows and dating the flows themselves using the latest isotopic techniques. Working in
collaboration with Indian scientists they are looking at dinosaur nests, pollen grains, leaves,
wood and even fossil frogs to see if Deccan eruptions had any affect, and if so what that
effect was. So far the work is showing that "life found a way" even under the most apocalyptic
of conditions. So, the researchers point out, preliminary results show that we might have to
look again for reasons why the dinosaurs died out.
END-PERMIAN EXTINCTION
The biggest extinction of the past 500 My, the end-Permian event (251 My ago), witnessed the
loss of as much as 95% of all species on Earth. Key questions to geologists concern what
combination of environmental changes could possibly have has such a devastating effect.
Evidence on causation is equivocal, with support for either an asteroid impact or mass volcanism.
Three key pieces of evidence for the K-T impact are the Chicxulub crater in Mexico, the
iridium spike and shocked quartz. All three phenomena were reported from Permo-Triassic
(PTr) beds in the 1980s and 1990s, and all three have been either rejected or greeted with
lukewarm enthusiasm at best. The evidence is far weaker and more limited than for an impact
at the K-T boundary.
At the end of the Permian, giant volcanic eruptions occurred in Siberia, spewing out some 2
million km3 of basalt lave, and covering 1.6 million km2 of eastern Russia to a depth of 400-3000
metres, equivalent to the area of the European Community. It is now accepted widely that
these massive eruptions, confined to a time span of <1 My, were significant factor in the end-
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Permian crisis.
The suggestion to this effect was first made in the 1980s. Russian geologists had explored the
Siberian Traps long before then, but were unsure of their age. The Siberian Traps are
composed of basalt, a dark-coloured igneous rock, which is generally not erupted explosively
from classic conical volcanoes, but usually emerges more sluggishly from long fissures in the
ground (as seen in Iceland). Flood basalts typically form many layers and can build up over
thousands of years to considerable thicknesses. They produce a characteristic landscape, called
trap scenery, where the different lava flows erode back through time, producing a layered,
stepped appearance to the hills (the word ‗trap‘ comes from the old Swedish word trapp,
meaning a staircase).
Early efforts at dating the Siberian Traps produced a huge array of dates, from 160 to 280
Mya, with a particular cluster between 230 and 260 Mya. According to these ranges, geologists
in 1990 could conclude only that the basalts might be anything from Early Permian to Late
Jurassic in age, but probably spanned the PTr boundary. More recent dating using newer
radiometric methods, yielded dates exactly on the boundary, and the range from bottom to top
of the lava pile was, 600 000 years, highlighting that the event occurred, geologically speaking,
overnight. In addition, this kind of time duration for the eruptions matches the evidence from
China of a rapid extinction.
Consequently, with increasingly precise dating, the Siberian Trap eruptions have moved from
having been allotted a relatively minor role in the PTr crisis as part of a complex web of
interacting processes, to being the most probable trigger for the catastrophe. However, there
are still unresolved debates concerning the accuracy of the new dates. Some scientists have
even suggested recently that the massive flood basalts were actually themselves caused by a
giant extra-terrestrial impact, which tore deep into the continental crust of that part of
present-day Siberia. However, the nature of the eruptions casts doubt on such a model, and
there is no evidence that any volcanism on the Earth, or indeed on any other planet, was
triggered by an impact.
To investigate the faunal, floral and environmental changes in more detail, continuous
fossiliferous rock sections through the PTr crisis need to be studied. In the late 1980s, few
such sections were thought to exist and those that had been studied previously were thought
to contain significant gaps right at the crucial extinction interval. Reanalysis of these sections
by Tony Hallam and Paul Wignall, among others, in the early 1990s led to the realization that
the records through the extinction event were much more complete than was believed
previously.
The rocks contain a huge diversity of fossil shells and skeletons, showing that the latest
Permian seas teemed with life. In particular, the Permian sediments are intensely bioturbated,
full of burrows made by a plethora of benthic animals living, feeding and moving through the
sediment. The communities were diverse and ecologically complex. By contrast, sediments
deposited immediately after the extinction event, in the earliest Triassic, are dark-coloured,
often black and full of pyrite. They largely lack burrows, and those that do occur are very
small, and fossils of marine benthic invertebrates are extremely rare. These observations, in
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association with geochemical evidence, suggest a dramatic change in oceanic conditions from
well-oxygenated bottom waters to widespread benthic anoxia. Before the catastrophe, the
ocean fauna was differentiated into recognizably distinct biogeographical provinces. After the
event, a cosmopolitan, opportunistic fauna of thin-shelled bivalves, such as the ‗paper pecten‘
Claraia and the inarticulate brachiopod Lingula spread around the world.
On land too, life was hugely diverse in the latest Permian. Terrestrial tetrapod (amphibian and
reptile) faunas had reached high levels of complexity, arguably as complex as modern
mammalian communities, with four or five trophic levels among carnivorous forms. For example,
the sabre-toothed gorgonopsians fed on thick-skinned, rhinoceros-sized herbivores, whereas
several ranks of smaller flesh-eaters fed on smaller prey. Numerous groups of plants provided a
diversity of habitats, and some floras were endemic, indicating geographical differentiation
relating to climatic zones. The decline and loss of tetrapods has been documented in some
detail in South Africa, where the disappearance of taxa is indicated to be rapid. Comparison of
the timings of species loss on land and in the sea suggests that they were coincident. In many
places, it seems that soils were washed off the land completely, and the only organisms to
survive appear to have been fungi.
Geochemistry gives additional clues about the nature of the environmental changes. Exactly at
the PTr boundary, there is a dramatic shift in oxygen isotope values: a decrease in the value of
the 18O ratio of about six parts per thousand (ppt), which corresponds to a global temperature
rise of ~6°C. Climate modellers have shown how global warming can reduce ocean circulation and
the amount of dissolved oxygen to create benthic anoxia. A dramatic global rise in temperature
is also reflected in the types of sediment and ancient soil deposited on land.
Can the evidence for oceanic anoxia, global warming, a catastrophic reduction in the diversity
and abundance of life be linked to the co-occurrence of the Siberian eruptions in a coherent
killing model? The key might come from further study of carbon isotopes. Values of 13C show a
sharp negative excursion during the PTr interval, dropping from a value of +2 to +4 ppt to -2
ppt at the mass extinction level. This drop implies a dramatic increase in the light carbon
isotope (12C), and geologists and atmospheric modellers have tussled over trying to identify its
source. Neither the instantaneous destruction of all life on Earth and subsequent flushing of
the 12C into the oceans, nor the amount of 12C estimated to have reached the atmosphere from
the carbon dioxide released by the Siberian Trap eruptions are sufficient to explain the
observed shift. Something else is required.
Not only must this new source of 12C be identified, but that source must also be capable of
overwhelming normal atmospheric feedback systems. The only option so far identified is the
methane released from gas hydrates, an idea that has been accepted with alacrity.
The assumption is that initial global warming at the PTr boundary, triggered by the huge
Siberian eruptions, melted frozen gas hydrate bodies, and massive volumes of methane rich in
12
C rose to the surface of the oceans in huge bubbles. This vast input of methane into the
atmosphere caused more warming, which could have melted further gas hydrate reservoirs. The
process continued in a positive feedback spiral that has been termed the ‗runaway greenhouse‘
phenomenon. Some sort of threshold was probably reached, which was beyond where the
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natural systems that normally reduce carbon dioxide levels could operate effectively. The
system spiralled out of control, leading to the biggest crash in the history of life.
Life came close to complete annihilation 251 Mya. A fortunate 5% of species did, however,
survive and understanding how these few taxa recovered from the severest of evolutionary
bottlenecks is crucial to understanding the subsequent evolution of the biosphere. It took 100
Ma for global biodiversity at the family level to return to pre-extinction levels. However,
ecological recovery was somewhat quicker, with complex communities such as reefs becoming
re-established by the Middle Triassic (some 10 Ma after the PTr boundary).
Details of the recovery of the marine ecosystem in the aftermath of the extinction are known
only from two sites, northern Italy and the western USA, both of which were located in
tropical regions during the Early Triassic. Initial benthic low-diversity communities were
composed of small-sized, epifaunal suspension-feeding opportunists, which were living under
suboptimal environmental conditions of low oxygen and low food supply. Microbial mats covered
much of the sea floor. A scarce infauna of small, deposit-feeding animals burrowed feebly just
below the sediment surface. This lasted for maybe a million years. With the disappearance of
benthic oxygen restriction and the increase in food supply, larger and more diverse communities
reappeared slowly. Epifaunal communities increased in complexity as crinoids and bryozoans
returned and began to reach up into the overlying water column. Infaunal communities saw the
return of suspension feeders and finally crustaceans, and the size and depth of burrowing
returned to pre-crisis levels by the Middle Triassic.
Little is known currently of the recovery pattern from elsewhere in the oceans, although work
is ongoing. On land, for millions of years, virtually the only tetrapod was the plant-eating
Lystrosaurus, subsisting on the few surviving herbaceous plants. Forest communities were
absent until the Middle Triassic. Life was clearly tough in the ‗post-apocalyptic greenhouse‘.
If the runaway greenhouse model is correct and explains perhaps the biggest crisis on Earth in
the past 500 My, it is a model worth exploring further. It appears to indicate a breakdown in
global environmental mechanisms, where normal systems that equilibrate atmospheric gases and
temperatures took hundreds of thousands of years to come into play. Perhaps the combination
of global warming and anoxia from gas hydrate release was a cause of other extinction events.
This scenario certainly has been postulated recently for the end-Triassic mass extinction and
for smaller events in the early Jurassic, Cretaceous, and Tertiary.
ALTERNATIVE INTERPRETATIONS OF EVOLUTIONARY PATTERNS
Does evolution occur in rapid bursts or gradually? This question is difficult to answer because
we can‘t replay the past with a stopwatch in hand. However, we can try to figure out what
patterns we‘d expect to observe in the fossil record if evolution did happen in bursts, or if
evolution happened gradually. Then we can check these predictions against what we observe.
What should we observe in the fossil record if evolution is slow and steady? If evolution is slow
and steady, we‘d expect to see the entire transition, from ancestor to descendent, displayed as
transitional forms over a long period of time in the fossil record.
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In the above example, the preservation of many transitional forms,
through layers representing a length of time, gives a complete record
of slow and steady evolution. In fact, we see many examples of
transitional forms in the fossil record. For example, to the right you
can see just a few steps in the evolution of whales from land-dwelling
mammals, highlighting the transition of the walking forelimb to the
flipper.
What would we observe in the fossil record if evolution happens in
―quick‖ jumps (perhaps fewer than 100,000 years for significant
change)? If evolution happens in ―quick‖ jumps, we‘d expect to see big
changes happen quickly in the fossil record, with little transition
between ancestor and descendent.
In the diagram below, you can see the descendent preserved in a layer directly after the
ancestor, showing a big change in a short time, with no transitional forms.
When evolution is rapid, transitional forms may not be preserved, even if fossils are laid down
at regular intervals. We see many examples of this ―quick‖ jumps pattern in the fossil record.
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Does a jump in the fossil record necessarily mean that evolution has happened in a ―quick‖ jump?
We expect to see a jump in the fossil record if evolution has occurred as a ―quick‖ jump, but a
jump in the fossil record can also be explained by irregular fossil preservation.
This possibility can make it difficult to conclude that evolution has happened rapidly. We
observe examples of both slow, steady change and rapid, periodic change in the fossil record.
Both happen. But scientists are trying to determine which pace is more typical of evolution and
how each sort of evolutionary change happens. In many cases, we seem to observe ―bursts‖ of
evolution in the fossil record. For example, imagine that in a lower rock layer, you see ancestor
1. In the next rock layer, you see species 2 and 3. Species 2 looks the same as ancestor 1.
Species 3 is morphologically distinct, but is clearly also descended from ancestor 1. What
happened?
The fossils we have are only ―time slices‖ of actual history. There are several different
hypotheses about what happened which are consistent with these fossil time slices. In order to
determine which of the following hypotheses most accurately explains the pace of evolution,
we‘d need more evidence.
Phyletic gradualism - slow steady divergence of lineages: The ―burst‖ of evolution is a
geological illusion. It only looks like a burst because a lot of time—say, 5 million years—passed
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between the times when the two rock layers were laid down. In this period of time, species 3
gradually diverged from ancestor 1 through a series of transitional forms, but these
transitional forms were not preserved.
Punctuated equilibrium - a large amount of change in a short time tied to a speciation event:
Species 2 and 3 are only 100,000 years younger than ancestor 1, and all the evolutionary change
connecting them took place in this short time. The ―burst‖ of evolution is really a burst.
Transitional forms between ancestor 1 and species 3 did exist, but for such a short amount of
time that they were not preserved in the fossil record.
Macromutation - a big mutation produces sudden evolutionary change skipping over transitional
forms: The ―burst‖ of evolution is really a burst - there was a lot of evolutionary change in a
very short amount of time. Species 3 was produced by a mutation that radically changed the
offspring of ancestor 1 in many ways. Such extreme mutants are sometimes called ―hopeful
monsters.‖ This hypothesis is consistent with the fossils; however, based on other
observations, we do not have clear evidence that such extreme yet adaptive mutations generally
occur. Nevertheless, it is possible that mutations affecting development have far-reaching
phenotypic effects and have played an important role in the evolution of life.
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Punctuated equilibrium is an important but often-misinterpreted model of how evolutionary
change happens. Punctuated equilibrium does not imply that evolution only happens in rapid
bursts. Punctuated equilibrium predicts that a lot of evolutionary change takes place in short
periods of time tied to speciation events. Here's an example of how the model works:
1. Stasis: A population of molluscs is experiencing stasis, living, dying, and getting fossilized
every few hundred thousand years. Little observable evolution seems to be occurring judging
from these fossils.
2. Isolation: A drop in sea level forms a lake and isolates a small number of molluscs from the
rest of the population.
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3. Strong selection and rapid change: The small, isolated population experiences strong
selection and rapid change because of the novel environment and small population size: The
environment in the newly formed lake exerts new selection pressures on the isolated molluscs.
Also, their small population size means that genetic drift influences their evolution. The
isolated population undergoes rapid evolutionary change. This is based on the model of
peripatric speciation.
4. No preservation: No fossils representing transitional forms are preserved because of their
relatively small population size, the rapid pace of change, and their isolated location.
5. Reintroduction: Sea levels rise, reuniting the isolated molluscs with their sister lineage.
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1. Expansion and stasis: The isolated population expands into its past range. Larger population
size and a stable environment make evolutionary change less likely.
The formerly isolated branch of the mollusc lineage may out-compete their ancestral
population, causing it to go extinct.
7. Preservation: Larger population size and a larger range move us back to step 1: stasis
with occasional fossil preservation.
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This process would produce the following pattern in the fossil record: Evolution appears to
happen in sharp jumps associated with speciation events.
We observe similar patterns in the fossil records of many organisms. For example, the fossil
records of certain foraminiferans (single-celled protists with shells) are consistent with a
punctuated pattern. However, it is also important to note that we observe examples of gradual,
non-punctuated, evolution in the fossil record too. The question that needs answering is: what
are the relative frequencies of punctuated and gradual change?
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GLOBAL CAUSES OF CLIMATE CHANGE THROUGH GEOLOGICAL TIME
During the history of the Earth, the climate has swung to hot and cold extremes compared with
today‘s climate. Below is a graph that shows science's best guess at how temperature has
varied over the past 540 million years during the Phanerozoic Era.
Variation in average global temperature during the past 540 Ma (the Phanerozoic)
The key feature is the wide variation in temperature over time, sometimes colder than the
average 15°C of today but much of the time considerably warmer. Note the three places where
temperatures drop below the line, representing periods of ice house conditions. Even though
there have been several extensive ice ages during the Phanerozoic (Ordovician, Carboniferous &
Quaternary Ice Ages), for the majority of the past half billion years there have been no
permanent ice caps in either hemisphere. In that sense, the total melting of the Greenland and
Antarctic glacial ice sheets that is predicted to happen due to global warming today, would
mark a return to historically normal conditions for our planet.
MILANKOVITCH CYCLES
It is now apparent that the climatic fluctuations of the past 500 million years or so have
followed a series of distinctive patterns, and hence explanations of long-term climatic change,
in recent years, tended to focus on the factors that have given rise to both the regularity and
frequency of climatic fluctuations. The hypothesis that has attracted the greatest attention is
undoubtedly the ―Astronomical Theory‖ developed by Croll over 100 years ago and subsequently
elaborated by the Serbian geophysicist Milankovitch in the 1940s. The Milankovitch theory is
now the name given to this astronomical theory of climate variations. The theory is based on
the assumption that surface temperatures of the Earth would vary in response to regular and
predictable changes in the Earth‘s orbit and axis. Due to planetary gravitational influences, the
shape of the Earth‘s orbit is known to change over a period of approximately 100 ka from
almost circular to elliptical and back again, a process referred to as the eccentricity of the
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orbit.
In addition, the tilt of the Earth‘s axis varies from ~21.5° to ~24.5° and back, over the space
of ~41 ka. Because the angle of tilt is measured relative to an imaginary line representing the
plane of the ecliptic (the plane described by the Earth‘s elliptical path around the sun), this
phenomenon is known as obliquity of the ecliptic. Obliquity does not influence the total amount
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of solar radiation received by the Earth, but affects the distribution of insolation in space and
time. As obliquity increases, so does the amount of solar radiation received at high latitudes in
summer, whilst insolation decreases in winter. Changes in obliquity have little effect at low
latitudes, since the strength of the effect decrease towards the equator. Consequently,
variations in the Earth's axial tilt affect the strength of the latitudinal temperature gradient.
Increased tilt has the effect of raising the annual receipt of solar energy at high latitudes,
with a consequent reduction in the latitudinal temperature gradient.
The third variable arises because the gravitational pull exerted by the sun and the moon causes
the Earth to wobble on its axis like a spinning top, and is known as precession. The consequence
of this is that the seasons (or equinoxes) seem to move around the sun in a regular fashion,
hence the term precession of the equinoxes or precession of the solstices. In effect this
means that the season during which the Earth is nearest to the sun (perihelion) varies. At
present, the Northern Hemisphere winter occurs in perihelion, while the summer occurs at the
furthest point on the orbit (aphelion). In about 10,000 years‘ time, the position will be
reversed, while in approximately 21,000 years‘ time the cycle will be complete. In fact, it now
appears that there are two separate interlocking cycles, a major one averaging around 23 ka
and a minor one at approximately 19 ka. Like obliquity, precession does not affect the total
amount of solar energy received by the Earth, but only its hemispheric distribution over time.
If the perihelion occurs in mid-June i.e. when the Northern Hemisphere is tilted toward the
Sun, then the receipt of summer solar radiation in Northern Hemisphere will increase.
Conversely, if the perihelion occurs in December, the Northern Hemisphere will receive more
solar radiation in winter. It should be clear that the direction of changes in solar radiation
receipt at the Earth's surface is opposite in each hemisphere.
These three variables, in combination, exert a profound effect on global temperatures. The
total amount of radiation received is determined largely by the eccentricity of the Earth‘s
orbit, while the other astronomical variables affect the way in which that heat energy is
distributed at different latitudes. In general it seems that solar radiation receipt in the low
and middle latitude regions is governed mainly by precession and eccentricity variations, while in
higher latitudes the effects of eccentricity are modulated or amplified by changes in obliquity.
Patterns of change through time can be calculated from astronomical data and Milankovitch was
therefore able to obtain estimates for radiation inputs at different latitudes, and hence infer
temperature changes through time.
The theory was first published in 1924 and initially found favour with many European geologists,
however, by the mid-1950s, the Milankovitch hypothesis as an explanation of climatic had been
almost universally rejected. In the late 1960s and early 1970s, however, work initially on sealevel changes and subsequently on deep-ocean sediments reawakened interest in the
Milankovitch hypothesis. The data from oxygen isotope variations in marine microfossils
provided the first unequivocal evidence of the 100 ka eccentricity, the 41 ka obliquity cycle and
the 23 ka and 19 ka precessional cycles in the geological record, and were an impressive
demonstration of the role of the astronomical variables in determining patterns of long-term
climatic change. Subsequently, evidence of the influence of the astronomical variables has been
detected in a wide range of proxy records including coral reef sequences, pollen records and ice
cores. Collectively these data would seem to confirm the hypothesis that changes in the Earth‘s
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orbit and axis, which have become known as orbital forcing, are the primary driving mechanism
in Quaternary climatic change.
Although the Astronomical Theory offers a coherent explanation for the sequence of major
Quaternary climatic oscillations over the last 2 million years, the sequence of changes shown on
the graph on page 140 cannot be accounted for solely by the Milankovitch model of orbital
forcing. It is now apparent that other factors have influenced the course of global climatic
change during the Phanerozoic Era. The major additional elements in the climatic equation which
serve to modulate or amplify the effects of the astronomical variables appear to be changes in
the distribution of the continental landmasses, oceanic circulation, tectonic activity, mountain
belt uplift and, possibly, variations in greenhouse gases from natural processes.
CONTINENTAL DISTRIBUTION
Over the history of the Earth there have been five, or possibly six, ice ages, three of which
were in the last 540 Ma (Ordovician, Carboniferous & Quaternary Ice Ages). They occur at
intervals of hundreds of millions of years. Plate tectonics acts on a similar timescale, and so
would seem an obvious candidate for a major forcing factor implicated in the coming and going
of ice ages. Plate motions are most unlikely to be the only factor, however, as they alone cannot
bring about climate change. What they can do though, is change the distribution and relative
positions of continents and oceans in ways that enable other influences to come into play. The
ways in which the distribution of land and sea might affect climate can be explored by looking
at models that use greatly simplified situations to simulate conditions that might have been
obtained in the past, when distributions of continents and oceans were different.
Although contrasting global climates in the geological past appear to be associated with
different continental distributions, the only way the relationship between configuration and
climate can be quantitatively investigated is through the use of computer simulations. In the
diagrams on the next page, idealised extremes of continent-ocean distribution have been used
to investigate the effects of continental configurations on climate by means of a model
consisting of a set of simple, connected balances of heat and moisture in different latitude
bands. The model simulates some of the familiar features of the Earth‘s climate system
including seasonal and latitudinal variations of incoming solar radiation, surface temperatures,
cloud cover, precipitation and evaporation, and snow and ice cover. The model was run using two
idealised continental geometries based on the present-day land area:


Two ―caps of land extending from the poles to 45° of latitude (both with and without ice
caps.
A tropical ―ring‖ of land extending 17° north and south of the Equator.
These idealised continental distributions are nothing like those of today, but the geography of
the Earth may have approximated a tropical ―ring world‖ between 700 Ma and 600 Ma, and
could have approximated very roughly a ―cap world‖ with one polar ice cap in the late
Carboniferous at 300 Ma.
As you might expect, for a polar cap world with ice caps the variation along the line of latitude
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in temperature looks rather similar to a simulation of the temperature distribution for the
present-day Southern Hemisphere and both of the cap worlds give average equatorial
temperatures close to those of the present day. By contrast, the tropical ring world is
significantly warmer than either of the cap worlds. In terms of radiation balance the addition
of ice caps leads to a cooler world. This is because ice has a much higher albedo than exposed
or vegetated continental crust, so an ice cap world reflects more solar radiation back into space
than an ice-free cap world.
As demonstrated so far, computer models can be useful in helping to isolate the effects of
different factors (in this case, changes in albedo and moisture balance), while keeping other
variables (e.g. the total continental area, and total incoming solar radiation) constant. They can
also be used to combine the effects of different factors in a quantitative way, and to test
hypotheses against physical laws. Climate models are based on a myriad of approximations and,
while they can often demonstrate that certain events may possibly have occurred, it is much
more difficult to attempt to prove that they must certainly have occurred.
In this case, two important components of the climate system not incorporated into the model
are: Heat transport by surface ocean currents and by deep thermohaline circulation; and
changes in the concentration of greenhouse gases in the atmosphere.
Cap, ring and slice worlds are very simplistic representations of the Earth, but do help to show
how continental configurations might influence ocean current patterns and hence climate. Cap
world, for example, would affect the surface current patterns and hence climate. This is due to
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the fact that ocean currents carrying heat form lower latitudes would be unable to penetrate
to very high latitudes, with the result that there would be a
―Slice world‖
stronger temperature contrast between equatorial and Polar
Regions. In cap world, the Polar Regions would be thermally
isolated, making the development of polar ice caps more likely:
strong eastward currents, comparable with today‘s Antarctic
Circumpolar Current, could flow around the polar continents under
the influence of westerly winds, further isolating the polar
continents from warm currents flowing from low latitudes.
In today‘s world, the northern Polar Regions are not similarly isolated, but they are
nevertheless largely cut off from warm currents flowing from lower latitudes. The Arctic
Ocean is almost completely surrounded by land, with the result that the only warm water
penetrating the region is the North Atlantic Drift (the downstream extension of the Gulf
Stream), which flows northwards through the Norwegian and Greenland Seas. By contrast, in a
tropical ring world, ocean circulation between tropical and Polar Regions would be possible, so
heat could be transferred from mid-latitudes to high latitudes, resulting in lower latitudinal
temperature contrasts.
Of course, discussing the transport of heat around the Earth by means of currents alone is
unrealistic as the redistribution of heat by winds, including the effects of evaporation,
transport and condensation of water vapour, is being ignored. Nevertheless, the high specific
capacity of water (the amount of heat needed to raise the temperature of 1kg of water by 1°C)
means that heat transport in the ocean is extremely important influence on climate. Linked to
this is the effect of gateways on the pattern of surface and deep ocean currents.
Gateways are gaps between the continents that permit significant longitudinal or latitudinal
connections to be made between oceans. Their opening and closing, and the resulting changes in
heat transport, can cause the climate for a particular land mass to change much more rapidly
than the slow drift of the continent across climatic belts. An example of how the opening of a
seaway can affect global climate is the opening of the Drake Passage between the southern tip
of South America and Antarctica.
The climate of modern Antarctica is extreme. Located over the South Pole and in total
darkness for six months of the year, the continent is covered by glacial ice to depths in excess
of 3 km in place. Yet this has not always been the case. About 50 Ma ago, even though
Antarctica was in more or less the same position over the pole, the climate was much more
temperate – there were no glaciers and the continent was covered with lush vegetation and
forests. So how did this extreme change come about?
The modern climate of Antarctica depends on its complete isolation from the rest of the planet
as a consequence of the Antarctic Circumpolar Current that completely encircles Antarctica
and gives rise to the stormy region of the Southern Ocean known as the roaring forties. The
onset of this current is related to the opening of seaways between obstructing continents.
Antarctica and South America were once joined together as part of Gondwana and were the
last parts of this original supercontinent to separate. By reconstructing continental positions
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from magnetic and other features of the sea floor in this region, geologists have shown that
the Drake Passage opened in three phases between 50 Ma and 20 Ma. At 50 Ma there was
possibly a shallow seaway between Antarctica and South America, but both continents were
moving together. At 34 Ma the seaway was still narrow, but differential movement between the
Antarctic and South America Plates created a deeper channel between the two continents that
began to allow deep ocean water to circulate around the continent. Finally, at 20 Ma there was a
major shift in local plate boundaries that allowed the rapid development of a deep-water
channel between the two continental masses.
While these plate motions were taking place the effect on Antarctica was profound. By 34 Ma
the climate cooled from the temperate conditions that previously existed. This was sufficient
for glaciers to begin their advance, and was followed by a period of continued cooling until, at
about 20 Ma, glaciation was complete. The Drake Passage had both a local and a global effect,
initially cooling the climate of Antarctica from temperate to cold and, then, ultimately playing
an important role in the change from global ―greenhouse‖ conditions 50 Ma ago to the global
―icehouse‖ of today.
This example shows how plate tectonics, continental drift and the opening and closing of
seaways can have a profound influence on both local and global climate. Throughout the
Phanerozoic there were long periods when the Earth was much warmer than today – often called
a ―greenhouse‖ climate – and other times when it was cold – called an ―icehouse‖ climate. These
cycles occur over periods of 100 Ma, reflecting the timescale of plate movements and the
growth and destruction of oceans. Given the clear link between ocean circulation and climate,
and the similar timescales of global climate change and plate motions, it is inescapable that one
of the chief controls on long-term changes in the global climate must be plate tectonics.
MOUNTAIN BELTS
Another aspect of how plate tectonics can influence climate is through the process of
orogenesis (mountain building). The Tibetan Plateau is the highest and largest plateau on the
Earth‘s surface. Its southern edge is marked by the Himalayas and several of Asia‘s other
great mountain ranges – the Karakoram, Hindu Kush and Kunlun – decorate its western and
northern margins. If continental relief can affect global climate, this is the obvious region to
study. The elevation of the Tibetan Plateau is a result of a head-on collision between the
continental margins of two plates: India, which was migrating northwards, and Eurasia, which
was stationary. The collision, which has been dated at around 50 Ma, led to crustal thickening
and uplift of the Himalayan Mountains and the plateau. The area then reached its maximum
present elevation between about 15 to 10 Ma ago. Mountain ranges and plateaux, such as the
Himalaya and the Tibetan Plateau, affect the climate physically by redirecting air masses
around and/or over them. Moisture-laden winds release their precipitation on the windward
side, whereas the leeward side is dry. The rise of the Himalaya and Tibet is believed to have
intensified the strength of the southwest monsoon by providing a source of heat at a critical
position in the atmospheric circulation.
To a large extent, the climate of southern Asia is dominated by the monsoons. In the northern
winter, the region is affected by cold, dry north-easterly winds blowing out from the intense
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high pressure region over the Eurasian continent. By contrast, in the northern summer when
there is a strong low pressure region over the Eurasian continent, the area receives moistureladen air in the southwest monsoon. The whole of southern Asia benefits from the monsoon
rains. In particular, the areas of tropical rainforest along southwest India, Myanmar (formerly
known as Burma) and Sri Lanka are the results of the south-westerlies releasing much of their
moisture over the high land. On approaching the southern slopes of the Himalaya, the stillmoist air mass is driven further upwards, causing it to cool, condense and precipitate as summer
rainfall over northern India. As a result, annual rainfall in Darjeeling on the southern slopes of
the Himalaya is over 3000mm, of which 88% falls between June and September, i.e. during the
south west monsoon. Less than 100km away, to the north of the Himalayan Mountains, the
Tibetan town of Gyantse receives an annual rainfall of only 270mm.
The seasonal shift in the Intertropical Convergence Zone (ITCZ) means that seasonally
changing winds (monsoons) affect large parts of the globe at low latitudes. The extreme change
in pressure over a large part of central southern Eurasia, from intense high pressure in winter
to very low pressure in summer, means that seasonal changes in the vicinity of southern Asia
and the Arabian Sea are by far the most dramatic. It seems that the reason for this may lie in
the existence of the Tibetan Plateau.
In 1989, the results of a series of experiments were published using a sophisticated computer
model of the global climate (a general circulation model or GCM) designed to investigate the
effect on climate of such an extensive high-altitude plateau. Starting with a simulation of the
present-day climate, the researchers changed just one variable: the topography of present land
masses. When the Tibetan Plateau was ―removed‖, the heavy summer rainfall in northern India
all but disappeared. In contrast, an even larger and higher plateau in central Asia greatly
increased the area of summer monsoon rainfall throughout extensive regions south of the
plateau, and decreased summer precipitation much further west in the Mediterranean region in
Europe.
Of course, such experiments have their critics, many of whom emphasised that no model can
take into account all the possible variables. After all, if you think about climate change over the
past 60 Ma, uplift of the Tibetan Plateau is not the only change to have taken place within the
Earth system. Nevertheless, the results of the modelling clearly suggested that the uplift of
Tibet could have had a dramatic effect on atmospheric circulation and precipitation throughout
much of the Northern Hemisphere, and may well have affected the strength of the monsoon
winds, particularly those of the southwest monsoon over southern Asia.
The strength of the southwest monsoon is determined by the pressure difference between the
high over the tropical Indian Ocean and the low over the southern part of the continent. At the
end of winter, the large rocky mass of the Tibetan Plateau heats up fast, once its high-albedo
covering of snow has melted. This heat warms the overlying air, allowing it to rise, causing the
pressure over the continent to fall. The warming of the air over the Himalaya and Tibetan
Plateau has a particularly large effect because the air is thin at these high altitudes and its
temperature, therefore, is more sensitive to changes in heat. The monsoon rains indirectly help
to warm the air over Tibet. The condensation of moisture to form rain as air rises over the
southern Himalaya releases latent heat originally taken up from the ocean as latent heat of
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evaporation. Thus, air that is rising because it has been warmed by the underlying continent has
an extra heat source, and rises more vigorously, intensifying the low pressure over the region
and drawing in moist air from the south even more strongly.
The overall effect of the Tibetan Plateau and the Himalaya on atmospheric circulation,
therefore, is determined by both their high elevation and their geographical position. Indeed,
the Tibetan Plateau is so high that the subtropical jet stream passes either to the north or to
the south of it. The climate implications of this are not well understood, but it is clear that the
jet stream would not have been diverted in this way before the uplift of Tibet. Uplift must
have caused major changes in atmospheric circulation across the northern Hemisphere. As far
as the southwest monsoon is concerned, because summer heating of the atmosphere over Tibet
has increased as the plateau has risen, it is possible that, at some stage during its elevation, a
threshold was reached above which the monsoon winds were greatly strengthened.
Although it is not possible to be precise about the timing of either the uplift of the Tibetan
Plateau or climate change in southern Asia, there is evidence that summer rainfall associated
with the southwest monsoon increased between 9 Ma and 6 Ma. By this time, the southern
plateau had already stabilised at its present elevation for at least six million years, but
northern regions were just reaching their maximum elevation. Geologists have speculated that,
by about 9 Ma, the area of the elevated plateau was sufficient to heat the lower atmosphere
during the summer months and thus trigger a dramatic increase in the intensity of the monsoon.
So far, a possible link between uplift of the Tibetan Plateau and a strengthening of the
southwest monsoon has been discussed. Although this may have affected a large part of the
globe, it essentially had a regional effect on the climate. If the uplift of this (or any other)
high plateau had had an effect on global climate, you would have expected the change to have
taken place not over the past 10 Ma as the monsoon was strengthening, but over the past 50
Ma, i.e. from the beginning of uplift of the plateau, initiated by continental collision. Before
looking at other possible factors which could cause global climate change, it is important to
understand how global climate has changed since the collision between India and Eurasia.
Bottom-water temperature °C
20
10
Quaternary
0
120
168
100
80
60
Ma before present
40
20
0
The variation in global average temperature over the past 120 Ma, deduced from oxygen
isotope studies of the remains of deep-sea benthic foraminiferans is shown in the graph over
the page. Despite the fluctuations, there is a clear downward trend, with a net cooling of
nearly 20°C over the period concerned. What caused this decrease in global temperature? One
potential cause may have been the arrangement of the continents (as already discussed).
However, irrespective of how the continents are rearranged, computer-based climate models
cannot reproduce the long-term cooling pattern indicated by the graph below. Furthermore, the
climate-modelling experiments that investigated the effects of the Tibetan Plateau on global
climate failed to show that the uplifting of Tibet could, by itself, cause long-term cooling. It
seems that something else is required to explain this trend. A change in the composition of the
atmosphere, especially greenhouse gases such as carbon dioxide, would be another factor that
could cause global climate change.
CHANGES IN GREENHOUSE GASES FROM NATURAL PROCESSES
The Earth‘s atmosphere is made up of a number of naturally occurring gases. The table below
lists the gases in the atmosphere in order of abundance. By far the most abundant gases are
nitrogen and oxygen, which together make up ~99% of the total. Apart from nitrogen and
argon, all of these gases affect the Earth‘s climate through their interaction with incoming
radiation, outgoing long-wave radiation, or both.
Gas
Nitrogen (N2)
Oxygen (O2)
Argon (Ar)
Water vapour (H2O)
Carbon dioxide (CO2)
Ozone (O3)
Methane (CH4)
Nitrous oxide (N2O)
Approximate concentration by volume
78%
21%
1%
0.3% (3000 parts per million)
0.038% (380 parts per million)
0.01%
1.8 ppm
0.3 ppm
As incoming solar radiation passes into the Earth‘s atmosphere most of the very short
wavelength radiation (ultraviolet light) is absorbed by the ozone. As solar radiation passes down
through the lower layers of the atmosphere, it is further absorbed or reflected. Of the total
solar radiation arriving at the outer edge of the atmosphere, about 25% is reflected from
clouds and particles in the atmosphere and about 25% is absorbed by clouds and atmospheric
gases. The 50% of the incoming radiation which reaches the ground directly is called the
insolation. Most of this is in the visible spectrum. About 5% is reflected back into space and
the remainder is absorbed by the land and ocean surfaces.
When an object absorbs solar radiation it becomes warmer. It then begins to radiate heat
energy to its surroundings. This radiated heat energy is in the form of long-wave radiation
(infra-red light). It is invisible to the human eye, though it can be felt by the body as heat. It
is this infra-red radiation from the land and ocean surfaces which takes part in what is known
as the ―greenhouse effect‖.
169
The greenhouse effect occurs because of differences in the way the gases in the Earth‘s
atmosphere transmit and absorb visible and infra-red radiation. The main constituents of the
Earth‘s atmosphere are nitrogen and oxygen; both are almost completely transparent to both
visible and infra-red radiation. But within the remaining 1%, there are several gases which are
also transparent to the visible part of the spectrum but which absorb infra-red. These are
called ―greenhouse gases‖; the most important is water vapour; carbon dioxide comes next;
other greenhouse gases include methane, nitrous oxide and CFCs (human-made gases).
The greenhouse effect is so-called because of its similarity to what happens in a greenhouse.
Sunlight coming in through the glass roof and walls heats up the interior of the greenhouse and
the glass prevents the excess heat escaping. Glass behaves like the greenhouse gases in that it
transmits visible radiation and absorbs infra-red. But, in fact, most of the heat-trapping effect
in a greenhouse is simply because the glass prevents convection and droughts from dissipating
the heated air.
As a result of the greenhouse gases in the Earth‘s atmosphere some of the infra-red radiation
which would otherwise escape into outer space is trapped and re-radiated back down to the
Earth. The effect is to raise the global average temperature by about 33°C over what it would
otherwise be. Without the greenhouse effect, the Earth would have an average temperature of
about -18°C compared with its actual average of 15°C, and would be far too cold for any known
form of life to survive. The natural greenhouse effect is, therefore, entirely beneficial.
The temperature variations shown in the graph on page 166 could have resulted from a
decrease in the concentration of carbon dioxide in the atmosphere. To account for the overall
fall in temperature, therefore, the carbon dioxide concentration would have to have declined
from a value of about eight times that of the present day. Many geologists believe that the
building of large mountain belts in central Asia could well have played a role in bringing about
such a dramatic change in atmospheric composition. This particular aspect of mountain building
is thought by some to be an important mechanism for changing atmospheric CO2 concentrations,
and hence climate.
In climate models proposed in recent years, the role of mountain building in the global climate
system has been treated in two contrasting ways. These approaches differ in their assumptions
concerning what primarily determines the CO2 concentration of the atmosphere, and hence
what ultimately drives carbon changes between the various carbon reservoirs, namely the
atmosphere, the oceans and the Earth‘s crust.
The GEOCARB model (originally conceived by Bob Berner and his colleagues at Yale University)
rests on two important assumptions: 1. Global temperatures are determined by the
concentration of CO2 in the atmosphere. 2. The concentration of CO2 in the atmosphere is
determined primarily by the volume of gases emitted from volcanoes.
An approximate measure of the global rate of emission of volcanic gases can be obtained from
the rate of production of new sea floor. There is evidence to suggest that over the past 110
Ma, the production rate of oceanic crust has decreased. Therefore, over the past 110 Ma, less
CO2 has been supplied by volcanism, so, according to the GEOCARB model, there would have
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been global cooling as a result of lower concentrations of CO2 in the atmosphere.
The second approach, referred to as the mountain-forcing model, is preferred by scientists
who suggest that changes in atmospheric CO2 concentrations are driven by changes in chemical
weathering rates. This model was pioneered by two geoscientists from the USA, Bill Ruddiman
and Maureen Raymo, who extended the idea that uplift of the Tibetan Plateau strengthened
the southwest monsoon and suggested that high weathering rates over the region of steep
topography affected by high summer rainfall were at least partly responsible for the global
cooling that followed the collision between India and Eurasia. If this is true, then uplift of
Tibet effectively set the scene for the glacial periods that have characterised the climate in
recent geological time.
It is important to emphasise that the Tibetan Plateau and the Himalaya appear to play quite
different roles in the climate system. The strengthening of the monsoon seems to be the result
of a large high plateau (the Tibetan Plateau) being uplifted, as it was this that affected the
atmospheric circulation. In contrast, the link between changes in atmospheric CO 2
concentrations and weathering requires increases weathering rates, and it is over the Himalaya
that these rates become particularly high and not on the Tibetan Plateau, which experiences
little rainfall (much of which evaporates) and has relatively modest relief.
Ironically, the very fact that the uplift of Tibet and the Himalaya appears to have such a
strong effect on the levels of atmospheric CO2 could potentially be a problem for supporters of
the mountain-forcing model, because it seems to leave open the possibility of runaway cooling
occurring, for which there is no evidence in the relevant geological records despite the overall
decline in global temperature during the past 100 Ma. According to the GEOCARB model, levels
of atmospheric CO2 would be maintained at more or less the same level over time by a negative
feedback process as follows: if more CO2 is released into the atmosphere as a result of
increased rates of production of sea floor, global temperatures will rise. This temperature rise
would lead to increasing weathering rates, which in turn would remove CO 2 from the
atmosphere, thus decreasing and returning the temperature towards its original value; and so
on (see flow diagram below).
Atmospheric CO2
concentration
increases
Chemical weathering
rates decrease
Volcanism; production
of new sea floor
Average global
temperature
increases
Average global
temperature
decreases
Chemical weathering
rates increase
Atmospheric CO2
concentration
decreases
Negative feedback loop in the climate cycle derived from the
GEOCARB model.
171
On the other hand, if (as assumed by the mountain-forcing model) the concentration of CO2 in
the atmosphere is primarily controlled by mountain uplift, then there is no direct link between
the rate of operation of the CO2 sink (i.e. preservation and burial of carbon in the oceans) and
the rate of operation of the CO2 source (i.e. volcanic emission). As a result, there can be no
direct stabilising feedback loops of the type described above to prevent runaway cooling.
Estimates of the rate at which the weathering of the Himalaya has depleted CO 2 in the
atmosphere according to the mountain-forcing model show that weathering alone would exhaust
all the CO2 in the atmosphere in only a few million years – something that obviously has not
happened. It is important to remember, however, that the processes that eventually lead to
mountain building can, at the same time, provide a source of atmospheric CO2.
Throughout the period of subduction, deep-sea sediments will have been scraped off the
surface of the subducting oceanic plate, building up great wedges of sediment at the side of
the subduction zone trench. As the continents collide, these sediments (including silica and
calcareous remains of dead organisms) will be subjected to metamorphism due to increased
temperatures and pressures. If calcium carbonate and silica are heated together at about
400°C, a decarbonation reaction occurs that releases CO2.
A further source of CO2 associated with the collision of continents is the oxidation of organic
carbon in the crust. After collision of India and Eurasia, buried carbon-rich sediments were
uplifted and eroded. Once exposed at the surface, this organic carbon was oxidised to CO 2
through bacterial activity and by simple chemical reactions with atmospheric oxygen.
In summary, there are at least two mechanisms that would have counteracted any tendency
towards a runaway loss of CO2 from the atmosphere through the weathering of the Himalaya:
decarbonation reactions in rocks heated up as a result of continental collision within the crust,
and oxidation of organic carbon at the surface of the crust.
So, which is right: the GEOCARB model or the mountain-forcing model? Well the answer is
both! The GEOCARB model for the carbon cycle is a steady-state model, which means that it
assumes the system has attained equilibrium, e.g. it assumes the flux of CO2 into the
atmospheric CO2 reservoir equals the flux out of it. If such a steady-state system is
perturbed, then steady-state conditions will no longer apply and the system will adjust until the
balance of fluxes is re-established. Uplift of the Himalaya and the Tibetan Plateau can be
considered a transient event that disturbed the steady-state carbon cycle. It is believed that,
during this transient event, the chemical weathering rates rather than the rates of production
of new ocean floor determined atmospheric CO2 levels.
So the GEOCARB model provides a satisfactory mechanism for long-term climate change over
hundreds of millions of years, while the mountain-forcing model is appropriate for the relatively
short-term disturbance to global climate caused by the uplift of Tibet and the Himalaya over a
period of several million years. Although the two models have been presented as opposing each
other, more recent revisions of the original GEOCARB model now include a feedback link
between Tertiary mountain uplift and global cooling. In other words, the two models are
converging despite operating on different timescales, thus providing a good example of how
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science advances through the testing of competing hypotheses.
VOLCANOES
Major explosive eruptions blast huge amounts of volcanic ash high into the atmosphere, often
reaching the stratosphere. Awe-inspiring though these eruption columns are, it is not the ash
particles that cause the most significant climatic effects. Ash fragments are relatively large
(ranging from tiny shards to fist-sized lumps of pumice) and they fall back to the ground over
timescales ranging from seconds to days. The longest-lasting environmental effects are
produced by the volcanic gases.
Volcanoes produce large quantities of sulphur dioxide as well as carbon dioxide. After eruption,
sulphur dioxide reacts with water vapour in the atmosphere to form aerosols (i.e. tiny airborne
droplets) of sulphuric acid, H2SO4. These acid aerosols form by a complex series of
photochemical reactions that may continue for months, replenishing the aerosol cloud so that
new particles form while older, larger ones settle out of the atmosphere. Aerosol particles are
tiny so any reaching the stratosphere may remain suspended there for months or years, far
longer than solid ash fragments.
Volcanic aerosols often produce gloriously colourful sunsets (and sunrises), which can be
enjoyed around the world. During the day, the sky looks blue because gas molecules scatter
sunlight in all directions. These molecules are smaller than the wavelength of light and scatter
blue wavelengths more effectively than red. Thus, blue light from the sun actually reaches your
eye even when you are not looking directly at it, whereas the other wavelength of light, as in
aerosols and smoke, the scattering process is more complex but the dominant effect is a shift
to longer visible wavelengths, around orange to red in the visible spectrum. It is less intense
than the blue-sky effect, and is most noticeable at sunset and dawn when the sun is low in the
sky and its rays follows a much longer atmospheric path.
In detail, the physics of light scattering by aerosols is complex, but the overall effect is that
when a layer of volcanic aerosols gets between the Earth and the sun, a fraction of the sun‘s
radiation is scattered back to space. In other words, the presence of the aerosols effectively
increases the Earth‘s albedo (i.e. the amount of light reflected back to space) causing cooling.
For example, the eruption of Mt Pinatubo in the Philippines in 1991 caused an observed
Northern Hemisphere temperature change of -0.5°C.
However, due to the long response time of the Earth system, it is unlikely that such short-lived
volcanic events have had a sustained impact on global climate. A different type of eruption that
might have had more serious environmental effects than the explosive eruptions is the effusion
of huge volumes of lava in the geological past. These are termed flood basalts.
Accumulations of flood basalt lava flows cover large parts of the Earth. Some of the statistics
are impressive: India‘s 65 Ma Deccan Traps cover 0.5 million km 2, and may have covered 1.5
million km2 when first erupted. They have an average thickness of at least one kilometre. Most
of their huge volume may have erupted in less than 0.5 million years and they consist of many
hundreds of individual lava flows. What would be the impact of such massive eruptions on global
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climate?
The Cretaceous Period (145-65 Ma) was an exceptionally productive period in terms of creation
of ocean crust and eruption of massive basaltic plateaux. The Ontong-Java Plateau, in the
Pacific Ocean, erupted 122 Ma and is estimated to be 25 times bigger than the Deccan Traps.
The effects of all this volcanic activity helped increase the atmospheric CO2 level to at least
four times the modern pre-industrial level of 280 ppm, leading to global greenhouse warming of
between 3°C and 8°C. This would have been accompanied by other changes such as an increase
in sea level. A runaway greenhouse effect was counteracted by increased deposition and
preservation of carbon in the warm oceans and shallow seas.
It is important to consider whether this increase in CO2 concentration could be the direct
effect of the eruption of flood basalt or whether other sources need to be considered. Initial
estimates based on the knowledge of how mantle CO2 is actually released into the atmosphere
suggest that these volcanic increases are fairly low. For example, the estimated annual flux of
CO2 to the atmosphere from the eruption of 1000 km3 of basalt over a decade is about 3% of
the present-day natural land-atmosphere flux – too small to have a significant effect on the
CO2 concentration in the atmosphere. Such a modest flux would probably allow carbon cycle
processes to offset the volcanic addition. This will result in some global warming, however, and
could trigger other mechanisms such as the destabilisation of methane hydrate deposits and
the release of methane into the atmosphere.
At low temperatures and high pressures, gases such as methane (which is 20 times more potent
as a greenhouse gas than CO2) can combine with water to form solid, ice-like crystalline
compounds known as gas hydrates. Within the hydrate structure, water molecules are organised
into hydrogen-bonded cages that enclose the ―guest‖ gas molecules (e.g. methane), forming a
solid stable at temperatures well above the melting point of water ice. When hydrate melts, or
dissociates, liquid water and a large volume of gas – up to 164m3 for 1m3 of gas hydrate – is
produced.
In the natural environment, suitable conditions for gas hydrate formation are to be found in
sediments of the deep sea, deep-lake bottoms and subsurface permafrost. Here, low
temperatures, high pressures and plentiful water (ice) promote the growth of gas hydrates in
the presence of a suitable gas source. Generally, hydrates are found at water (or permafrost
sediment) depths of greater than 300-500 metres. The gas required for formation of gas
hydrates, such as methane, comes from the bacterial breakdown of organic matter.
The large volumes of methane trapped as gas hydrates in fairly shallow water, and its potency
as a greenhouse gas, means natural gas hydrates may play an important role in the global carbon
cycle and climate change. Isotopic analysis of carbon from the shells of fossil foraminifera
provide evidence that periods of global warming in the geological past are linked to injections of
isotopically light carbon into the atmosphere. Isotopically light carbon is characteristic of
biogenic gases, as found in seafloor gas hydrates.
The general theory is as follows. During warm climate periods, rising sea-levels should act to
stabilise seafloor gas hydrates by increasing pressure, but may destabilise continental gas
174
hydrates through flooding of low-lying permafrost areas with relatively warm seawater. This
destabilisation of continental hydrates and associated methane release could act as a positive
feedback on warming, increasing temperatures even further. The opposite is the case for
periods of global cooling, e.g. during glaciation. A reduction in sea levels would act to destabilise
seafloor gas hydrates, releasing methane, and potentially acting as a negative feedback effect,
warming the climate and restricting glaciation. In this case, continental gas hydrates would be
stabilised by lower surface temperatures and increased pressures associated with burial under
ice sheets.
The role of methane release from gas hydrates plays in global climate change is still poorly
understood. However, it is now recognised that gas hydrates represent a major source of
greenhouse gases held precariously in a relatively unstable geological environment, and that
large-scale release of methane could have a major impact on global temperatures.
EVIDENCE OF GLOBAL CLIMATE CHANGE
FOSSIL RECORD
The fossil record provides evidence of different climatic zones, including corals and land plants.
In favourable situations in tropical seas, corals together with all the organisms to which they
give shelter and attachment, grow in such profusion that they build up reefs and islands of very
considerable size. Although coral reefs are one of the most familiar and most studied of
modern carbonate environments, there are many other types of reef developing at the present
time and preserved in the geological record.
There are many factors controlling the growth of modern coral reefs and it is likely that these
same factors exerted an influence on coral and other reefs in the past. For coral-reef growth,
these factors are: a) water temperature – optimum growth is around 25°C, b) water depth –
most growth takes place within 10 metres of the surface, c) salinity – corals cannot tolerate
great fluctuations and d) turbidity and wave action – coral growth is favoured by intense wave
action and an absence of terrigenous silt and clay. The majority of reefs occur along shelf
margins, an agitated zone where waves and currents of the open sea first impinge on the sea
floor. However, there are deeper water corals forming coral banks today; these are mostly
solitary corals.
The coral species found today (Scleractinian) are not the same as those of the past (Tabulate
and Rugose); because of this some geologists have suggested that Palaeozoic reef-building
corals may have preferred cooler water, like some of the solitary species today. However,
whether corals could survive in cold waters or not: they could only flourish with continual and
plentiful supply of sunshine.
Land plants make excellent climate indicators and, broadly speaking, two approaches are
commonly adopted; nearest living relative technique and physiognomy. The first technique is
175
based on the climatic tolerances of nearest living relatives (NLRs) and is known as the NLR
technique. This assumes that ancient plants and plant communities lived under similar conditions
to those of their nearest living relatives. The success of the technique, which is widely used in
Quaternary studies, depends on the correct identification of the fossil and, of course, limited
evolutionary change. For pre-Quaternary fossils that may represent extinct groups with very
different environmental tolerances from their living relatives, and for vegetation for which
there are no living counterparts, a different approach has to be adopted.
The second method for using plants as climate indicators is based on the architecture, or
physiognomy, of the plant or community, and is applied when fossils are in the form of
vegetative organs, particularly leaves. As plants cannot move around once they have taken root,
they must be well adapted to their local environment or they will die. In the course of evolution,
most plants have become honed for successful exploitation of particular environmental niches,
and many display specially adapted physiognomies. One extreme example is seen in desert
plants, which have a low surface area to volume ratio and thick cuticles so as to conserve water
(e.g. cacti). Rainforests by contrast, are characterised by plants with large leaf-area, forming a
vertical succession of layers within forests.
As a result of physiognomic adaptations to environment being so consistent, quantitative
comparisons for determining pre-Quaternary climates can be made, within certain limits. One
of the most successful applications of this approach was devised as long ago as 1915, when two
American botanists noted that the leaves of modern woody, broadleaved flowing plants (such as
alder, willow and figs) tended to have smooth (entire) leaf margins in warm climates, but
toothed, jagged margins in cool climates.
SEDIMENTARY ROCKS
The distribution patterns of climatically sensitive deposits, such as evaporates, desert
sandstones (red beds), glacial diamictites (non-sorted mixtures of rock interpreted as being of
glacial origin, previously known as tillites), fossilised wood and coals can provide important clues
to past climates.
The formation of evaporates (salt deposits) requires that evaporation exceeds precipitation.
Ideal conditions occur in arid regions, where formation can take place in enclosed basins with
high temperatures and low rainfall. Modern evaporates occur mostly in subtropical regions
centred around latitudes 25° to 35° in both hemispheres. Much of the same pattern can also be
seen for desert sand dune deposits.
Red beds are sedimentary rocks containing haematite (Fe 2O3) that formed under oxidising
conditions in hot climates. The original source of the iron in these rocks was often exposed
igneous or metamorphic rocks that had been intensely chemically weathered. This iron was then
remobilised (as Fe2+) in anoxic groundwater and re-precipitated in desert sediments as iron
oxide, as a result of evaporation drawing the water up towards the surface. Modern red beds
form largely within 30° of the Equator, and most Palaeozoic red beds seem to have had a similar
distribution, being commonly associated with evaporate deposits.
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In contrast to these indicators of climatic warmth, most diamictites are considered to have
been deposited by glaciers. Their widespread occurrence in southern areas of Pangaea during
the Permo-Carboniferous indicates that large areas experienced glacial conditions for at least
some of the time. On the deep-ocean floors, sediments have been accumulating in a relatively
undisturbed manner for thousands, or even millions of years. They consist partly of terrigenous
deposits, i.e. detrital material derived from erosion of the land masses surrounding the ocean
basins, and partly of biogenic sediments composed largely of accumulations of the calcareous
and siliceous skeletal remains of micro-organisms that formerly lived in the ocean waters.
Terrigenous detritus arrives on the ocean floor by a number of different pathways, but the
principal transporting agencies are turbidity currents, bottom currents, wind and ice. In the
mid and high latitudes, both coarse and fine terrigenous detritus appears to have been
delivered to the ocean floors mainly during glacial periods, reflecting in particular the icerafting of glacially eroded debris and, to a lesser extent, the transport of aeolian (wind-blown)
sediments from the greatly expanded periglacial regions. Indeed, wind-blown sediment may
have constituted a major proportion of the fine detrital input in the low latitudes during glacial
times. In many ocean sediment sequences, therefore, a broad correlation can be detected
between the deposition of terrigenous material and former glacial episodes, and this is
reflected most clearly in the large volumes of ice-rafted debris that are found in deep-ocean
sediments.
In the deeper oceans, the sediments tend to be finer grained and are often dominated by
biogenic material consisting of the accumulations of the carbonaceous and siliceous remains of
micro-organisms that formerly lived in the ocean waters. Such sediments are known as marine
oozes and are frequently characteristic of interglacial or warmer episodes. They contain
recognisable fossil remains which provide a record of ocean circulation, ocean water
temperature and, by implication, atmospheric temperatures throughout geological time.
OXYGEN ISOTOPE RATIOS IN FOSSIL SHELLS
Oxygen has three stable isotopes with relative atomic masses 16, 17 and 18 ( 16O, 17O and 18O).
Over 99% of natural oxygen is made up of 16O, with most of the balance being 18O. Water that
evaporates from the ocean eventually condenses as cloud and falls as rain or snow. When
seawater evaporates from the ocean, water molecules with the lighter oxygen isotope (H216O)
evaporates more readily, so atmospheric water vapour is relatively enriched in the lighter
isotope. When water vapour condenses and is precipitated back into the ocean, the water
containing the heavier isotope (H218O) condenses preferentially. Both processes deplete water
vapour in the atmosphere in H218O relative to H216O. When 18O-depleted water vapour is
precipitated as snow in Polar Regions, the snow will also be depleted in 18O relative to the
oceans. The larger the ice caps, the higher the relative proportion of 18O in seawater and the
lower the relative proportion of 18O in ice caps. The diagram on the next page summarises the
relative proportions of 16O and 18O during a cold phase in global climate.
The variation in the isotopic composition of ocean waters over time can be reconstructed from
the 18O values of carbonate shells and skeletons preserved in deep-sea sediments. Many marine
organisms secrete (or build) carbonate structures and oxygen is abstracted from the sea
177
waters for this purpose. Thus, the oxygen isotope ratios in fossil carbonates buried in
sediments on the ocean floors should reflect the ratios prevailing in the oceans at the time of
their secretion. Analyses have been carried out on the remains of a range of marine microorganisms, but by far the most widely used fossils are the tests of the planktonic and benthic
foraminiferans.
The amount of 18O in the shells of
these organisms is very small, but it
can be measured accurately by mass
spectrometry.
The
higher the
measured 18O value in marine fossils,
the greater the enrichment of 18O in
seawater and the larger the ice caps
on land at the time the organisms
were
alive.
Therefore,
useful
information about past climate can
be
obtained
from
these
foraminiferans; their 18O values
reflect changes in ice volume and
global surface temperature and, of
course, when global temperatures are lower, ice caps are larger. The graph below shows the
variations in surface water temperatures in the vicinity of Antarctica during the past 60 Ma, as
determined from oxygen isotope ratios of the remains of planktonic foraminiferans.
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ROCK DEFORMATION
Rocks deform due to forces put on them. The force applied to a rock is known as stress. Stress
is defined as force per unit area, i.e. stress becomes greater if the force gets bigger, or if the
area over which a given force acts gets smaller. For example, dancers do much more damage to
wooden floors with pointed stiletto heels than through wearing broad flat-heeled shoes! When
stresses that are not equal in all directions act upon a rock, they will deform it. Such stresses
generally produce a change in size or shape. The name strain is given to these changes. So
stress produces strain! Materials can respond to stress (i.e. deform) in three different ways:
Brittle deformation materials fracture (break) under stress e.g. faulting in rocks. Plastic
deformation materials bend under stress, but do not return to original shape e.g. folding in
rocks. Elastic deformation materials deform under stress, but recover their original shape
after the force has been taken away e.g. elastic band. Rock structures cannot represent elastic
deformation, because the deformation we see in folds and faults is permanent. For example, a
spring balance stretches when weights are loaded onto it. The amount that the spring stretches
(strain) is proportional to the load (stress) put on it. This sort of strain is recoverable. This
means that the material returns to its original shape after the removal of the stress. Have a
look at a simple stress-strain curve to help you understand these ideas.
In the diagram above there are two simple stress-strain curves showing how rocks react to
stress. It shows that two rocks can respond quite differently when subjected to similar stress.
Rock A has gone beyond its elastic limit and has been permanently deformed in a ductile
fashion (plastic deformation). Rock B is a much more brittle rock and fractures (breaks) at
the elastic limit.
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CAUSES OF ROCK DEFORMATION
Rock deformation results when rocks undergo permanent strain (change in shape) in response to
applied stresses (forces). These stresses acting on the crust are due to tectonic processes at
plate boundaries and can be of three types:



compressive - due to plate collision
tensional - due to crustal stretching
shear - due to continental collision or conservative plat emotion
FACTORS THAT INFLUENCE HOW ROCKS RESPOND TO STRESS
How rocks respond to stress (i.e.) brittle deformation or plastic deformation depends on a
number of factors:



temperature
speed of deformation
properties of rock material itself.
Temperature - Hot rocks generally behave in a ductile (plastic) way, cool rocks in a brittle way.
Rocks deformed at low temperatures at or near the Earth's surface show far more evidence of
brittle failure (joints and faults) than those deformed at depth. At depth, major faults pass
into shear zones (where movement is taken up within the rock rather than taking place across a
single plane as in faults). The higher temperatures that exist at depth favour ductile rather
than brittle deformation, and shear zones can be considered as ductile equivalents of faults.
Speed - The speed at which rocks are deformed is known as the strain rate. Rocks deformed
quickly (high strain rates) show more brittle behaviour than those deformed slowly (low strain
rates). When an elastic band is stretched slowly it will expand (plastic/ductile deformation),
however if the force is applied too quickly it will snap (brittle deformation). Real rock
sequences naturally deform over millions of years, not seconds e.g. the Alps took 20 million
years to form! This slow rate of deformation is the reason why we can get folds in brittle
crustal rocks.
Rock Composition - Some rocks are almost always rigid and tend to resist deformation; others
almost always flow during deformation. This property is known as competence. Competent rocks
are stiffer, do not flow and break more easily than incompetent rocks. Sandstones, limestones
and igneous rocks tend to deform in a competent way (brittle failure) during deformation,
whilst mudstones and shales generally behave in an incompetent way.
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HOW BRITTLE STRUCTURES FORM
Faulting results when stresses applied to a rock exceed the fracture strength of that rock.
When these stresses are not equal in all directions, a fracture or fault forms. Theoretically,
the stresses acting on a rock can be broken down into three parts. These three principal
stresses always act at right angles to one another. Fault type is determined by the orientation
of these principal stresses. There are three different fault types:



normal faults - which form due to extension of the crust.
reverse faults - which form due to compression of the crust.
tear faults - which form due to shearing (rocks moving passed each other).
The principle stresses are labelled:



P1
P2
P3
In the diagram opposite the three principle
stresses are in this instance all equal (i.e.)
P1=P2=P3. As the stresses are equal in all
directions (usually from the pressure from the
weight of surrounding rocks) the rock will be
squeezed (compressed) and no fracture will
occur.
In the diagram opposite the stresses are not
equal so:



P1 becomes P min
P2 becomes P int (intermediate)
P3 becomes P max
P max means the most compressive force (but
not necessarily the highest overall force!) and
P min is always the least compressive force
(but not necessarily the lowest force!). You
are doing well if you are still with me! As one
stress is more compressive than the others
in this diagram, the rock will at first be
squashed until its elastic limit is reached and
then it will fracture. We know from
181
experiments that under unequal compressive stress like this a fracture will always form at an
acute angle to P max (usually 20º- 40º).
a). Normal Faults
Although the crust is being pulled apart by
tectonic forces (P min), the greatest
compressional forces (P max) are at right
angles to the bedding. Since failure planes
always make a small angle with P max,
normal faults should form as planes dipping
steeply through the sedimentary layers. If
two fractures form a rift valley will be
develop.
A good way of remembering which fault is which is by looking to see if the ground has been
lengthened (due to tension), if it has it is normal (i.e. they both have the letter L in them).
b). Reverse Faults
Here the crust is being compressed by
tectonic forces (squeezed) so P max (principle
stress of maximum compression) is in a
horizontal plane and P min (principle stress of
maximum tension) is in a vertical plane. As the
beds are squeezed and thicken brittle
deformation will eventually occur. Since failure
planes always make a small angle with P max,
reverse faults should form as planes dipping
gently through the sedimentary layers.
A good way of remembering which fault is
which is by looking to see if the ground has
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been shortened (due to compression), if it has it is reverse fault (i.e. they both have the letter
S in them).
A thrust fault is a low-angled reverse fault and usually has an almost horizontal thrust plane.
c). Tear Faults
Strike-slip faults (tear faults) should form
when the principle stresses of compression
(P max) and tension (P min) both lie in a
horizontal plane. Tear faults can be
described as dextral or sinistral. If you
stand on the outcrop of a particular bed
facing the fault plane then for dextral
movement the corresponding outcrop on the
other side of the fault plane will be displaced
to the right and for sinistral movement it will
be to the left.
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DESCRIBING FAULT GEOMETRY
1. Fault Dip Angle & Direction – The dip of the fault is the angle between the horizontal and
the fault plane. Strike-slip faults are vertical and so have a dip angle of 90°. Normal faults tend
to have less steep dips, between 45° and 70°. Reverse faults tend to be less steep again, usually
less than 45°. Thrust faults are low-angled reverse faults and usually have a dip angle of less
than 10°. The direction of the fault dip is just the compass direction the fault plane is inclined
towards, such as North, South or SW. This can only be worked out from a cross-section.
2. Fault Orientation – The orientation or trend or strike of the fault is the two directions the
fault line points towards as it outcrops onto the surface, such as North-South, East-West or
NE-SW.
3. Fault Type – Faults are classified into groups based on their dominant slip direction. Thus if
the main displacement seems to be along the strike of the fault it is a strike-slip fault,
whereas a dip-slip fault shows displacement chiefly in the dip direction of the fault. Dip-slip
faults are further divided based on the sense or direction of displacement with respect to the
fault dip. Where the fault dips towards the downthrow side it is a normal fault and where the
fault dips towards the upthrow side it is a reverse fault. Alternatively, a normal fault has its
hangingwall downthrown and a reverse fault has its hanging wall upthrown.
a. Normal Fault
Normal faults are dip-slip faults produced by tension in the Earth's crust. Normal faults dip
towards the downthrown side or towards younger beds on a map! Normal faults result in
crustal extension (i.e. the crust has been lengthened).
b. Reverse Faults
Reverse faults are dip-slip faults produced by compression within the Earth's crust. Reverse
faults dip away from the downthrow side. Reverse faults may cause repetition of beds in
boreholes. A reverse fault results in crustal shortening (i.e. the crust has been shortened).
Normal Fault
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Reverse Fault
c. Thrust Faults
Thrust faults are low-angled reverse faults
where the angle of dip of the fault plane is low
or nearly horizontal. In a thrust fault
compression has caused the hangingwall to be
thrust up over the top of the footwall. Thrust
faults are reverse faults. The diagram is an
example of a faulted nappe.
d. Tear or Wrench Faults
Tear faults are strike-slip faults where the movement has been horizontal. The terms dextral
and sinistral are applied to strike-slip faults in order to describe their displacement. Dextral
Displacement: Is right hand movement. Sinistral Displacement: Is left hand movement
(remember when your Granny was at school it was a sign of the devil to write with your left
hand and you were regarded sinister, hence the origin of the term!)
Golden Rule for working out strike-slip displacement:'If you stand on the outcrop of a
particular bed facing the fault plane then for dextral movement the corresponding outcrop on
the other side of the fault plane will be displaced to the right and for sinistral movement it will
be to the left'.
Dextral tear fault
Sinistral Tear Fault
Distinction between Normal and Strike-slip faults




There may be lateral (sideways) displacement of the axial plane trace of a fold.
There may be displacement of a vertical feature such as a dyke.
There may be displacement of a granite mass.
The outcrop pattern of beds may be off-set.
4. Fault Throw – The throw of the fault is the vertical displacement along the fault.
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5. Fault Sense of Displacement – The relative movement directions on each side of a fault is
called the sense of displacement, and is shown on a cross-section by paired half-arrows. The
side that has gone up relative to the other is known as the upthrow side, and the side that has
gone relative to the other is the downthrow side. Often in the field it is not possible to simply
use displacement of beds across the fault plane to determine in detail how a fault has moved. It
may be possible to establish, for instance, the downthrown side of the fault; the dip of the
fault plane; and whether the fault is normal, reverse, or strike-slip. However, the exact path
one wall took relative to the other cannot be established from this information.
Clues to the movement direction lie along the fault plane itself. Many fault planes show a
polished surface, called a slickenside surface, containing either aligned grooves (slickenlines) or
elongate crystalline fibres (slickenfibres), commonly of quartz or calcite. These grooves or
fibres are aligned with the movement direction of the fault and often give a smooth feel along
the fault plane surface in the direction of movement. Slickenfibres form when the walls are
held apart by mineral-rich fluids under pressure, whereas the slickenlines form where they are
not. Although theses slickenside features can give an accurate measure of the direction of slip
of the fault the movement history of the fault may be a protracted one with periods of
reactivation, and generally speaking the slickenside features only record the last episodes of
movement. So, although they provide an accurate movement direction they rarely provide a
complete record of movement.
HOW FOLDS FORM
There is no single mechanism that explains how all types of folds form. Folds form in two
fundamentally different ways: as a mechanical consequence of shortening along layers due to
compression (layer-parallel compression) or an uneven loading oblique to perpendicular layers
(such as a mechanical consequence of fault displacement, shear folds or flow/slump folds).
Buckle folds are extremely common. They form as a
mechanical consequence of shortening across layers.
For example, gently pushing the two extremities of
a paper sheet on a table towards each other
reproduces this folding mechanism. When the force
is small the sheet remains flat. As the force is
slowly increased, it suddenly becomes curved. The
presence of layers of contrasting competence is an
important requirement for buckling. Some rock
types are almost always rigid and tend to resist deformation (e.g. sandstone, limestone and
most igneous rocks); others almost always flow during deformation (e.g. shale, mudstone and
clay). This attribute is termed competence. Competent lithologies are stiffer, flow less easily
and break more readily than incompetent lithologies at the same temperature and pressure. If
a competent layer set in a less-competent matrix is shortened along its length, a mechanically
unstable situation develops where the competent (stronger) layer is deflected sideways and the
weaker matrix fills the gap, forming a fold. These folds are called buckle folds because they
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have formed by buckling – they have shortened along the length of the layers. They are
generally not related to faulting.
The diagram here illustrates some of the principles of
buckle folding in a layered composite of materials of
different competences shortening along its length. If
competent layers are separated from other competent
layers by a considerable thickness of incompetent material,
the wavelength of the buckle fold is controlled by both the
thickness of the buckled layer and the competence
contrast between layers. An increase in the thickness of
the competent layer or an increase in the competence
contrast both lead to an increase in the wavelength of the
fold. Where the competent layers are thicker, as in A, the limbs are longer and so are the
wavelengths. Where the competent layers are thinner, as in C, the limbs are shorter and the
wavelengths are also shorter.
Bending is a second fault mechanism, which this time involves
forces acting at high angles to the layers and not along the
layers as in buckling. The variation in competence between the
layers is not important here. A layer subjected to bending is like
a notebook supported at the ends and loaded in the middle. The
notebook bends downward when the load is placed in the middle.
Bending mostly produces very gentle folds at a variety of scales.
Lithospheric-scale folds produced by bending are common in
cratons where vertical stresses produce broad basins and
domes. Flexural bending of lithospheric plates also occurs at
subduction zones and adjacent to oceanic ridges, at smaller accumulations of volcanic rocks in
the oceans, and where loading by ice or sediments causes the lithosphere to undergo flexural
readjustment. Flexural bending is also involved in foreland basins adjacent to mountain belts.
At a crustal-scale bending produces gentle up-warping and down-warping such as the arching of
cover rocks above an intruding pluton (often termed a forced fold). Other forced folds include
drape folds which are formed when sediments, which cover a more rigid basement, flex and
drape the deep fault scarps in response to vertical movements along basement faults. Other
common types of forced fold are the result of faulting processes. Fault-bend folds are created
as a consequence of movement along a fault
line. Imagine trying to push a rug up or down a
set of shallow steps: it would crumple (i.e.
develop folds) as it moved over the steps. The
same is true of sheets of rocks moving on
stepped faults and thrust faults – both
anticlines and synclines may form as a hanging
wall moves over a fault ramp.
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All of the folds discussed above are tectonic in origin: that is they are caused by some form of
tectonic force acting on already consolidated sediments, i.e. rocks. However, a last type of fold
mechanism can occur in materials that are not yet consolidated such as flowing lava or watersaturated, unconsolidated sediments. In many areas of deformed sedimentary rocks it is
difficult to distinguish between structures formed during deposition when the sediment was
unconsolidated, and those formed after lithification in response to tectonic forces. On first
examination many synsedimentary structures (those formed before lithification) such as slump
folds have superficial geometric similarities to tectonic folds. It is therefore extremely
important when studying folds to distinguish between prelithification structures and postlithification, tectonic structures.
Synsedimentary folds or slump folds are generally tight to isoclinal with variable shapes at low
fold amplitudes. The orientation of the fold axes in a slump fold sequence tends to vary widely
and recumbent folds are dominant. The fold axial planes may be stacked up (imbricated) one on
top of the other. Axial plane cleavage may be developed, particularly in the hinge regions but
these are probably due to later compaction during burial. The upper surfaces of slump folds are
very often truncated due to erosion and fractures and veining are usually absent. In general,
slump folds have no geometric relationship to the large-scale folding of an area (regional fold
pattern). Synsedimentary folding is commonly associated with disturbed sedimentary
structures such as dewatering structures (flame structures and load casts). Slumps are also
characterised by extensional structures at the rear, whereas the front of the slump fold will
be marked by localised compression, with the development of folds, thrust faults and
imbrications. The table below shows the criteria that could be used to distinguish between
synsedimentary folds and tectonic folds. A is the most reliable criterion and C is the least
reliable. Several criteria must be used in conjunction in order to determine the origin of a
particular fold.
Synsedimentary Folds
Truncated by overlying beds
Burrowing or boring organisms
Cut by dewatering structures
Undeformed clasts or fossils
Folds with weak or no axial
planar cleavage
Fold axes vary in orientation
Dominantly recumbent folds,
which may be imbricated
Both
extensional
and
compressional features with
no veins developed
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Reliability
Index
A
A
A
A
B-C
B-C
C
C
Tectonic Folds
Fold geometry relates to regional
structure
Axial planes symmetrical around
large folds
Limited spatial distribution
Symmetrical fracture patterns
and veins
Slickensides and striations down
fold limbs and around fold hinges
Associated
brittle
fault
structures
Continuity of axial planes across
several beds
Well-developed
axial
plane
cleavage in mudstones
Reliability
Index
A
A
A
A
B
B
B
C
DESCRIBING FOLD GEOMETRY
1. Fold Type - Upward-arching (or upward-closing) folds are antiforms; those that arch or
close downwards are synforms. In the absence of way-up criteria, folds cannot be classified as
anticlines or synclines. In anticlines, the oldest rocks are found in the core of the fold; in
synclines, the youngest rocks are in the fold core. Neutral folds close sideways, so they cannot
be defined as antiforms or synforms. However, neutral folds can still be identified as anticlines
or synclines.
Antiform
Synform
If two adjacent folded surfaces have the same shape, the fold they enclose is a similar fold;
the layer will be thicker in the hinge than the limbs. If the distance between the two surfaces
remains constant, in other words the thickness of the layer does not vary, the deformed layer
constitutes a parallel fold. The deformation of a sequence of rocks displaying a range of
competencies will generate parallel folds in the competent layers and similar folds in the
incompetent layers. Buckle folds are very common and form when layers are compressed along
their length. This type of compression (or buckling) leads to the development of symmetrical
folds. Shear folds on the other hand occur when compression is oblique to the layering and
results in asymmetrical folds.
2. Fold Orientation – It is the fold axis which is used to determine the orientation (direction
facing or trend) of the fold. The two directions the line of the fold axis points to is the
orientation or trend of the fold. This trend is described by using two (opposite) points of the
compass, such as East-West, North-South or NE-SW.
3. Fold Style – Style relates to the inter-limb angle of a fold. The inter-limb angle, as the
name implies, is the angle between the limbs. In a non-plunging fold, it can be calculated by
adding together the dips of the two limbs and subtracting them from 180°: the remainder is
the inter-limb angle. This gives five categories of fold:





Gentle folds –inter-limb angle of 120° - 180°
Open fold - interlimb angle of 70° - 120°
Closed fold - interlimb angle of 30° - 70°
Tight fold - interlimb angle of less than 30°
Isoclinal (literally, ―equal dip‖) fold – interlimb angle of 0° i.e. the limbs are parallel.
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Gentle
Open
Closed
Tight
Isoclinal
4. Fold Symmetry - The diagram on the next page illustrates the concept of fold symmetry.
When one limb of a fold is the mirror image of the other, and the axial plane is a plane of
symmetry, a fold is said to be symmetrical. It is a common misconception that the dips of limbs
of a symmetrical fold must have equal dips either side of the axis. This need not be the case.
Irrespective of the orientation of a fold, a symmetric fold has limbs of equal length, while an
asymmetric fold has limbs of different length. It is not easy to identify the symmetry of a
fold from the evidence on a geological map or cross section without knowledge of the length of
the fold limbs.
Symmetrical fold
Asymmetrical fold
5. Fold Attitude - The variation in limb dip either side of a hinge is a function of the attitude
of the fold. Attitude is determined by the orientations of the axial plane and the limbs, as
shown in the diagrams below.
Upright
Steeply
inclined
Moderately
inclined
Gently
inclined
Recumbent
In the first fold in the diagram above the limbs have similar dips, so the axial plane is more or
less vertical – the fold is therefore described as upright. In the second fold the axial plane has
moved away from the upright, but the younging direction in both limbs is upwards – the fold is
therefore said to be steeply inclined. The fourth fold in the diagram above shows an axial plane
so inclined that the younging direction in one limb is now downwards, allowing the fold to be
defined as overturned. In the last fold the axial plane is more or less horizontal and the fold is
recumbent. It is the inclined fold that is most frequently mis-identified as asymmetric.
6. Fold Dimensions – Fold dimensions can be measured just like waves. The wavelength of a fold
is measured between two adjacent antiformal (or synformal) hinges. The amplitude of a fold is
half the distance between antiformal and synformal hinges, measured at right angles to the
wavelength.
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Amplitude
Wavelength
Folds are described according to their shape and
orientation. The sides of a fold are called limbs; when the
limbs dip in towards the hinge or axis, the fold is a syncline,
and younger rocks are found in the middle. If the limbs dip
away from the hinge or axis, the fold is described as an
anticline; here older rocks are at the centre, in the core of
the fold. When the age of the rocks cannot be worked out
then a downfold is termed a synform and an upfold is
termed an antiform. These two terms describe the shape
whereas the terms anticline and syncline describe the age
of the rocks being folded.
With 'equal opposed compression' upright folds develop producing alternating synforms and
antiforms that are symmetrical. The term upright means the axial plane is vertical. With
'unequal opposed compression' axial planes are inclined and the folds produced are
asymmetrical about the fold axis. When these inclined folds are formed the central limb
begins to become inverted (upside down) and in the diagram above the grey bed (Y) appears to
be the youngest when it is actually the oldest. With continued compression overfolds may form.
Here the limbs of the folds dip in the same direction but at different angles. The central limb
of the fold is inverted i.e. it is the wrong way up. Sedimentary structures could be useful for
indicating the way up of this fold in the field, e.g. graded bedding, flute casts, tool marks,
desiccation cracks, cross laminations, dune bedding etc. Under extreme compression,
recumbent folds form. Here the axial planes are horizontal or nearly so. If compression
continues the fold (or series of folds) may be faulted to produce a thrust fault which is a low
angled reverse fault. Very large scale recumbent folds are known as nappes (French: sheet or
table cloth because these large-scale crumpled folds are analogous to those that form when a
tablecloth is pushed across a table) and are associated with mountain regions that have
undergone orogenesis (Mountain Building) e.g. the Alps. Nappes indicate that very complex
191
deformation has occurred in an area and are commonly thrust faulted. These thrust faults may
have moved tens of miles.
The diagrams below show some fold terminology and how folds should be described.
192
UNCONFORMITIES
Unconformities represent significant time breaks in the rock record (known as stratigraphic
succession). They may also represent periods of no deposition and erosion. The time gap is
usually represented by formations of different stratigraphic age which are separated by a
boundary known as the surface of unconformity. There are four types of unconformity:
1) Angular Unconformity
2) Disconformity
In an angular unconformity the lower,
older beds dip at a different angle to the
younger, upper beds. This usually occurs
where unfolded, younger beds rest on
folded, older beds.
A disconformity is minor type of
unconformity, representing a short and
often local period of non-deposition of
sediment.
3) Parallel Unconformity
In parallel unconformity the lower and
upper series of beds dip at the same
amount and in the same direction. This is
due to a period of erosion removing layers
of rock in a succession of horizontal
strata.
193
4) Non-Conformity
A non-conformity is an unconformity in
which
the
upper
layers
overlie
metamorphic or igneous rocks.
194
GEOLOGY OF THE LITHOSPHERE
1. Earth‘s Heat Flow






What is the Earth‘s heat flow?
How is heat lost through the Earth‘s lithosphere?
What is the origin of this heat flow?
How does heat flow vary with depth?
How is rock strength related to temperature?
How and why does heat flow vary across the Earth‘s surface?
2. Evidence for the Structure of the Crust & Upper Mantle



What is the lithosphere and what is the structure of the lithosphere?
What evidence supports a layered internal structure of the lithosphere?
What is the thickness of the lithosphere and why does it differ between continents
and oceans?
3. Formation and Destruction of Oceanic Lithosphere






What is the structure of the oceanic lithosphere?
What evidence supports a layered internal structure of the oceanic lithosphere?
How does oceanic lithosphere form?
How can the rates and directions of sea-floor spreading be calculated?
Where and how is the oceanic lithosphere reabsorbed into the mantle?
How do ocean basins evolve?
4. Formation and Deformation of the Continental Crust






Why are the Earth‘s oldest crustal rocks found in continental areas?
What are the large-scale features of the continents?
How did the large-scale features of the continents form and how are they related to
tectonic setting?
How did the continental areas form?
What forces are acting on the continental crust?
How do these stresses cause brittle & ductile deformation on all scales in crustal
rocks?
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1. EARTH‘S HEAT FLOW
Heat is generated within the Earth and lost at the Earth's surface. The Earth's heat flow is
the rate at which heat is moving upwards through a particular part of the Earth's crust. This
rate is measured in Watts produced per square metre (Watts/m²). There are four mechanisms
of heat transfer in the Earth: conduction, convection, advection and radiation. In every case,
energy moves from a higher level to a lower level, and the process by which it does so is
different for each mechanism. As the heat is generally moving upwards through solid rock, the
main mechanism of heat transfer is by conduction. Conduction then is a process by which heat
in a material is transferred by the interaction of atoms and molecules within the material. The
low thermal conductivity of rocks means conduction of heat through the crust is very slow. In
fact the thermal properties of rocks are such that any heat present in the Earth at its
formation could only have been conducted a distance of a few hundred kilometres throughout
the planet's whole lifetime! Because of this conduction is unlikely to be the only process by
which heat in the Earth is transported. Within the Earth there are mass movements of
material, and these moving masses carry heat with them. This process of heat transfer is called
convection. Although the mantle is solid it is convecting. In order for convection to occur, hot
material must underlie cooler material. The hotter and cooler materials do not simply change
places by leaking through each other. During convection, columns of relatively hot material rise
up, and columns of relatively cool and therefore denser material sink. This convection transfers
heat outwards from the interior towards the base of the lithosphere and is much faster than
conduction in transferring heat. However, even this process could not have transferred to
space all the heat generated in the Earth's interior, so the heat being generated on the
196
surface today must be the result of heat generated long in the past. Advection is rather like
the upward part of convection; movement of a whole region takes heat up with it. For example,
when molten material (magma) moves up through fractures in the lithosphere it either spreads
out at the surface as a lava flow or, if it is injected, it cools and crystallises without returning
to its sources as in convection. So heat is transferred by molten rock from deeper levels where
melting is taking place to shallower levels where it solidifies, losing its heat by conduction into
the overlying crust or by radiation into space. Any planetary body that exhibits, or has
exhibited, volcanic activity must have lost some of its internal heat in this manner. In the case
of radiation, photons (electromagnetic radiation) carry away the heat energy from the surface
of the Earth into space.
So, most of the heat reaching the Earth's surface does through what is termed solid-state
convection through the mantle, where hot mantle rises at speeds of few centimetres a year and
transfers its heat to the base of the lithosphere. Having lost heat, that bit of the mantle
becomes slightly denser, and sinks downwards to complete the loop. Most of the heat deposited
at the base of the lithosphere trickles through to the Earth's surface by conduction, but some
is carried higher by molten bodies of rock that can intrude high into the crust, or even reach
the surface at volcanoes. The total average global rate of heat loss is 4 x 10 13 Watts, which
works out at an average heat flow of about 0.08 Watts/m², though it varies considerably from
place to place. To help you appreciate the magnitude of this figure think about this; the heat
flowing out of an area the size of a football pitch could illuminate continuously, in theory, 5.6
100-Watt light bulbs!
Heat flow from ocean basins (0.095 Watts/m²) is higher than the global average (0.08
Watts/m²), although this too varies across an ocean basin. All oceans show a similar
distribution of heat flow. For oceanic lithosphere up to 35 Ma old, the heat flow is high (>0.3
Watts/m²), particularly near spreading oceanic ridges, such as the East Pacific Rise in the
Pacific Ocean. For oceanic lithosphere older than 65 Ma, the mean heat flow is lower (<0.05
Watts/m²). The lowest heat flow values (<0.04 Watts/m²) of all come from ocean trenches,
such as the Peru-Chile Trench off the west coast of South America.
There are clear reasons for this decrease in heat flow from ridge crest to ocean basin to deepsea trench which is directly related to plate tectonic and mantle processes within the
lithosphere. As the oceanic lithosphere is pulled apart at constructive plate margins, it thins
allowing the partially mobile asthenosphere to rise into within 50km of the surface and
undergoes decompression melting. So oceanic ridges like the East Pacific Rise sit atop a rising
plume of hot rock and magma that carries heat from the deeper mantle. This magma cools,
solidifies and becomes attached to the oceanic lithosphere. As the lithospheric plate spreads
away from the ridge, it loses heat by conduction to the sea floor and gradually cools. So heat
flow decreases with age and hence distance from a mid-oceanic ridge. The oldest parts of the
sea floor should have the lowest heat flow, and these parts should be found farthest from
ridges and closest to deep-sea trenches, where the cold slabs sink back into the mantle. The
process of cooling of the recently created oceanic lithosphere thus dominates oceanic heat
flow. It is thought that this form of cooling may account for as much as 60% of the total heat
flow from the Earth and that this may represent a major mode by which the Earth is cooled.
197
The oceanic lithosphere is not solely cooled by conduction. There is a much more efficient
mechanism of heat transfer, which removes heat faster than is possible by conduction. The
formation of oceanic lithosphere at ridge crests involves the contact between hot rocks and
cold seawater. The rocks crack and fracture as they cool, allowing seawater to penetrate them,
down to depths of at least a few kilometres. The interaction between hot rocks and cold
seawater has the effect of setting up a very powerful system of convection currents of water
within the oceanic lithospheres, which act to remove heat from the rocks much more quickly
than could be achieved by conduction alone. Dramatic conformation of these convecting systems
has come from underwater exploration near ridge crests, where in some places hot plumes
(black smokers) erupt from the ocean floor at temperatures exceeding 620º C. Hot spots such
as at Hawaii may also increase the heat flow above the average, but this is related to a mantle
plume and not directly to plate tectonics.
The lowest values of heat flow are found in ocean trenches such as the Peru-Chile Trench. As
the old, cold oceanic lithosphere is being subducted, it is heated partly by friction along its
upper surface and partly by conduction from the hot mantle into which it is descending. There
is also dehydration of the minerals formed by metamorphism of oceanic lithosphere and this
has a cooling effect on the upper part of the slab. Because rocks have a very low thermal
conductivity, the main bulk of the slab will take a very long time to heat up by conduction and
reach thermal equilibrium with the mantle into which it is descending. All these factors
contribute to lowering temperatures at oceanic trenches and therefore making heat flow to the
surface much lower. However, on the other side of the oceanic trench there will be an increase
in heat flow, due to the effects of subduction on the upper parts of the descending slab and
the immediately overlying mantle. The basaltic rocks forming the upper part of the descending
slab are heated releasing a watery fluid into the overlying mantle. This causes hydration melting
of the basaltic slab and in the overlying mantle forming andesitic magmas, which rise through
the lithosphere convecting heat upwards through the lithosphere. Some of these intrude into
the crust forming volcanic island arcs like the Indonesian islands and continental volcanic chains
such as Mt St Helens in the Rockies. The increase in heat flow in island arcs and continental
volcanic chains is therefore a result of intrusion and eruption of igneous rocks.
Measuring heat flow on the continents is much more difficult than on the ocean floor. The
average heat flow in all continents is about 0.06 Watts/m². Just like in oceanic crust there is a
decrease in heat flow with increasing age of continental rocks. However, the rate of decrease
is very different between the two types of crust. In oceanic areas heat flow decreases quickly,
whereas in continental areas heat flow decreases much more slowly. The fact that heat flow
declines so slowly with increasing age in continental rocks implies that there is an additional
source of heat in continental lithosphere, which is not available in the ocean. The continental
rocks are enriched with heat-producing isotopes (K-40, Th-232, U-235 and U-238) in granitic
rocks relative to the oceanic crust, and these provide the additional heat source.
198
The crust and the upper mantle beneath it together form a single rigid layer known as the
lithosphere, and is split into a number of tectonic plates. The mantle below the lithosphere is
solid, but it is not rigid and is convecting at speeds of the order of centimetres per year (this
is how heat is transported from the Earth's interior to the base of the lithosphere). The upper
part of this convecting region in the mantle is particularly weak, and is called the
asthenosphere. It is this weak layer that enables the plates to move.
We know from boreholes and deep mines that the temperature in the Earth increases with
depth. The average rate of temperature increase (the geothermal gradient) in regions remote
from volcanic activity is about 20 to 30ºC per kilometre. If this rate were maintained through
the outer layers of the Earth, the melting point of basaltic rocks would be reached at the base
of the continental crust. However, we know this cannot be true due to the transmission of Swaves all the way through the mantle until the outer core is reached. The solution to this is
that something like 50% of the heat escaping to the Earth's surface is generated in the crust,
and that only a half comes from below. This means that the geothermal gradient gets less (10ºC
per kilometre) as you go deeper. One reason for this lower value below the crust is that the
mantle is convecting over long periods of time. Convection is much more effective than
conduction at transporting heat. If convection could magically be stopped, then heat would build
up within the mantle and its temperature would rise at a rate of 1ºC or so every million years
until it melted. The solid-state convection within the mantle is actually what stops it becoming
hot enough to melt.
So what is the origin of the Earth's heat? Between 20% and 50% of this heat is to be found in
the Earth's core and is thought to be residual heat left over from the time of the Earth's
formation (including the accretion of large planetary bodies in the Earth's early history). This
ancient heat continues to leak out, but is not being generated today. The remainder represents
heat that is still being generated today by the decay of radioactive elements. All radioactive
decay produces heat, which is described as radiogenic heating, but there are only three
elements today whose decay produces significant amounts of heat. These are uranium, thorium
and potassium. Most potassium atoms are stable and non-radioactive. Each of these contains a
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total of 39 heavy particles (protons and neutrons) in its nucleus, and is described as potassium39. However, about one potassium atom in every 10,000 contains an extra neutron, making a
total of 40 heavy particles. These atoms of potassium-40 are unstable, and undergo
radioactive decay. All atoms of uranium and thorium are unstable.
The abundances of all three heat-producing radioactive elements, and therefore the rate of
radiogenic heat production, are much greater in granite (continental crust) than in basalt
(oceanic crust) and peridotite (mantle) which explains why about 50% of the Earth's radiogenic
heat is generated in the continental crust. So heat flow in continental regions is dominated by
the decay of radioactive elements in the upper part of the crust so that the geothermal
gradient in the lower crust and upper mantle is much less than at the surface.
The likely temperature distribution in the outer parts of the Earth (the geotherm) is shown in
the diagram over the page.
This diagram shows why the lithosphere is rigid and brittle, and why the asthenosphere is weak
and plastic. Above 100 to 150km depth the temperatures have not increased enough to cause
the rocks of the lithosphere to start to melt. However, after a depth of approximately 150km,
temperatures have continued to increase enough that they are now close to the melting point of
the rocks found here. This is why the asthenosphere is mobile. The heat causes rocks to melt
ever so slightly (only 1-5%) and become weak and behave in a plastic manner.
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2. STRUCTURE AND EVIDENCE FOR THE STRUCTURE OF THE CRUST & UPPER MANTLE
(LITHOSPHERE)
The lithosphere (from "lithos", the Greek word for rock) is the outer most part of the Earth,
and includes the crust and upper mantle. It constitutes a single rigid unit which is relatively
cold (<1300ºC) and brittle. Under stress the lithosphere will generally fracture. In contrast,
below the lithosphere the mantle is relatively weak, and is termed the asthenosphere (using the
Greek word for "weak"). Although still solid the asthenosphere is partially molten (probably only
about 5%) and responds in a ductile fashion when subjected to stress.
The evidence for this structure has come from the study of earthquake body waves. P-waves
and S-waves are known as body waves, because they travel through the Earth's interior (i.e.
through the body of the Earth). P-waves (push waves) involve the transmission of compressions
and expansions. As a P-wave passes through a material, the particles of the material vibrate
backwards and forwards along the direction of wave travel. This is a longitudinal movement. Pwaves are 1.7 times as fast as S-waves. P-waves range in speed from 1 km/second through
water to 14km/s at the base of the mantle. S-waves range in speed from 1 km/s in loose
sediment to 8km/s at the base of the mantle. S-waves (shear waves) involve side to side
displacements. The motion of the material particles as the wave passes through is at right
angles to the direction of wave travel. This is a transverse movement. The fact that seismic
waves travel at speeds which depends on the properties of the materials (elastic and density)
they are passing through allows the investigation of the internal structure and thickness of the
lithosphere. The waves from an earthquake source are transmitted outwards in all directions.
The source is known as the earthquake focus. The point on the Earth's surface vertically above
the focus is called the earthquake epicentre. Because waves travel outward in all directions
from the focus, they are potentially capable of being detected by seismometers anywhere,
although how far the waves travel depends on the size of the earthquake in the first place and
on how rapidly the wave energy dissipates.
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1. Seismic Velocity: The time it takes for a wave to be picked up on a seismograph can give
information about the properties of the materials it is travelling through. The velocity of
seismic body waves is related to the elastic properties of a rock and the density of a rock. The
elastic properties can be further split up into the incompressibility (the bulk modulus) and the
rigidity (shear modulus) of a material. So how do these properties affect the speed with which
a body wave can travel through it? Firstly, let us take p-waves. The following equation
summarises how p-wave velocity is related to elasticity (incompressibility & rigidity) and
density.
Vp = P-wave velocity
κ = bulk modulus (a measure of incompressibility or resistance to change in volume without
change in shape)
μ = shear modulus (a measure of rigidity or resistance to change in shape without a change in
volume)
ρ = density
Incompressibility (or bulk modulus – denoted by the Greek letter ―Kappa‖) is a measure of the
amount a rock can resist a change in volume if put under pressure. A rock is more able to resist
a change in volume than, for example, a sponge, since a sponge is easily compressed. Therefore,
a rock is said to be more incompressible. P-waves, which are push and pull waves, move through
rocks by compression and dilation. As a rock gets harder to compress P-waves can move through
more quickly. The formula backs this up and shows that P-waves will increase in velocity as
incompressibility increases.
The rigidity (or shear modulus – denoted by the Greek letter ―Mu‖) is a measure of the degree
to which a rock is able to resist change in its shape without changing its volume. For example, if
you grab the top and bottom of a sponge it is quite easy to distort its shape. If you tried the
same with a rock it is very difficult to distort it and change its shape. The rock is therefore
much more rigid than a sponge. S-waves, which are shear waves, move through a rock by shaking
from side to side. As a rock gets more rigid the vibrations can pass through it more quickly. An
analogy for this is the twanging of ruler over the side of a table. If a wooden ruler is used the
ruler vibrates quickly, however, if a plastic one is used it is less rigid and vibrates less quickly.
The formula below backs this up and shows that S-waves will increase in velocity as rigidity
increases. Note that there is no incompressibility component in this equation because S-waves
do not change the volume of a rock when they pass through it they just change its shape. An
important point to mention here is that fluids (liquids and gases) have no rigidity and do not
resist shear, but will deform without a return to their original shape. Therefore all fluids are
incapable of transmitting shear waves and as S-waves are shear waves they are not transmitted
through liquids or liquid rock.
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Vs = S-wave velocity
μ = shear modulus (a measure of rigidity or resistance to change in shape without a change in
volume)
ρ = density
Density (denoted by the Greek letter ―Rho‖) also affects the velocity of both P and S-waves. In
both instances density decreases their velocity. So when rocks are put under high confining
pressure such as when deeply buried both density and the elastic moduli properties of
incompressibility and rigidity increase. The formulae show that the velocity of P-waves increase
as incompressibility increases, and the velocity of S-waves increase as the rigidity increases,
but both decrease as density increases. To explain the conundrum that body waves increase in
velocity with depth inside the Earth even though density increases as well tells us that
incompressibility and rigidity increase faster than density with depth.
In summary, P-wave velocity depends upon the incompressibility of the rock as well as the
density. So the more incompressible the rock, the greater the P-wave velocity will be. S-wave
velocity depends upon the rigidity of the rock as well as the density. So the more rigid the
rock, the greater the S-wave velocity will be. Geologists can therefore get a very good
indication of the properties of a rock deep underground. The faster a body wave travels the
more rigid and less compressible the material it is travelling through is.
The fact that body waves travel at speeds which depends on the material properties (elastic
moduli and density) they are travelling through allows geologists to use seismic body wave
observations as evidence for the variation in thickness and mechanical properties of the
lithosphere and asthenosphere. The depth of the lithosphere varies from its thinnest at about
10km at mid-oceanic ridges, thickening with age and distance from these ridges to about
100km. The thickest areas are under old, shield areas of continental lithosphere which could be
up to 350km in thickness. Geologists can work out the depth of the lithosphere from the time
taken for body waves to reach the surface at specific points on the Earth. The two diagrams
below show seismic S-wave velocity varies at two different locations. The first diagram on the
left shows S-wave velocities reducing at about 25 km depth due to the rocks becoming more
compressible and less rigid. This then indicates the base of the lithosphere and probably
represents oceanic lithosphere. The second diagram (right hand one) shows seismic S-wave
velocities reducing at about 35-40km depth. This then represents the base of the lithosphere
and probably represents continental lithosphere.
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In 1926, the seismologist Gutenberg noted that seismic waves from earthquakes with foci
located at depths between 50 and 250km in the mantle took longer to arrive than they should
have done according to simple calculations based on a homogeneous mantle (i.e. a mantle made
up of materials with the same mechanical properties). Since the travel times are long the Pwave and S-wave velocities must be lower. The layer has therefore come to be known as the low
velocity layer (LVZ). The reason why seismic velocities are reduced in the LVZ is due to its
mechanical properties. Remember that the P-wave velocity is reduced and S-wave velocity falls
to zero in a liquid. However, if the LVZ was partially liquid, say 5% liquid and 95% solid (rather
like slush after the snow has just begun to melt) both P- and S-wave velocity will fall, but the
latter not to zero. This is because the incompressibility and rigidity moduli fall if some liquid is
present, but the latter not to zero.
As can be seen from the seismic velocity diagram below, P-wave velocities increase through the
lithosphere as it becomes more rigid and less compressible. P-waves travel at approximately
6km/s in the upper crust, 7km/s in the lower crust and 8km/s in the upper mantle. Below the
rigid lithosphere P-wave velocities slow down to 7.8km/s indicating that the material there is
much less rigid than those above, and probably contains some partially molten mantle material
(1-5%). This area of reduced P-wave velocities known as the low-velocity zone is made of the
same material above and below it but because it is slightly molten it is able to deform and flow,
and roughly equates to the location of the asthenosphere. It is important to realise the
significance of the low-velocity layer in terms of the Earth‘s physical properties. Because this
partially molten layer exists, this region allows the relative movements to take place between
the lithosphere and asthenosphere and therefore all the processes of plate tectonics.
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2. Seismic Refraction & Reflection: The study of the way seismic waves are refracted can
also give information about the structure of the lithosphere. The basis of the seismic
refraction method is to set off seismic waves at one point and then determine how long it takes
for the waves to reach a series of observation points. In travelling from the source to the
detectors, the waves may be refracted or reflected at any boundaries or discontinuities which
happen to lie in their path. A discontinuity is a boundary between two layers in which the
seismic-wave velocities are quite different. This could arise because the two layers are made of
quite different materials.
The focus of an earthquake will emit seismic waves in all directions. However, one of the wave
paths will pass through Layer 1 into Layer 2 at an angle of incidence less than a certain critical
angle and will therefore be refracted. As Layer 2 is more rigid and less compressible than
Layer 1 the wave speeds up and will refract at an angle larger than the angle of incidence. The
exact relationship between the angle of incidence and the angle the wave is refracted by is
given by Snell's Law.
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One of the wave paths will also pass through Layer 1 into Layer 2 at an angle of incidence
greater than the critical angle and so no refraction takes place and the wave is reflected.
Because an earthquake emits seismic waves in all directions, there is bound to be one wave path
that strikes the boundary between Layer 1 & 2 equal to the critical angle. Here the wave is
refracted just inside the lower layer and parallel to the boundary. It is this wave that is of
particular importance in seismology. Another important property of seismic waves is that all
points on a wave-front emit secondary waves in all directions. Thus as the seismic wave travels
parallel to the boundary, it will continuously emit secondary waves, some of which will be
refracted back upwards into Layer 1. Each of these newly refracted waves will obey Snell's
Law, but be refracted back into Layer 1 at the critical angle.
A direct wave which has travelled entirely in Layer 1 will also be seen on a seismometer as well
as all the above (reflected & refracted).
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All this explains what Croatian seismologist Andrija Mohorovicic was seeing on his seismometers
in 1909. At seismograph Station 1 he saw a set of P-waves (P1) arrive then a set of S-waves
(S1), followed by another set of P and S-waves (P2 & S2). He realised that P1 & S1 were seismic
waves direct from the focus, and that P2 & S2 were waves refracted along the boundary
between Layer 1 and Layer 2. This observation informed him of the possibility of a discontinuity
between two layers of different materials. From further investigations his seismometers at
Station 3, further away from the focus again showed the two sets of P and S-waves arriving
but this time it was the refracted waves that were arriving ahead of the direct ones even
though they had travelled further. This then led him to believe that not only was there a
discontinuity or boundary between two different layers of different material but that Layer 2
was more rigid and less compressible, because the seismic waves had travelled much quicker
there than in Layer 1. At seismograph Station 2 he found the location (point of transformation)
where both sets of P and S-waves (direct & refracted) arrived at the same time. The diagram
overleaf attempts to illustrate his findings.
With this information Mohorovicic used a mathematical equation to calculate the depth to this
boundary and named it the Mohorovicic discontinuity. It is now known simply as the Moho for
obvious reasons and is the boundary between the crust and the mantle within the lithosphere.
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depth to boundary
distance
from
transformation
focus
to
point
of
velocity of P-waves in Layer 1 (7km/s)
velocity of P-waves in Layer 2 (8km/s)
Using the same techniques further sub-divisions of the lithosphere were found. For example,
Conrad in 1923 found evidence for a discontinuity between the upper crust and the lower crust,
which he called the Conrad Discontinuity. This then gave us the layered structure of the
lithosphere as we know it today (upper crust, lower crust & upper mantle).
So what is the thickness of the lithosphere? The asthenosphere is defined as a zone of partial
melting in which the material can flow (as opposed to the lithosphere above, which is rigid) and
occurs when mantle rocks start to reach their melting point at approximately 1300ºC. So the
base of the lithosphere is the point at which these temperatures are reached and therefore
does not have the same thickness everywhere. Beneath the continents, the thickness of the
lithosphere varies between 40 and 400km thick depending to some degree on its age. More
recent continental lithosphere, younger than 0.5 billion years, ranges in thickness from 40 120km. However, the much older Precambrian (3.5 to 0.5 billion years) continental lithosphere
ranges in thickness from 140 - 400km, and interestingly shows a bimodal thickness distribution
centred at 220km and 350km. Beneath the oceans varies with age, with young oceanic
lithosphere being as little as 10km thick, and old oceanic lithosphere about 100km thick. The
reason for these differences in lithospheric thickness is due to temperature. Beneath oceanic
ridges, where temperatures are high, the lithosphere is comparatively thin, whereas beneath
old oceanic lithosphere and continents, where temperatures are lower, the lithosphere is much
thicker as can be seen in the summary diagram below.
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3. FORMATION AND DESTRUCTION OF OCEANIC LITHOSPHERE
1. Structure of the Oceanic Crust
The oceanic lithosphere has a four part layered structure – Layer 1, Layer 2a & b, Layer 3 and
Layer 4. Each layer is compositionally and mechanically distinct from each other.
Layer 1 – Sediments: The thickness of sediments forming layer 1 of the oceanic lithosphere
increases with distance from the spreading ridges. The further the crust is from the ridge, the
older it is and the more time has elapsed for sediments to accumulate. Near spreading ridges,
sediments are no more than a metre or two thick, except where an ridge lies close to land (for
example in the Gulf of California). By contrast, in abyssal plain sediment thicknesses of a
kilometre or more are commonplace. The continental shelf-slope-rise region may be blanketed
by more than 10km of sediment. The type of material deposited near the mid-oceanic ridges is
sediment that settles from suspension in the open oceans and are known as pelagic sediments.
These pelagic sediments are dominated by the remains of skeletal and shelly hard parts of
planktonic organisms made of calcite or silica. Other materials that could be deposited in the
deep ocean areas include very fine wind blown material and volcanic ash. In contrast material
deposited near continental landmasses will be eroded clastic material.
Layer 2a - Basaltic Lavas: When basaltic lava erupts under water, the lava commonly forms
pillow-shaped bodies. The top of the oceanic crust is therefore characterised by pillow lavas perhaps the most abundant volcanic rock on Earth. Lavas that erupt under water cool more
rapidly because water has greater thermal conductivity and specific heat than air. This rapid
cooling forms the distinctive pillow-like shapes. As soon as the magma reaches the sea-floor, a
chilled skin forms on the outer surface, which is still flexible enough to change shape as the
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lava flows. As more magma arrives at the surface to feed the flow, pressure inside the skin
causes bulbous protrusions to form, which commonly detached from the main flow when they
reach a critical size, and form separate pillows. BY repeating this process, piles of pillows, up to
hundreds of metres thick, can develop, to produce a characteristic hummocky topography on
the sea-floor. Typical pillows range from about 10cm to over a metre in diameter, and up to
several metres in length. When completely solidified, they have dark glassy rims (the original
chilled skin) enclosing a more crystalline interior which cooled more slowly.
Layer 2b - Sheeted Dykes: Tension across the spreading ridge episodically produces a
vertical fissure, up which magma will escape from the underlying magma chamber. Any magma
that reaches the sea-floor will spread out to form lava flows that make up the topmost layer of
the oceanic crust. However, magma that solidifies within the fissure creates a vertical sheet of
rock known as dyke, which is typically about 1 metre wide. A new dyke is intruded after each
episode of extensional fracturing of the crust, so that the crust immediately overlying the
gabbro in the centre of the spreading zone consists of nothing but adjacent sheets of nearvertical parallel dykes, often described as sheeted dykes.
Layer 3 – Gabbro: The comparatively large size of the crystals in the gabbro layer shows that
they are a product of slow cooling. It was once thought that in most places the magma collected
in large, shallow magma chambers, several km in size, before crystallising. Accumulation of
crystals on the floor and walls of these chambers would give rise to the solid gabbro part of
the crust. However, it is now believed that, often, the magma has begun to crystallise during its
ascent from the source region, and arrives in the crust as a mixture of crystals and liquid. Also
it seems to arrive in small batches, which mix with the even more crystallised magma that
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arrived in earlier batches. Thus, in most places, the gabbro layer is a crystal mush, consisting
of crystals mixed with only about 10% of magma, and lacks any large bodies of liquid magma.
Layer 4 – Peridotite: This is the upper mantle part of the oceanic lithosphere that is too cold
to be part of the asthenosphere. Peridotite is an igneous rock consisting predominantly of the
mineral olivine.
2. Evidence for the Structure of the Oceanic Crust
Our knowledge and evidence for the structure of the oceanic crust and upper mantle (oceanic
lithosphere) comes from four main sources:

Geophysical techniques, notably seismic refraction, reflection and velocities, but also
magnetic and gravity surveys and heat flow measurements.

Examination and measurement of physical properties of rocks dredged from the
seabed and cored by drilling programmes from the upper parts of the crust.


Direct observation and photography of the seabed using submersibles.
Land-based studies of ophiolites.
Geophysical Techniques: It was seismic refraction studies which first revealed that the
oceanic lithosphere consists of layers whose seismic velocity increases with depth. The seismic
layers are numbered 1 to 4, and match up with the rock types present.
The following points summarise the layered structure of the oceanic lithosphere:
 The igneous layers (2-4) are formed by processes at spreading axes, and are progressively
buried by increasing thicknesses of sediments constituting layer 1, which eventually
become consolidated into sedimentary rock.
 Layer 2 and 3 constitute the igneous oceanic crust. The normal total thickness of these
layers is 6-7km. It may be less near transform faults and fracture zones, or more near hot
spots such as Iceland.
 The two most important changes in seismic velocity are at the base of the layer 1
sediments and at the base of layer 3. The latter is the Moho, where gabbro at the base of
the crust rests on peridotite at the top of the mantle.
 Within layer 2 and 3 there are no definite seismic discontinuities, but rather a general
increase of seismic velocity with depth. Subdivisions within layer 2 and the distinction
between layers 2 and 3 are defined by changes of velocity gradient (changes in the rate of
increase of seismic velocity with depth), and not by abrupt changes of seismic velocity.
 Despite their different appearances, newly formed rocks of layers 2 and 3 are very similar
in their overall mineralogical (and chemical) composition; in order of decreasing abundance,
the constituent minerals are plagioclase feldspar, augite and olivine. However, the crystals
forming the gabbro are larger than those in the dykes, which in turn are larger than those
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in the lavas. The top of layer 4 accounts for the seismic discontinuity of the base of layer
3 (the Moho). The principle minerals of the peridotite of layer 4 are olivine and augite.
 The subdivisions of layer 2 cannot always be recognised, but where they are, drilled
samples suggest the following explanation. In the sub-layer 2a, the lavas have many
fractures and voids. In sub-layer 2b, the lavas are generally less broken and the voids are
mostly filled with clays and other alteration minerals. This is partly because as soon as
oceanic crust is formed it becomes subject to chemical alteration and metamorphism, by
interaction with seawater. Thus, it seems likely that the seismic velocity distinctions
between layers 2a and 2b, and between layers 2b and 3, are often due to differences in
the intensity of alteration and metamorphism, and/or progressive closure of cracks and
pore spaces at greater depths and therefore greater pressure, rather than to differences
in rock type. Fractured lavas have lower seismic velocities than similar unfractured rocks
due to the presence of water. The speed of sound in water is less than that in rock. Water
will occupy fractures in rocks on the seabed, and so will reduce the speed of seismic
velocity through fractured rocks relative to unfractured rocks, in which there is no water
to slow the seismic waves.
 Seismic velocities increase more rapidly with depth in layer 2 than in layer 3. Seismic
velocities in layer 2 also tend to increase with age, i.e. with distance from the ridge, as the
remaining cracks and pore spaces become filled with minerals formed as a result of the
interaction between rock and seawater.
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Deep-sea Drilling Programmes: The Deep Sea Drilling Project (DSDP) was begun in the
1960's using the Glomar Challenger, a ship that was specifically designed for drilling into the
ocean floor beneath several kilometres of seawater. Many hundreds of holes have now been
drilled through the ocean-floor sediments into the underlying igneous rocks, and have added
greatly to our understanding of the structure of the oceanic lithosphere. The Glomar
Challenger was succeeded by a similar but better-equipped vessel, the JOIDES Resolution,
which began operations in the international Ocean Drilling Program (ODP) in 1985. Examination
of the oceanic crust has been further helped by the use of instruments lowered down drill
holes to measure the physical and chemical properties of the sediments and underlying volcanic
rocks without bringing them to the surface. Drilling has been supplemented by increasingly
sophisticated techniques such as underwater photography from cameras towed close to the
sea-floor.
Direct Observation: Since 1963, geologists have been able to undertake direct observations of
the sea floor using manned submersibles equipped (such as Alvin) with precise navigation
systems, sampling apparatus and cameras. Nowadays, remotely controlled submersibles are used
carrying cameras, side-scan sonar equipment and sampling gear. They can reduce the time taken
for deep-sea exploration and survey from years to weeks, because they can dive deeper and
stay down longer than manned submersibles. They are also cheaper and involve virtually no risk
to human life.
Ophiolites: Sometimes during subduction a piece of oceanic lithosphere is obducted onto the
overlying continental crust rather than being subducted beneath it. A piece of ocean floor
which has been preserved in this way is called an ophiolite or ophiolite complex. Worldwide,
ophiolites usually seem to have been formed at, or shortly before, the onset of continental
collision (mountain-building). Newly formed oceanic crust, while still hot and therefore less
dense, may be sufficiently buoyant to get caught up in the mountain-building episode
(orogenesis). Ophiolite complexes consist of thin layer of sedimentary rock overlying basic
lavas, dykes and gabbro, which in turn overlie peridotite. This sequence is similar in thickness
and structure to that inferred for oceanic lithosphere from seismic and drilling studies of the
ocean basins. Geologists therefore interpret ophiolites as being fragments of ancient oceanic
lithosphere. So not only do ophiolites give evidence for the structure of the oceanic lithosphere
they can also be dated radiometrically or by examining the fossils in the oldest associated with
them. The ophiolite forming the Troodos Mountains in Cyprus is one of the most intact and
best-known globally, while the Lizard in Cornwall is the best example in Britain.
However, although ophiolites are very useful a number of differences have been found between
the rocks in ophiolite complexes and present day oceanic lithosphere. The main differences are
due to either weathering or metamorphism. After ophiolites are obducted onto the continental
crust they come in contact with the atmosphere and the some minerals (e.g. feldspar) in the
gabbros, dolerites and basalts react with water and alter into clay (hydrolysis).
The peridotite can also be altered when the olivine minerals react with water to form
serpentine (clay-like mineral), forming a new rock referred to as serpentinite. As a consequence
the seismic velocities measured in rock samples from the bottom parts of an ophiolite complex
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are lower than in the corresponding parts of oceanic lithosphere due to the original minerals
being converted to less rigid and more compressible clay-like minerals such as serpentine (the
green appearance of these minerals was thought to be like that of a snake or serpent hence the
name) by the effects of water at some stage of an ophiolites' history. The term ophiolite,
which was introduced in the 1800s to describe serpentinites, was derived from the Greek word
ophi which means snake or serpent.
3. Formation of the Oceanic Lithosphere
The axes of the ocean ridge systems are the most active volcanic zones on Earth, where
spreading zones generate new oceanic lithosphere at rates of between 10 and 200km per million
years (1 - 20 cm/year). Volcanic activity at ocean ridges is episodically continuous for much of
the lifetime of an ocean basin, which may be as long as 200 million years. The process that
forms oceanic lithosphere is known as sea-floor spreading. Although much still remains
uncertain about the sea-floor spreading process, it has become apparent that it is the diverging
motion of plates at oceanic ridges that causes the igneous activity, rather than the igneous
activity that forces the plates to move.
As two oceanic plates are pulled apart (perhaps due to the forces of slab pull during
subduction) the lithosphere is stretched and thinned. Beneath the thinning lithosphere the
asthenosphere is drawn upwards so as to prevent a gap opening and fills the space where the
lithosphere once was. As the asthenosphere rises past a depth of a few tens of kilometres, the
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steadily decreasing pressure triggers up to 10% partial melting, a process known as
decompression melting. Although the material that is being partially melted has the
composition of peridotite, the resulting magma has the composition of basalt. This basaltic
magma is less dense and more mobile than the unmelted mantle, so it rises upwards into the
fractures in the brittle lithosphere and either erupts onto the sea-floor or cools and solidifies
in the fractures as dykes, adding new material to the oceanic lithosphere.
The theory of sea-floor spreading was tested by the use of palaeomagnetism. From the mid1950‘s magnetic surveys using magnetometers towed behind ships built up a substantial
database of the intensity of the present-day magnetic field across the oceans. The magnetic
intensity depends entirely on the magnetic properties of the crustal rocks under the oceans,
and this in turn depends largely on the mineralogy of those rocks. The oceanic crust is made of
basalt which is composed of magnetic minerals such as magnetite and would be expected to
produce a significant local magnetic anomaly. But basalt is basalt (or so it was thought!), and the
investigators were expecting uniform magnetic properties across the oceans. What they
discovered was that the magnetic field was anything but uniform, instead containing an
unexpected pattern of linear magnetic ―stripes‖. These defined long, linear zones of rapid
anomaly change, bordering an area of relatively uniform positive anomaly on one side and an
area of relatively uniform negative anomaly on the other. These stripes could be traced
individual anomalies for hundreds of kilometres along the ocean floor, until they stopped at
prominent fault-like zones. The pattern of anomalies did not actually stop at these cross-faults,
now known as transform faults, but appeared to be displaced, sometimes by distance of over
1000km. The anomaly pattern was also symmetrical across the mid-ocean ridges.
These stripes represent variations in the direction and amount (or both) of the Earth‘s
magnetic field. They can be explained if the magnetic field has been subjected to frequent pole
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reversals over time. Each positive anomaly represents a time during which a magnetic field of
one polarity was established. When the polarity changed, the field would rapidly decay, to
become re-established shortly afterwards with the opposite polarity, giving a corresponding
stripe of negative anomaly. Therefore, the ocean-floor record must represent magnetic field
reversals, but how could the symmetry and linear nature of the magnetic reversals be
explained?
In 1963, two British geoscientists, Fred Vine and Drummond Matthews, proposed an
hypothesis which elegantly explained magnetic reversal stripes, linking these to sea-floor
spreading and laid the foundation for plate tectonics. They suggested that new oceanic crust
was formed by the solidification of basalt magma extruded at mid-oceanic ridges. As it cooled,
the lava acquired a magnetisation in the same orientation as the prevailing global magnetic field.
As the forces that drive sea-floor spreading continued to operate, the ocean margins would
move apart and more oceanic crust would be generated at the ocean ridge. If the magnetic
field reversed, then newly erupted basalt would become magnetised in the new magnetic field
which would be significantly different from the pre-existing one. Not only would the new
oceanic crust show a different magnetic polarity from its neighbour, but it would also be
slightly younger. As new oceanic crust continued to form, so stripes of basalt of the same age
and magnetic polarity would spread away from the ridge. Ocean crust can therefore be
considered as the ―tape‖ in a global ―tape recorder‖, in which the reversal history of the Earth‘s
magnetic field was constantly recorded.
4. Calculating Rates and Directions of Sea-floor Spreading
There are a number of ways relative plate motion can be estimated at plate boundaries.
Spreading occurs away from an ocean ridge at constructive plate boundaries, so it is possible to
refer either to the movement of one plate away from the other plate (the spreading rate), or
to the movement of one plate away from the ridge (the half-spreading rate). Both of these are
relative rates of movement, since they measure the movement of one side of a constructive
boundary relative to the other, or to the ridge. The average rates of plate movement are given
in mm/year (i.e. distance/time). So to calculate the rate of sea-floor spreading you need to
know two things; the distance the plate has travelled and the age of the sea-floor. The distance
from the ridge axis or distance a volcanic island is from a hot spot can be measured from maps
produced by surveys from ships or satellites (using GPS for example). However, the precise
location of ridge axis cannot be accurately measured to metres but to within 1km. Spreading
rates are more accurate closer to the ridge. The time component is measured using radiometric
methods of dating.
Radiometric dating is the most reliable method of obtaining an age for a rock. This absolute
dating method depends on being able to measure the tiny amount of radioactivity trapped in
some rock, especially igneous rocks. Some elements are unstable and thus known as radioactive.
An example is the element potassium (K). Most potassium atoms are stable and non-radioactive.
Each of these contains a total of 39 heavy particles (protons and neutrons) in its nucleus, and is
described as potassium-39. However, about one potassium atom in every 10,000 contains an
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extra neutron, making a total of 40 heavy particles. These atoms of potassium-40 are
unstable, and undergo radioactive decay.
Once a mineral crystallises into a solid, any elements within it that are unstable start to break
down to form a new product. This is known as radioactive decay. The steady rate at which a
radioactive element decays can be measured in the laboratory. The rate of decay is given in
half-lives. The half-life is the time taken for half of the original radioactive elements to decay
into a new product. In the example above, potassium-40 breaks down to form the inert gas
argon (argon-40). This is known as potassium-argon dating. The time it takes for half the
original unstable parent element (K) to from to the new stable daughter element (Ar) is about
1,300 million years!
By measuring the relative proportions of the surviving parent elements and its decay products
trapped within the same rock or mineral, the time since the rock or mineral was formed can be
calculated. So in the case of K-Ar decay, which has a half-life of 1,300 million years, if a rock
showed it contained 50% of the original parent element and 50% of the new daughter element
then the rock would be 1,300 million years old.
However, having explained all that, K-Ar dating is very rarely used these days as it is widely
regarded as being very inaccurate particularly for MOR basalts which may have interacted with
seawater. The modern methods for dating ocean floor basalts are 40Ar-39Ar dating (if the
rocks are older than~2Ma) or U-Th-series disequilibria dating (using short lived intermediate
isotopes in the Uranium-Thorium decay series) for more recent eruptions. (This is not to be
confused with U-Pb dating which is only suitable for very old rocks). Another issue relevant to
all methods of radiometric dating is the level of accuracy. At present the absolute age of a
rock specimen is accurate to +/- 2%, depending on the method used.
Another method of measuring the time component needed to work out the rates of plate motion
is by magnetostratigraphy. Magnetostratigraphy is a technique used to date volcanic sequences
such as ocean floor basalts which are not easily sampled. The method works by collecting
igneous rock samples from more readily accessible locations such as the layered piles of lava
from land volcanoes. In Hawaii for example, geologists were able to sample each successive
layer of lava, date it radiometrically and then work out its magnetism using a magnetometer.
Palaeomagnetic results from these lavas showed that through geological time the Earth's
magnetic field direction has not been constant, but has periodically reversed in direction
(―flipped‖). In fact the field has reversed polarity many times. If the ancient magnetic field
was oriented similar to today's field (North Magnetic Pole near the North Rotational Pole) the
strata retain a normal polarity. If the data indicate that the North Magnetic Pole was near the
South Rotational Pole, the strata exhibit reversed polarity. Putting all the data together they
were able to date the age of magnetic reversals in the layers reaching back 4.5 million years.
From this a Global Magnetic Polarity Time Scale was developed showing all the magnetic
reversals over this period.
As mentioned previously, because of the vast area and considerable depth of the world‘s oceans
it is no easy task to collect samples of rock from the sea floor to be radiometrically dated.
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However, the development of magnetostratigraphy has made this task much quicker and easier.
Since the 1950s geologists have been towing magnetometers from ships all over the world‘s
oceans. These magnetometers detect the parallel bands of strong and weak magnetism in the
basaltic ocean floor. Lava preserves the direction of the Earth‘s magnetism at the time of
eruption because tiny magnetite minerals act like compass needles and are frozen in place when
the basaltic lava solidifies. When the magnetic field of each of these minerals is pointing in the
same direction as present day‘s magnetic field the total magnetic field is stronger than normal,
forming a positive magnetic anomaly. When the magnetic field of the oceanic crust is opposite
to the polarity of the present day magnetic field, the total intensity is weaker than normal,
forming a negative magnetic anomaly. So the alternating bands of sea-floor magnetism are the
records of past reversals of the Earth‘s magnetic field.
When the pattern of magnetic banding from the ocean floor is compared to the Global
Magnetic Polarity Time Scale the match is found to be almost perfect. The most recent
reversals occur at the mid-oceanic ridge, and progressively older ones are found progressively
further away toward the flanks of the ocean basins. Reversals of short duration are matched
by relatively narrow bands and those of long duration by broad bands. Because we know the age
of the magnetic reversals from radiometrically dating terrestrial lavas we can determine the
age of oceanic crust from the match between the magnetic bands and the Global Magnetic
Polarity Scale. The velocity of sea-floor spreading can be calculated by measuring the distance
from the ridge crest to a given magnetic band and then dividing it by the age of that band as
determined by its position in the Global Magnetic Polarity Scale.
Magnetic Anomalies: A simple method for estimating the average rate of sea-floor spreading
ocean ridges can be developed by using the width of magnetic anomalies of the same age
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between two ocean basins. Wide magnetic anomalies suggest a fast rate of sea-floor spreading
whereas narrow magnetic anomalies suggest a slower rate.
Where ocean-floor magnetic anomalies of known age (either by radiometric dating or
magnetostratigraphy) can be identified at a measured distance from a mid-oceanic ridge then
the relative rate of plate motion can be determined. If ocean-floor magnetic anomalies of
known age can be identified at a measured distance apart and on opposite sides of a
conservative plate boundary, they can be used to determine a rate of relative motion between
two plates. These methods are best applied to oceans that have formed during periods
measured in tens of millions years.
In the example above the half spreading rate is 20mm/year, and is calculated from the fact
that rocks originally at the ocean ridge have moved 100km in 5 million years. This works out at
20km every million years which equates to 20 000 000mm (or 20mm/year)! The full spreading
rate is the two half spreading rates added together i.e. 40mm/year.
Ocean-floor Depth: Ocean ridges have an average width of around 1 000km and rise above the
surrounding abyssal plains by 2 to 3km. The ocean floor is therefore higher at constructive
boundaries than it is further away from the ridge towards the continents. This is due to the
hot magma erupting at ocean ridges which cools, contracts and becomes denser with time. The
ocean floor therefore "sinks" as it moves away from the ridge. So the depth to the ocean floor
can be used to give a rough age of the oceanic crust itself. Therefore if the distance from the
spreading ridge, and the depth to a particular piece of ocean floor are both known, then the
average rate of plate motion relative to the ridge can be calculated.
Global positioning by satellite (GPS): The rate of relative plate movements at the present day
can be measured extremely accurately by using modern space technology. This, of course, gives
us no insight into plate movements in the past. GPS uses several high-altitude satellites with
orbital periods of 12 hours, and each satellite broadcasts its position and time. When multiple
satellites are tracked, the location of the receiver can be estimated to within a few metres on
the Earth‘s surface. However, accuracy of the results depends on many factors including the
length of time over which measurements have been accumulated. To reach accuracies of 1 –
2mm for sites that are thousands of kilometres apart requires many years of data
accumulation.
So far we have been concerned with the relative rates of motion between plates. This does not
tell us anything about the actual motion of plates with respect to some fixed point on the
Earth. Just because one plate appears to be moving relative to another, we cannot be sure
whether both plates are really moving or whether one is actually stationary and the other plate
is moving! Magnetic anomalies tell us that some type of movement is taking place, but not
absolute or true plate motion.
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Hot spots: Hot spots
represent
the
surface
expression of a narrow
heat source consisting of
material flowing upwards
from deep in the mantle
and are thought to be
fixed in relation to the
moving lithosphere above
it. Hot spots can therefore
be used as a fixed frame
of reference against which
to
measure
relative
movements of plates. This
should then give us the
true or absolute movement
direction and speed of each
of the Earth's plates.
In this example it has taken the last 43 million years for the plate to move 3000km. This
equates to about 70km in 1 million years or 70,000,000mm (a rate of 70mm/year). In the
previous 35 million years (78 - 43 = 35) the plate has moved about 2000km, which is about
57km in 1 million years or 57,000,000mm (a rate of 57mm/year). Therefore, the absolute rates
and direction of movement for the Pacific Plate are 57mm/year in a northerly direction
between 78 and 43 million years ago, and 70mm/year in north-westerly direction in the last 43
million years.
Identifying the Hawaiian hot spot as a stationary features detached from the moving Pacific
Plate is critical in calculating the true movement amounts and directions of the Pacific Plate.
How do we know that the Hawaiian hot spot is stationary? If the hot spot itself has moved its
location within the Earth over the past 45 Ma, then our calculations are incorrect. We can
approach this using data from other hot spots. If we can recognise another hot spot on the
Pacific Plate which has a chain of volcanic islands like Hawaii, and if we can calculate the
velocity of the Pacific Plate over the past 45 Ma from this new information, we can compare
the rate and direction of movement for the Pacific Plate. If the two answers are the same,
then either the two hot spots have been stationary over the past 45 Ma or they have both
moved relative to the Pacific Plate by exactly the same amount and direction over that time. If
the two answers are different, then one hot spot has definitely moved relative to the other,
and we cannot be sure which one is actually moving. If we can find three, four, five or more hot
spots within the Pacific Plate and they all give the same results, we can become more confident
that the hot spots are stationary and that we know the true motion of the plate.
In 1972, an American geologist noted the similarity of the orientation of the chains of major
volcanic islands and seamounts on the Pacific Plate. He noticed that the geometry of the
Hawaiian-Emperor and the Austral-Marshall Island seamount chain shown in the map below can
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be matched to the expected by moving a rigid Pacific Plate over the hot spots fixed at the
present locations of Hawaii and the Macdonald seamount, respectively.
Subsequent studies have extended that early work by incorporating the distribution of the
ages of the volcanoes as well as the geometry of the volcanic chains, and these confirmed the
fixed relative positions of two of these Pacific hot spots over the past 45 Ma. The coincidence
of two hot spots giving the same result for the Pacific Plate strictly means that these hot
spots have not moved relative to each other. However, we cannot eliminate the possibility that
they are coupled together in some way, and drifting as a group of hot spots!
5. Destruction of Oceanic Lithosphere
Oceanic lithosphere cools by conduction as it moves away from the oceanic ridge, leading to a
thickening of oceanic lithosphere. Oceanic lithosphere is reabsorbed into the mantle by sinking
at subduction zones. The subducting lithosphere however does not generally melt unless it is
quite young (< 2-3 million years) and therefore hot. It is the dewatering of the slab which
lowers the melting point of the overlying mantle wedge and causes partial melting of the
asthenosphere. As the old, cold lithosphere descends into the asthenosphere less dense
minerals such as water are lost due to dehydration which further cools the descending slab
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further increasing its density. So the older and colder the oceanic lithosphere the denser it will
be and the faster it will subduct. Also at depth phase changes occur in the slab converting some
minerals into denser varieties. At about 400km depth olivine transforms into a form in which
the atoms are packed more closely, known as spinel, with a consequent increase in density of
about 10%. There is further phase change at about 670km where the spinel structure
transforms into an even-higher pressure form known as perovskite with another density
increase of about 10%. These increases in density all increase slab pull forces, helping the
oceanic lithosphere to subduct, as does the steepness of the descending plate (the steeper the
subducting plate the quicker it subducts). These subducted slabs descend to 700km or
sometimes all the way to the core-mantle boundary where they eventually melt and rise up again
as mantle plumes to form hot spots.
6. Evolution of Ocean Basins
The oceans of the world have been forming since the beginning of the geological record.
However, the shape of most past ocean basins has to be worked out from observations of
remains preserved in continental areas. That is because ocean basins are relatively short-lived
features on Earth: no oceanic crust older than 180 million years is known from the present
oceans. A complete cycle of ocean basin formation, opening and closure seems to take about
400 million years (about 200 million years to open, about 10 million years to switch from a
passive plate margin to a destructive plate margin, and about 200 million years to close).
The evolutionary sequence of an individual ocean basin growing from an initial rift, reaching a
maximum size, then shrinking and ultimately closing has been termed the Wilson Cycle after J.
Tuzo Wilson the Canadian geologist who put forward the idea in the 1960's. The cycle starts
off with a continental craton which has probably been stable for thousands, if not millions of
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years, with no associated tectonic activity. Then a mantle plume may develop at the core-mantle
boundary and initiate uplift and rifting above it in the continental crust as anomalously high
heat flows interact with the lithosphere.
Stage 1: Embryonic Ocean
During crustal extension, the ductile lower part of the crust is stretched, but the brittle upper
part is fractured and rifted. Blocks of crust slide down fault planes, and sediments accumulate
in the lakes and valleys which occupy the resulting depressions. The whole area is uplifted as
well as extended due to the heat form the mantle plume the probably initiated the rift in the
first place. An example of a continental rift system is the East African rift valley, here a
down thrown rift valley also has igneous activity (basaltic) associated with the underlying
mantle plume partially melting the asthenosphere.
Rift systems can be classified into active and passive types. Active rifts are formed by
processes in the mantle below the lithosphere related to mantle plumes. Hot, buoyant mantle
plumes cause doming and uplift of the lithosphere above. This doming causes the lithosphere to
thin due to extension and erosion of the uplifted area. This thinning causes partial melting of
the underlying asthenosphere (decompression melting) generating basaltic magma. This form of
rifting causes uplift and a volcanic basin, with little amounts of sediment deposited but large
amounts of volcanic material. Examples of active rifts include mid-oceanic ridges and back-arc
basins. In contrast, passive rifts are caused by stresses within the lithosphere which is
stretched and thinned by plate forces. This form of rifting causes subsidence and a
sedimentary basin flanked by faults, with large amounts of sediment deposited and little
volcanism.
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Stage 2: Young Ocean
When separation occurs, basaltic magma rises to fill the gap between the two continental
blocks. Because the new oceanic crust is both thinner and denser than continental crust, it lies
below sea-level. Initially, the young marine basin is fairly shallow. If repeated influxes of
seawater become evaporated, salt deposits (evaporites) will accumulate. Otherwise, there will
be normal marine sedimentation of muds, sands and limestones, depending on local conditions.
One of the clearest examples of a young ocean basin is the Red Sea. About 35 million years ago
(Mya) a crack formed on the continental crust between what is now Africa and Arabia, and by
25 Mya east-west extension had begun to be felt across the entire area of what is now the Red
Sea, and developed into a rift system within the continental crust. Along this Red Sea rift, the
continental crust continued to stretch and subside, and evaporites were deposited on top of it.
About 5 Mya the system was reactivated and as the continental crust was already so fully
stretched that rather than thinning and stretching even further it was pulled apart, and seafloor spreading began here for the first time. The Red Sea opened as a new ocean due to the
formation of a constructive plate margin. The new basaltic ocean crust that develops has
palaeomagnetic strip patterns within it due to iron minerals in the magma aligning with the
magnetic field present when it was formed. These areas also have high heat flows just like midoceanic ridges and develop narrow seas with parallel coasts and a central depression. The Red
Sea is of particular interest to geologists as a "living experiment", as the final stages of
continental extension and the beginning of sea-floor spreading can be observed in this one
location. At the northern end of the Red Sea, the basin floor consists of thinned continental
lithosphere that has been injected by numerous basaltic dykes, indicative of the very last
stages of continental extension prior to break apart. This contrasts with the southern end,
where volcanic sites are associated with rifting and the formation of new oceanic lithosphere
produced by sea-floor spreading.
Stage 3: Mature Ocean
The Atlantic is an excellent example of a mature ocean. Here spreading is dominant due to the
process of sea-floor spreading, and no subduction. These ocean basins have a very active
oceanic ridge in the middle of the ocean with the characteristic symmetrical pattern of
palaeomagnetic stripes either side of the ridge and very high heat flows.
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Stage 4: Declining Ocean
Stage 4 involves the development of one or more destructive plate margins. By the time oceanic
lithosphere is about 200 million years old, it is so cold and dense compared with the adjacent
continental lithosphere that it starts to sink into the underlying asthenosphere under its own
weight (helped by the continual push of newly formed oceanic lithosphere at the mid-oceanic
ridge). This starts the process of subduction. Recent research suggests that it can take as
little as 10 million years for subduction to be initiated. The declining ocean basin like the
Pacific Ocean is both spreading and shrinking. The Pacific has an active spreading centre known
as the East Pacific Rise, but it is not in the middle of the ocean anymore due to the speed of
subduction under the South American and North American Plates. In fact the Pacific has
subduction zones surrounding the entire ocean basin and the fast rates of spreading are due to
the process of slab pull (cold, dense oceanic crust pulling the oceanic plate with it due to
gravity).
Stage 5: Terminal Ocean
The Mediterranean is an ocean in the final stages of its life, with the African Plate being
consumed under the Eurasian Plate. Unless the world system of plates changes so as to halt the
northward movement of Africa relative to Europe, the continental blocks of Europe and Africa
will eventually collide, and new mountain ranges will form.
Stage 6: Relict Scar
The end of the Wilson Cycle occurs when two continental crusts collide at a collision zone. The
Indus suture in the Himalayas is an example of where two continental plates have collided
destroying the Tethys ocean basin in the last 40 million years. The two continental blocks are
too light to be subducted so thrust faulting, intense deformation and uplift occurs.
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Over millions of years these uplifted mountains are worn down and a stable continental craton
forms which may well be rifted in the future to start the cycle off all over again.
The oldest crustal rocks are found on the continents. The oldest continental rocks are around
4,000 to 3,800 million years old, and are on average 1,800 million years old. These dates have
been worked out from radioactive dating techniques. In comparison the oldest oceanic rocks
are only 170 million years old, and are on average only 65 million years old.
The reason for the vast differences in ages is due to the fact that oceanic crust has a similar
density to that of the mantle and can be subducted and destroyed. The continental crust on
the other hand is much less dense than the underlying mantle and is too light to be pushed
down. So once formed continental crust is hard to get rid of and explains why it is much older
than the oceanic crust. The only way to destroy the continental crust is by surface erosion. But
even here when eroded sediment is dumped onto the sea-floor of the deep oceans it is
eventually returned to the continents as it is scraped off at subduction zones to form an
accretionary wedge or prism.
4. Formation and Deformation of the Continental Crust
There are three groups of large-scale features found on continental lithosphere: cratons,
subduction zone orogenic belts and collision zone orogenic belts. All three groups are linked to
plate tectonics in their formation.
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1. Cratons
Cratons are areas of continental lithosphere that have been tectonically inactive or almost
inactive for hundreds of millions of years. They are named after the Greek word kratos
meaning "strength". The biggest cratons are several thousand kilometres across, though they
are highly variable in size and shape. They are usually situated well away from present-day plate
boundaries. They are stable areas - cratons show no evidence of tectonic activity over long
periods of geological time.
Typically, they contain a succession of nearly horizontal sedimentary layers, resting
unconformably on ancient metamorphic rocks. Over vast area, these sedimentary rocks are not
significantly folded or faulted, and frequently not even tilted. The succession of strata may
record deposition over many hundreds of millions of years.
The largest craton is the Canadian or Laurentian Shield, which makes up most of the eastern
half of Canada from the Arctic Ocean to the Great Lakes. Most of the rest of North America
is younger rocks that have become accreted around the craton as a result of subsequent
deposition of sediments, volcanic activity and plate collisions. Each of today's major continents
has a cratonic element within it little changed over the last 3 billion years. Australia has the
Pilbara craton and the Yilgarn craton to its south, Africa has the Kaapvaal and Zimbabwe
cratons in the south and several others further north and west, Europe has the Baltic shield,
Asia's largest is the Aldan shield in the far south-east of Russia, and South American cratons
occur in Brazil and Argentina.
In plate tectonic terms, cratons show that deformation is rare well away from active plate
boundaries. Cratons predictably show few tectonic structures. Major changes only happen if
the plate is disturbed by some event sourced from outside the plate itself, like for example the
effects of a mantle plume.
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2. Subduction Zone Orogenic Belts
The continental lithosphere is covered with long, narrow belts containing interlocking
sedimentary and metamorphic rocks, and igneous intrusions. These belts are characterised by
uplift, and by complex structures such as major thrust faults and large-scale folds with
cleavages. Typically, such belts form mountain chains, or at least the eroded remnants of
former mountain chains. They are known as orogenic belts from the Greek word oros meaning
"mountain" and genesis meaning "production". These are two different types of orogenic belt subduction zone orogenic belts and collision zone orogenic belts.
The diagram above shows the main large-scale features of a subduction zone orogenic belt.
Forebulge and Trench: The forebulge is a broad-scale upward bulging of the strong oceanic
lithosphere as it bends into the subduction zone. The downbend of the rigid oceanic lithosphere
forms an oceanic trench between the relatively high area of ocean floor either side of it
(forebulge and fore-arc ridge). These oceanic trenches are often thousands of metres deep,
with the deepest being the Mariana trench at 11 033m below sea-level.
Accretionary Prism: On the leading edge of the crust overriding the subduction zone, just
landward of the trench, there is often a wedge-shaped pile of material, mostly soft sediment,
that has been scraped off the downgoing plate. This is known as the accretionary prism or
accretionary wedge. The reason for these accumulations of material at subduction zones is
partly that young pelagic sediment (remains of dead marine organisms) is easily scraped off the
oceanic plate, and partly because the subduction trench area is a natural sediment sink of trap
for material eroded off nearby landmasses. Sediments from these two different sources (deep
ocean or the arc margin) are of contrasting character. The thin layer of sediment supplied
continuously by the subducting oceanic plate is dominated by fine-grained calcite oozes or clays
deposited in deep-sea environments. In contrast, sediment from the overriding plate is both
coarser and more voluminous than the pelagic sediment. The eroded volcanic arc material is
dominantly clastic material, much of it transported by high-energy turbidity currents plunging
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down towards the subduction trench. The resulting thick, graded deposits (greywackes) blanket
the sea-floor towards the rear of the prism, filling in many small basins that form on top of the
prism.
Fore-arc Basin: As the downgoing plate scrapes under the leading edge of the overriding plate,
it drives up the surface to create a fore-arc ridge (often forming a chain of islands between
the trench and the volcanic arc). The rising fore-arc ridge, plus down-flexing of the overriding
plate's leading edge, creates a fore-arc basin in which sediments derived from the volcanic arc
accumulate.
Volcanic Arc: Most of the Earth's active volcanoes on land are found at subduction zones,
where they are arranged in long chains lying parallel with the plate boundary. Typically, the arc
has a well-defined row of prominent volcanoes, known as the volcanic front, on average 100 300km from the plate boundary. This volcanic front is situated about 100km above the
subducting slab. Arc volcanoes exist due to the effect of water, released from the descending
oceanic lithosphere, lowering the melting point of peridotite in the overlying mantle wedge. Arc
volcanoes erupt a range of magma compositions, forming a basalt-andesite-dacite sequence
depending on how far the magma has evolved due to fractional crystallisation. The magmas at
subduction zones are distinctive in containing more water and more silica as they differentiate,
and therefore often erupt more explosively than magmas form other plate tectonic settings.
Back-arc Basin: Behind many subduction zones is a region that typically has lower elevation
than the arc itself, forming a basin. These regions are therefore referred to as back-arc
basins. There is considerable variety in these areas, not least the fundamental division between
those occurring within oceanic crust and those underlain by continental crust. Back-arc basins
are common in oceanic settings, especially in the western Pacific Ocean where many of which
are still active. Active back-arcs owe their existence to partial melting in the mantle and the
formation of new crust very similar to mid-oceanic ridges.
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The basic theory for the formation of an active back-arc suggests that a pre-existing volcanic
arc split down the middle and new oceanic crust was formed in the resulting spreading centre,
and the two halves have gradually separated. One half (the remnant arc) becomes inactive,
while the other experiences continuing magmatism, remaining an active arc. Due to the fact
that volcanic arcs have a high heat flow due to the production of magma this area of the crust
is relatively weak (remember the higher the temperature the weaker the material and the
cooler the material the stronger it is?). This area of weakness in the crust will tend to be
exploited first if tectonic forces put the region under tension, with ductile deformation acting
to thin the crust until a spreading centre and back-arc basin is formed.
There is no consensus yet on the driving force causing back-arc spreading itself, and a number
of theories have been put forward. However, many geologists believe that trench suction
initiates extension in the overriding plate, which is then further weakened by the ascent of
magma from the mantle. Trench suction is the gravitational force exerted on a plate above a
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subduction zone that literally "sucks" the plate into the oceanic trench causing the plate to
stretch and thin. The back-arc regions of continental areas are much less distinct, and much
more complex, than their oceanic counterparts, but again are probably due to lithospheric
extension
Paired Metamorphic Belts: The association of subduction zones with volcanic arcs has led to
the concept of paired metamorphic belts across subduction zone orogenic belts. This simply
highlights the juxtaposition of two linear regions of the same age with very different
metamorphic character. The country rocks in the volcanic arc above the subduction zone are
affected by large numbers of igneous bodies and intrusions which heat the crust on a regional
scale causing crustal temperatures to be higher than normal at shallow depths. Consequently
the country rocks will experience low pressure, high temperature regional metamorphism. In
contrast, where sediments have been dragged down into a subduction zone at the base of the
accretionary wedge they will experience particularly high pressures, but having been recently
at the Earth's surface they will not as hot as most rocks at the same depth. They will
therefore experience high pressure, low temperature regional metamorphism. When a paired
metamorphic belt such as this is found in ancient rocks, it is a good sign that there was once a
subduction zone there.
3. Collision Zone Orogenic Belts
Destructive plate margins may undergo a long history of oceanic plate subduction (such as the
present western margin of South America), but eventually they are terminated by continental
collision. For example, the northward subduction of thousands of kilometres of Indian oceanic
crust was terminated by the arrival and collision of the Indian subcontinent and formation of
the Himalayan orogenic belt. Where the convergence of two or more plates brings blocks of
continental crust into contact with one another it is known as a collision zone orogenic belt.
Collision orogenic belts may be young and mountainous like the Alps or Himalayas, or old and
relatively flat like the Highlands of Scotland, but all orogenic belts are largely made up of
intensely deformed and metamorphosed rocks.
The diagrams on the next page show a section through an idealised collision zone between two
major continents at two stages during its evolution. In the first diagram, the large-scale
features of a subduction zone orogenic belt before collision can be identified. Oceanic
lithosphere is being subducted beneath the upper plate and a volcanic arc is formed. The active
continental margin of this overriding plate is marked by an accretionary prism composed of
slices of oceanic sediments that have scraped off the descending plate. The passive continental
margin of the downgoing plate is marked by extensive shelf sediments lying in the ocean.
Eventually (as seen in the second diagram on the next page), the leading edge of the continent
on the downgoing plate enters the subduction zone but the buoyancy of the continental crust
prevents subduction from continuing. In order to even partially subduct, material must deform
extensively but, once subduction has ceased, the energy of the system has to be dissipated by
other processes. The result is further deformation that affects both plates. This deformation
231
starts at the suture line between the two plates but, over time, it migrates out to affect more
and more of the lower plate. Crustal shortening is accompanied by significant thickening, and
this thickened lithosphere is isostatically unstable, resulting in topographic uplift (i.e.
mountains). Variations in temperature and pressure through the collision zone result in
different types of deformation. In the upper crust, deformation will be dominated by the
processes of fracture and thrust faulting (brittle deformation). The higher temperatures and
pressures in lower crustal levels favour structures such as folds and shear zones which are
examples of ductile deformation.
232
Foreland: The foreland represents the undisturbed continental crust in front of a collision
zone. The foreland often forms a basin, known as foreland basin. These form as the rigid
lithospheric plate flexes to accommodate the weight of the mountain belt behind it. Although
the foreland is not directly involved in the deformation that produced the mountain belt the
huge amounts of material thrusted on top of each other in the Inner and Outer Zones can
depress the crust over a region beyond the actual limits of the collision zone. Such a
phenomenon is called non-local isostatic compensation - the effect of the load goes beyond its
bounds. Once a basin forms it becomes a natural sink or trap for material being eroded off the
newly uplifted mountain range.
Outer Zone: The Outer Zone is characterised by deformation of rocks at high crustal levels.
The Outer Zone has less intense deformation due to the fact that it is at some distance from
the suture zone and has therefore been deformed for a shorter period of time. Typically the
rocks in this zone are sedimentary rocks formed in shallow marine environments which have
been detached from the underlying crustal rocks and moved as a series of thrust sheets over
the foreland area. Thrust zones such as these are typical of the Outer Zone of orogenic belts
and are known as foreland thrust belts. Thrusting is an example of brittle deformation and is a
consequence of the low temperatures in the upper crust. The low temperatures and pressures
also mean the rocks in the Outer Zone show low-grade metamorphic change such as a weak,
sporadically developed cleavage. Granite plutons are rare or absent. The diagram above is a
cross-section through the Outer Zone of the Alps, near Grenoble in France, where the Eurasian
plate to the right of the diagram has collided with Italy which is the landmass to the left of
the diagram.
Shortening has been achieved by folding and thrust faulting. The displacement on individual
thrusts is at least 2km and the thrust planes dip towards the Inner Zone of the orogenic belt.
The folds are asymmetric, open, steeply inclined and gently plunging. These folds are thrust
ramp folds formed above a thrust plane as it moves up and over the rocks below. The principal
stresses that produce thrusting are also shown.
233
So, any folding that does occur in the Outer Zone is closely related to the thrusting. These
folds tend to be asymmetric, open, steeply inclined and flexural (parallel) in nature. The folding
of sedimentary rocks under low temperature and brittle conditions commonly produces flexural
folds (also known as parallel folds). The characteristic feature of flexural folds is that the
folded layers undergo little or no change in thickness, as they are traced around the foldhinges and or flanks. The geometry of flexural folds requires that they should change in shape
as they are traced up and down their axial plane. The outside of each antiform or synform
becomes bigger whilst they become tighter and more compressed in the opposite direction,
eventually forming simple cusps in the innermost part of the fold. The thickness of the most
competent (strongest) layer determines the fold wavelength. Flexural folding of sedimentary
rocks occurs by the process of flexural slip. This is the movement along the boundaries
between the layers. So flexural slip folding occurs wherever sedimentary layers slip over one
another along bedding planes. This type of folding is very common in competent rock types such
as limestone and sandstone.
Inner Zone: The Inner Zone of orogenic collision belts typically show complex ductile
structures in high-grade metamorphic rocks. Generally, these rocks have suffered repeated
phases of progressively more intense deformation, usually over a prolonged period of time and
at elevated temperatures and pressures associated with deep crustal levels. The continental
crust has been severely shortened and thickened. Every major collision zone orogenic belt has
its own distinctive organisation. However, several features seem to be common to the inner
zone of many orogenic belts.
234
First, the rocks involved are often much more varied than those in the Outer Zone, including
deep and shallow marine sedimentary deposits, ophiolite sequences, andesitic volcanic rocks and
plutonic igneous rocks. Secondly, there are repeated episodes of intense deformation as the
crust is shortened. High temperature conditions in the deeply buried rocks allows ductile
deformation such flow folding and shear zones to develop. The folds are commonly asymmetric,
tight to isoclinal, flat-lying to recumbent, and usually show axial plane cleavage.
Huge nappe structures form in the sediments first to be deformed close to the suture zone,
causing overturned stratigraphy. Often many separate phases of deformation occur which
develops minor folding on major folds, a feature known as parasitic folding.
Flow folds form under ductile conditions and have the same curvature on the inner and outer
surfaces of the folded layer, and are consequently also known as similar folds. Each layer is
thinner on the flanks and thicker on the axis of the fold. These folds extend indefinitely as
space is not a problem, unlike in flexural folds. Flow folds form due to movement within a
sedimentary rock layer rather than between the layers as in flexural folds. The high
temperatures exceed the yield strength of the rock and the minerals start to shear past each
(or in other words flow) and develop cleavage. This movement within the rock allows the whole
layer to bend and fold. Flow folds are common in incompetent rocks such as shale and clay
because their internal structure is weaker and less able to handle the pressure.
235
Thirdly, the metamorphic rocks in the Inner Zone indicate high to extreme pressures and
medium to high temperatures during deformation. Vast granite plutons are extremely common.
These may be formed by the very high temperatures in a collision zone causing sedimentary
material to be melted as a consequence of deep burial, or even due to decompression melting as
deeply buried material is brought nearer to the surface as erosion rapidly wears away the
uplifted rocks.
Lastly, brittle faults are much less common than ductile folds and cleavages. However, interior
zones frequently show large thrust faults that are sometimes associated with immense
recumbent folds.
Hinterland: The hinterland is the area beyond the suture line of a collision orogenic belt. Here
obducted ophiolite sequences may be present, mixed up with deformed ocean sediments. Also
andesitic lava deposits indicate the evidence of subduction prior to collision, as do the presence
of large batholiths which will decrease in age away from the suture line.
On the next page are two diagrams illustrating two general points about continental collision.
The first one (top of next page) shows a converging continental margin with one plate
subducting under another with thick deposits of sediment deposited on the ocean floor.
236
The second diagram (below) shows why orogenic belts are usually asymmetrical because one
continent must override another, so the direction of thrusting will typically be dictated by the
direction of subduction. Within the orogenic belt, deformation will be more intense and slightly
older in the inner zone less intense and slightly younger in the outer zone. The margins of both
continents are the first to be involved in deformation. At successive stages, the deformation
―front‖ moves outwards into the continental block. Sequentially younger thrusts develop in the
outer parts of the orogenic belt, carrying deformed rocks over the foreland. The belt as a
whole shortens and thickens.
The diagram on the next page shows a simplified geological cross-section of the Himalayan
collision zone. The Himalayas and Tibetan Plateau form a vast mountainous region (orogenic
belt) north of the Indian subcontinent. After the development of the theory of plate tectonics
in the 1960's it was recognised that the Himalayan mountains formed from the collision and
continued convergence between the Indian subcontinent and the rest of central Asia. The
sheer size of these mountains, where uplift rates greatly exceed those of erosion, and also the
abundant landslips and widespread seismicity, indicate that this plate convergence is still
operating today.
237
About
238
130
million years ago (Mya) the Indian Plate rifted away from the super-continent known as
Gondwanaland (Africa, Australia & Antarctica) and drifted north, due to sea-floor spreading,
toward Eurasia. Between the converging continents lay a wide sea called Tethys. By about 110
Mya the oceanic crust under the Tethys Sea started to be subducted under the Eurasian
continent to form an active continental margin similar to the present day Andes. The normal
subduction zone sequence of basalt-andesite-dacite magmas and lavas were produced in the
overlying continental crust. Between 40 to 50 Mya the two continents collided at what is now
termed the Tethyan Suture. This suture zone is marked by a chain of ophiolite fragments that
represent slices of the Tethys oceanic crust that were obducted and subsequently deformed as
collision continued.
Also in this area of the collision zone a vast body of plutonic rocks (The Trans-Himalaya
batholith) 50km wide, stretch for 2,500km along the Indian-Eurasian border. Their range in
age from 100 to 40 million years indicates that magmatism under the Eurasian continent
continued from the beginning of India's northward movement until about 10 million years after
collision. After collision India continued to move northwards with respect to the nearly
stationary Eurasian continent. So since the collision this convergence between the two
continents must have been taken up by huge deformation of the continental lithosphere as
India thrusted under the Eurasian continent for nearly 2000km.
5. How the Continents Formed
The continents have grown larger through geological time by the gradual accretion of material
derived from the upper mantle. Primitive crust was of an oceanic type and continents were small
or non-existent. Then through chemical differentiation of mantle material the continents slowly
grew. The earliest continental rocks came into existence at a few isolated island arcs. Here
denser and colder oceanic crust subducted under another oceanic plate. The descending cold
basaltic slab containing sea water, induced melting as it entered into the asthenosphere. Partial
melting of the basaltic crust caused intermediate magma (andesitic) to form. Volcanoes and
island arcs formed, as well as huge volumes of plutonic rocks. The generation of new crustal
material from the mantle at subduction zones is the single most important mechanism of net
crustal growth. Whether continental growth was more rapid at some stage in the geological
past, as is widely accepted, or involved mechanisms other than subduction-related magmatism,
are topics of debate amongst earth scientists!
These island arcs collided to form larger continental areas, while deforming and incorporating
the volcanic and sedimentary rocks that were deposited in the intervening oceans. Eventually
this process generated masses of continental crust of the size and thickness of modern
continents. The continent of North America, for example, has formed from the accretion of
colliding plates and island arcs over billions of years. These processes of accretion and
recycling sedimentary material involve no actual production of new crustal material from the
mantle reservoir, and thus no net growth of continental crust on a global scale.
239
GEOLOGICAL EVOLUTION OF BRITAIN
Contents:
1.
2.
3.
Plate Tectonic movements have resulted in the northward drift of the British area
through Phanerozoic time.
Changes in the latitude of the British area through geological time are interpreted
from evidence of former climates and from inferred palaeomagnetic pole positions.
The surface distribution of rocks in Britain has been determined largely by tectonic
activity related to former lithospheric plate movements.
Geological time can be divided into a number of Eons, Eras and Periods, with further
subdivisions into epochs. These are arranged chronologically, with the oldest at the bottom,
younging upwards to form the stratigraphic column (see geological timescale on back page of
this handbook). The stratigraphic column can be looked at in two ways. The first deals with the
order of rock units. This order has been established using the principle of superposition and
the principle of faunal succession, and produces the rock stratigraphic column, based simply on
the relative ages for rock successions. The second aspect of the stratigraphic column relates
to the geochronological dating of rocks using a variety of radiogenic isotopes. This forms the
time stratigraphic column and allows geologists to apply absolute ages to rock succession.
1. Plate Tectonic movements have resulted in the northward drift of the British area
through Phanerozoic time.
The following diagrams show how the Earth‘s continents have drifted across the globe over the
last 550 million years (Ma). Over the last ~550 Ma, the British Isles has formed a small portion
of series of different continental masses (being in fact part of two separate continents in the
Cambrian), and has slowly drifted northwards to its present position. Until the end of Silurian
and beginning of the Devonian, the northern and southern halves of what is now the British
Isles were on different continents, separated by an ocean – the Iapetus. The existence of this
ocean, along with the collision between these two continents is recorded in the Caledonian
Orogenic Belt, which contains rocks formed within and on the flanks of the now vanished
Iapetus. By ~375 Ma, this ocean had closed with the resultant continental collision produced a
series of major tectonic structures.
At a later date, the Variscan Orogenic Belt (which is found in the southern British Isles)
formed as a result of another period of continental collision, when the Rheic Ocean closed
between Laurentia and Gondwana. This tectonic activity led to the unification of all the globe‘s
main continental landmasses into one supercontinent called Pangaea.
The break-up of Pangaea resulted in the formation of new oceans (including the Atlantic), as
well as the formation of another extensive orogenic belt when Africa collided with Europe to
form the Alps, and India collided with Asia to form the Himalayas. These last two examples
illustrate that orogenic episodes in one region can occur at the same time as ocean spreading in
another region.
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Late Proterozoic ~ 550 Ma
The northern British Isles is located at
the passive margin of Laurentia, while
the southern British Isles is situated
behind the subducting margin of
Avalonia, a micro-continent on the edge
of Gondwana. Both Laurentia and
Avalonia are south of ~40° S and are
separated from each other by a
spreading ocean (which becomes the
Iapetus).
Early Ordovician ~ 490 Ma
The southern British Isles is still
located at the margin of Avalonia,
which has drifted southwards to ~60°
S. In contrast, Laurentia, carrying the
northern British Isles, has started to
drift northwards, residing at ~20° S,
separated from Gondwana by the
Iapetus Ocean (which is now beginning
to close).
Late Ordovician-Early Silurian ~ 450-440 Ma
The
Iapetus
Ocean
has
been
progressively closing, bringing the
micro-continent of Avalonia (including
the southern British Isles, ~30° S),
closer to Laurentia (including the
northern British Isles, ~20° S). At the
northern margin of the ocean,
subduction
is
occurring
below
Laurentia, whereas the southern
margin with Avalonia is passive. To the
south of Avalonia, the Rheic Ocean is
actively spreading.
241
Mid-Devonian ~375 Ma
―Zipper-like‖ continental collision has been
occurring between Laurentia and Avalonia,
uniting the British Isles along the Iapetus
suture zone. This collision is known as the
Caledonian Orogeny. At this time, the
British Isles are at ~20° - 25° S, located
within the southern desert latitudes.
Carboniferous ~302 Ma
As the Rheic Ocean closes between
Laurentia, Eurasia and Gondwana, the
Variscan Orogeny starts to affect the
southern British Isles. During the
Carboniferous, continental drift has
carried the British Isles northwards
across the equator, into subtropical
latitudes.
Triassic ~237 Ma
All of the landmasses have united to form
the supercontinent Pangaea. To the east,
Tethys is actively spreading, while the
British Isles continues to drift northwards
to ~20° - 30° N, equivalent to the modern
day Sahara latitudes.
242
Jurassic ~195 Ma
Break-up of Pangaea results in
Gondwana and Laurasia separating, as
the southern Atlantic Ocean starts
to rift open. The British Isles
continues to drift northwards to ~35°
- 40° N into more temperate
conditions,
with
lithospheric
extension
and
passive
rifting
occurring to the east (forming the
North Sea) and west (where later the
North Atlantic
will open).
Late Cretaceous – Early Tertiary ~65 Ma
Passive rifting has given way to active
rifting to the west of the British
Isles, allowing the northern Atlantic
Ocean to continue opening in a zipperlike fashion northwards. Active seafloor
spreading
is
occurring
throughout the Atlantic, Indian and
Pacific Oceans, whilst Tethys closes,
resulting in the eventual collision of
Africa, India and Eurasia
Summary:
Britain‘s northwards drift
through the Phanerozoic.
Latitude
55-60ºN
30ºN
0º
30ºS
60ºS
Date
0 Mya
200 Mya
300 Mya
400 Mya
600 Mya
243
2. Changes in the latitude of the British area through geological time are interpreted
from evidence of former climates and from inferred palaeomagnetic pole positions.
An insight into the position of the British area throughout geological time can be gained by
looking at five important lines of geological evidence.
First, there is evidence for the ancient climate as indicated by rock types and fossils. For
example desert sandstones, glacial boulder clays and fossilised coral reefs all tell about past
climates and probable past latitudes (see notes on sedimentary environments). This is done in
comparison with modern climates, assuming that conditions have not changed a great deal in the
last 500 million years. The geological sequence in Britain shows evidence of rocks produced in a
variety of climates from cold temperate conditions to sub-tropical reef environments, Saharalike deserts and tropical rainforests. Instead of proposing widespread global climatic changes in
the past geologists now simply say that Britain has drifted northwards during the last 500
million years from a position in the southern hemisphere, with southern Britain near the
Antarctic Circle, across the equator and into the northern hemisphere reaching its present
latitude around 50 million years ago. Since then movement following the opening of the North
Atlantic has been mainly eastwards. Britain probably spent 300 million years in tropical
latitudes, which explains the coral reefs, coal forests and Sahara-type deserts. However, this
assumes that the Earth has had the same climatic conditions and temperature differences
between the poles and the equator, a concept known as the principle of uniformitarianism (―the
present is the key to the past‖. In other words the processes and climate zones in the past are
still the same as present processes and climate zones.
Secondly, fossils can tell geologists, through what Darwin called faunal and floral provinces,
about ancient land bridges and land separation. Areas of land may be isolated from others by
wide oceans or impassable deserts, which can prevent animal and plant groups from mixing and
interbreeding, so that isolated communities arise. For example, trilobites from the Cambrian of
north-west Scotland show differences from those found in England and Wales, which suggests
a faunal province created by a wide ocean which restricted the intermingling of shallow water
animals. It had long been noticed that the trilobites in north-western Ireland and Scotland
showed more resemblance to those of North America, Greenland and Newfoundland than to
those found in England, Wales, Massachusetts, New Brunswick and south-east Ireland. These
trilobite faunal provinces are evidence for a wide ocean separating two continents. In the Lower
Palaeozoic, the present Atlantic Ocean had not formed, so North America and Greenland were
joined to Scotland and Northern Ireland. But these regions were widely separated by a huge
ocean from those places we now call England, Wales, Massachusetts and New Brunswick. This
long-vanished ocean has been called the Iapetus Ocean.
On either side of the Iapetus, separate trilobite faunas developed which were at their most
divergent in the Ordovician. The northern shores supported the so-called ―Pacific Province‖
trilobites, which included Olenellus, Bathynotus, Nevadia, Bonnia, Protypus and others. The
southern shores were characterised by the ―Atlantic Province‖, which included Callavia, Holmia
and Strenuella. Other fossil groups, such as brachiopods, also reflect this variation.
244
As the ocean began to close the faunas on its opposing shores came closer together. Their
migration of larval forms became possible across the narrowing sea. This narrowing process
continued until the two faunas were identical, and it led eventually to the complete closure of
the seaway in the Devonian.
These differences in trilobite populations prove that in the early Ordovician the Iapetus Ocean
was wide enough to have one side on the equator (Laurentian continent) and the other side
(Avalonia/Gondwana continent) on the Antarctic Circle (65ºS). This gives a maximum width of
about 7,000 km. Trilobites were shallow sea organisms that were influenced by climate. Some
liked it hot, others liked it cold, whilst some preferred lime and others muddy conditions.
Trilobites found in Wales and the Lake District such as Calymenella, Ogyginus and Neseuretus
lived in cold water environments. Whilst trilobites, such as Bathyuriellus, which are found in
Scotland, lived in warm, tropical seas. After further research a third group of trilobites have
been found in Scandinavia which are totally different from the Scottish or Welsh species.
These trilobites lived in temperate latitudes (backed up by palaeomagnetic data) and indicate
the presence of a third continent midway between Laurentia and Gondwana.
Thirdly, the presence of acidic and intermediate volcanic lavas and ashes, rhyolites and
andesites in a marine sedimentary sequence suggests that the rocks were formed near a
subduction zone in a closing ocean. Conversely the absence of andesitic volcanic rocks suggests
a widening ocean.
Fourthly, the presence of basaltic pillow lavas erupted on the ocean floor indicates old oceanic
crust. Being dense, these are usually subducted at plate boundaries but sometimes appear
within continents and are called ophiolites. When found in ancient rocks, these indicate that
sea-floor spreading has happened in that area.
Lastly, palaeomagnetism (ancient magnetism) allows geologists to establish where rocks were
when they took up their initial magnetic field in relation to the magnetic poles. This tells
geologists the ancient latitude of formation of many rocks. The Earth has magnetic field
derived from the convection of the electrically conducting liquid outer core. At the Earth‘s
surface, the magnetic field is similar to that produced by a short bar magnet with poles deep
inside the Earth and its axis roughly aligned parallel to the Earth‘s rotation axis. The magnetic
field can be visualised by means of lines of force, which are parallel to the Earth‘s surface at
the equator, dip steeply at the high latitudes and point straight down at the poles. In olden
days, navigators could tell roughly at what latitude they were by using this feature of magnetic
inclination and measuring the angle at which a compass needle dips.
Some rocks, particularly basalts, contain natural magnets in the form of the iron oxide mineral,
magnetite. Magnetite crystals become magnetised in the direction of the prevailing magnetic
field at the time of their formation. For example, young basalt samples from latitudes near the
magnetic poles show steep magnetic inclinations within magnetite crystals, while similar samples
from equatorial region show shallow inclinations. Magnetite-bearing rocks can therefore be
used to indicate the latitude that the site of the sample occupied at the time it formed.
Geologists term these rocks as good Palaeolatitude indicators. As the Earth‘s magnetic field is
axially symmetric, magnets are no use as indicators of longitude. Basalt contains a significant
245
proportion of magnetite crystals and is thus an excellent recorder of Palaeolatitude. When
liquid basalt erupts and cools below Curie point (~600°C for magnetite), magnetite crystals
become aligned according to the strength of the magnetic field and the latitude of the
eruption. As the lava cools, the mini-magnets become frozen into the rock, aligned in the
prevailing magnetic field. Samples of basalt with a known present-day orientation can be
collected and examined in the laboratory to give information about the magnetic field
orientation at the time the sample cooled. If a large expanse of basalt can be sampled,
palaeomagnetic readings from well-separated localities can be used both to investigate the
latitude at the time of the eruption and how much rotation has been experienced.
Magnetite-bearing rocks therefore are able to provide a great deal of data about the previous
latitudes of Britain. Another consequence of measuring palaeolatitudes from magnetite-bearing
samples is that such samples can be used to locate the fossil magnetic poles (palaeopoles), as
long as the palaeolatitudes of two or more samples of exactly the same age is known. If basaltic
rocks of various ages are collected from one continent and their remanent magnetism
measured, then the location of the magnetic poles at the time each rock was formed can be
calculated. If these locations are linked in chronological order and located on a map, they form
an apparent polar wandering curve. It is called this because it seems as though the poles have
shifted over time, when in actual fact it is the continents which have moved. If the position of
the magnetic poles is assumed to have been fixed over time then this palaeomagnetic data will
246
record the movement of the continent instead of the poles. However, apparent polar wandering
curves are useful in two ways. First, apparent polar wandering curves plotted for many places on
the same continent can be used as verification of the palaeolatitude data. The second use of
apparent polar wandering curves is to identify whether and when continents were joined and
when they drifted apart.
For example, when samples taken from igneous rocks in Europe and North America of the same
age but older than 100 Ma, they show different alignments of magnetite crystals depending
from which continent they were taken from. It is clear that the polar wandering curves for
Europe and North America do not match (see diagrams below). However, when North America is
rotated toward Europe, an almost perfect match in the two curves can be generated between
the two continents.
Some sedimentary facies, fossils assemblages and depositional environments indicate a
particular climatic zone e.g. desert environments, glacial environments and tropical shallow
seas.
Quaternary
Ground moraine, boulder clay & till deposited by ice. Sands & gravels from meltwater streams.
Scree slopes & solifluction deposits suggest periglacial conditions.
Deposits of glacial origin take characteristic forms. The most well-known is called boulder clay,
or more correctly, till. Boulder clay is a rock flour produced by the grinding action of ice sheets
and glaciers as they pass over the solid bed rock. The ice also picks up boulders and rock
fragment which it uses at its base as grinding tools, and these boulders themselves can
eventually be reduced to flour. When the ice retreats it leaves behind a mixture of clay and
boulders, hence the term boulder clay. The boulders can vary in size but are usually well
rounded, smoothed and occasionally marked with scratches. Many boulders have travelled
247
considerable distances from their original source where they were first picked up by the ice,
and are called glacial erratics. In East Anglia there are erratics from Norway, and in
Shropshire many different granite erratics have come from Scotland and the Lake District. Till
is the term used more often nowadays than boulder clay. Frequently it is sandy, or a mixture of
sand and clay. Clay tills (true boulder clays) are sticky, blue clays (orange when weathered)
which can be a gardener‘s nightmare in parts of Lancashire!
Glacial moraines are formed on the top of a glacier as material falls on to its surface from the
valley sides, and also at its snout. When the ice melts they are left behind on the ground
surface over which the glacier or ice sheet has been moving. Moraines are usually irregular
dumps of poorly sorted, angular and coarse-grained material. Examples of morainic gravels can
be found in the mountain areas of the Scottish Highlands and the Lake District. Drumlins are
well-known landforms left behind by moving ice sheets. Glacial moraine deposited by the ice is
further moulded into smooth elongated hills with the blunter end towards the ice. They have
the shape of half an egg, typical dimensions being 400m by 100m, and often occur in large
numbers in drumlin fields, and produce a type of hummocky land surface called a basket of
eggs, well known in Ribblesdale and the Eden Valley in northern England.
As well as glacial tills and moraines which are all deposited directly from the ice, melting ice
sheets and glaciers produce large amounts of fluvio-glacial sands and gravels which are washed
out of the base and the edge of the ice by meltwater streams. Kames are masses of gravel
dumped at the edge of an ice sheet and eskers are long, narrow sinuous ridges of sand and
gravel formed in tunnels within the ice by melt waters flowing between the ice and the ground
surface. The finest examples of these are in north-east Scotland near Inverness where the
highest esker in Britain, the Kildrummie esker, is found at Torvean (70m high). Fluvio-glacial
deposits are generally sorted, rounded and stratified (layered) deposits.
Areas adjacent to or above ice sheets during the Ice Age were affected by what is called a
periglacial climate (peri means around the edge of, as in perimeter). Here, extreme cold at
night can be followed by relatively high temperatures in the day. This causes freeze-thaw
action when water which has been trapped in joints of crags expands at night, causing the rocks
to become shattered as the process is repeated day after day. Large scree fields form around
the frost shattered crags. The screes above Wastwater in Wasdale in the Lake District are a
classic example. In areas where bare rock does not occur at the surface, freeze-thaw action
causes the whole hill surface to become broken up to a depth of a few metres, into a coarse
angular deposit called head. Large areas of ground are affected by very slow gravity
movements of these fractured surface deposits, a process called solifluction.
Tertiary
Plant remains (lotus flowers) in Skye suggest climate was often humid & sub-tropical. Between
eruptions a red laterite soil formed indicating rapid weathering in a tropical environment. In
southern England fossil evidence of rich vegetation such palm trees, as well as crocodiles.
248
Cretaceous
Chalk widespread. Very pure CaC03 so no terrigenous input so far from land in a clear, warm,
tropical sea.
Jurassic
Oolitic limestone in the Cotswolds indicates clear, warm, shallow, marine environment. Some
land fossils – reptiles & dinosaurs, and desiccation of shallow seas in southern England backs
this up.
Permo-Triassic
Desert sandstone, conglomerates, evaporites & desiccation cracks in Shropshire, Midlands &
Liverpool
Deserts conjure up an image of a vast area of sand dunes, but in fact this is only one of several
sub-environments found in deserts. Deserts are defined as areas where the average rate of
evaporation exceeds the average rate of precipitation. This lack of water is the reason why
they contain very little plant and animal life. Lack of vegetation and water also means deserts
are dominated by physical weathering and that when precipitation does come erosion is rapid,
thus sediments in deserts can be easily moved around. Deserts occur in both hot and cold
climatic areas of the world. Hot deserts cover a larger area of the Earth and it is here that the
constant heating during the day and cooling at night set up stresses within the rocks which lead
to increased physical weathering.
The dominant agents of transport and deposition – wind and water – can be used to divide
desert features and sedimentary deposits into two categories. The first category is those
features and sedimentary deposits formed mainly by wind; these include aeolian sand dunes,
extensive area of which may be called sand seas, together with large areas of bare rock and
rock debris. The second category is the features and sediments formed by the transient
passage of water; these include alluvial fans, river valleys and temporary lakes (playa lakes).
Aeolian processes are in many ways similar to water transportational processes. However, the
difference in density and viscosity between air and water leads to some differences in
sediment behaviour. Saltation of fine to medium-grained sands and suspension of silts are the
dominant processes during aeolian (wind) transport, whereas rolling and sliding, as well as
saltation and suspension transport are important processes during water transport. As sand
grains saltate to greater heights in the air than water, and strike the grains on the surface
beneath them at higher speeds, they cause coarser grains to move forwards by surface creep.
This process does not occur in water. As might be expected, these differences in mode of
transport lead to differences in the nature of the aeolian bedforms, and how they develop. The
main controls on aeolian bedform morphology are wind speed, wind direction, and sediment
supply.
Despite various experiments with sands in wind tunnels, the process of aeolian ripple formation
is not well understood. One suggestion is as follows. Sand grains begin to saltate as soon as the
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wind speed at a particular point on the bed is sufficient to start them moving. As aeolian sands
are well sorted, the grains will all be fairly similar in size and so they will tend to saltate around
the same distance, descending to hit the surface again in more or less the same place. On
impact, they set in motion the grains that were too coarse-grained to saltate. So the coarser
grains move by surface creep, and the impact makes the finer grains saltate again. As this
process is repeated, a series of regularly spaced impact sites develops on the bed. Sand is
pushed forward at these sites by surface creep and so impact ripples begin to form. These are
rarely preserved in the sedimentary record.
Finer-grained sand migrates rapidly over the ripples by saltation and some is deposited and
trapped in the ripple troughs, but coarser-grained sand moves more slowly as it is pushed up
the up-wind side of the ripples by surface creep. This process results in coarser-grained sand
becoming concentrated at the ripple crests. This contrasts with current-formed ripples where
coarser-grained material is located in the ripple troughs. Impact ripples are slightly
asymmetrical, like current-ripples, but they are flatter and their crests are long and relatively
straight with occasional bifurcations.
Unlike subaqueous dunes, aeolian dunes do not grow from ripples in response to increasing wind
speed. As wind speeds increase, ripples are washed out and eventually replaced by a flat bed
again. Aeolian dunes take much longer to develop than ripples and, provided there is enough
sand available, may grow to spectacularly greater heights than the subaqueous varieties,
because they are not constrained by flow depth. Furthermore, actively migrating sand dunes
may have migrating impact ripples superimposed on them, and smaller dunes may grow on the
backs of larger ones. For aeolian dunes to grow it seems that there has to be some obstacle in
the path of the sand to begin with, such as an isolated clump of vegetation or a rock. As the
airstream flows over the obstacle, a wind shadow is created on the down-wind side, containing
localised eddies where sand is deposited to build up drifts. When sufficient sand has
accumulated the drift becomes independent of the original obstacle and starts migrating as a
dune.
Saltating sand grains bounce their way up the shallow slope of a dune to the crest where they
periodically avalanche down the steep lee slope by a process known as grain flow. On this steep
lee slope of the dune shallow gullies occur, these are the sites of small avalanches where grain
flow is takes. The sand on this steep slope often exceeds its stable angle of rest, and so begins
to move under the influence of gravity. Once in motion, the flow is sustained by grains colliding
and ricocheting, and rolling over each other. During this process the larger grains are
preferentially moved upwards to the top of the flow and the finer grains move downwards.
Once the flow has stopped the grains, even though of very similar size, have been sorted, giving
rise to cross-bedding. Over time more and more sand is transported up the shallow windward
slope and deposited on the steep lee slope, where it becomes unstable and avalanches forming
another set of cross-beds on top of the previous set. Thus, the dune gradually migrates in the
direction of wind flow.
Aeolian dunes in the sedimentary record look very similar to those formed in water. There are a
number of pieces of evidence to distinguish between the two. To begin with, the aeolian crossbeds are much thicker, commonly in the order of 5-10m, whereas subaqueous cross-beds are
250
normally less than 2m thick. Secondly, although the cross-beds in aeolian dunes at the base of
each set make only a low angle with the horizontal, they curve upwards more steeply than
subaqueous cross-beds. This is because the steep slope of an aeolian dune may be stable up to
an angle of rest of about 35°, whereas the angle of rest of subaqueous dune slopes is usually
much less. Thirdly, in aeolian dunes cross-bedding is much better developed due to the fact
that the process of grain flow is much more efficient in air than in water, so the sorting
process is more pronounced. Other pieces of evidence would be the frosted surfaces to the
aeolian sand grains, due to the lake of water to cushion the contacts between grains. The
aeolian sandstones would also be texturally very mature, although sands which have been
subjected to long periods of wave transport would also be texturally mature. A lack of marine
fossils would also be a useful clue.
In order to preserve the dune sands, so that eventually they will become sandstones in the
geological record, there has to be net accumulation of sand. So the sands have to be protected
from the erosive and transportational power of winds. This can be achieved in several ways. (i)
Through deposition at or near the water table (wet sands simply do not blow away), combined
with subsidence, allowing more and more sand to accumulate. (ii) The growth of vegetation or
the formation of surface cements, protecting the sands and preventing them from blowing
away. The red iron oxide (haematite) coating around the grains of desert sandstones could be a
form of early cement that may well have helped to stabilize these desert sandstones so they
could be preserved. (iii) If the supply of sediment is exceptionally high, then the sands may
accumulate because there is insufficient wind to remove them.
So, in summary, wind processes are important in shaping, transporting and depositing sand
grains, together with moving all loose sediment away from an area and leaving behind bare rock
and clasts that are too large to be moved. The wind is also important in a desert in helping any
water that may be present to evaporate. Water causes some of the most spectacular and
dramatic features of desert erosion because most of the sediments are not bound together by
vegetation, fine-grained sediment or cement. Rainfall is sporadic, but when it occurs over a
short seasonal period or every few years it may well be torrential. In desert regions within or
near to mountain ranges, a rainstorm may be catastrophic. The sudden rainfall carries large
volumes of accumulated sediment down from the mountains and into the desert area and may
erode a steep-sided, narrow valley called a wadi (Arabic for watercourse). As the flood waters
subside by infiltration and evaporation, sediment is deposited. It is usually poorly sorted and
may well show only crudely defined bedding. For much of the time, most wadis are dry valleys
that are only periodically occupied by water. There are two other morphological features
formed by the presence of water in a desert; alluvial fans and playa lakes.
Alluvial fans are localised deposits that form where a river or stream loaded with sediment
emerges from a confined mountain valley (such as a wadi) onto a flat lowland plain. Alluvial fans
have a basic semi-conical shape. The apex of the cone points up the valley it has come from. As
the water flows out from the steep valley and onto the flat plain, the sudden decrease in
gradient is accompanied by extensive deposition from sediment gravity flows. Although alluvial
fans are most common in semi-arid and glacial climatic regions where vegetation is sparse and
run-off is seasonal, they can also form in humid tropical regimes. Alluvial fan deposits are
generally poorly sorted, matrix-supported and both texturally and compositionally immature,
251
but they may show varying degrees of layering or graded bedding depending on how much water
remains in the flow. The actual degree of grain roundness is a function of the sediment source,
how far the sediment gravity flow travelled and whether it was fairly viscous so the grains
were cushioned. The fan is traversed by a network of rapidly shifting braided streams and so
the sediments in the channels comprise gravels, sandy gravels and sands. The sands may be
arkosic if the nearby mountains contain feldspar-bearing rocks and chemical weathering and
grain transport are limited. A feature common to all alluvial fans is that the sediments become
finer-grained with increasing distance from the apex of the fan. Alluvial fans are a prominent
feature of most braided river systems in mountainous areas and where mountains meet flat
desert areas.
Temporary lakes, often termed playa lakes, can form in any depressions in the desert surface
as it rains and are common at the base of alluvial fans. The playa lakes become infilled with
fine-grained sediments that have been transported in suspension either by wind or water. If
sediment has been carried by water down alluvial fans, the finest-grained sediments will be
carried the furthest and will be deposited at the distal (furthest from the apex) part of the
fan in the flat area where the playa lake forms. The resulting sedimentary rocks are typically
laminated mudstones and siltstones. Desiccation cracks are very common and form as the finegrained sediments dry out. The water may support organisms, such that both trace and body
fossils may be found. If the water does not drain away, but is allowed to evaporate slowly,
evaporite minerals such as gypsum and halite will be deposited and a salt pan is formed.
Carboniferous
Carboniferous rocks and fossils show that Britain was close to the equator. In the Lower
Carboniferous a rise in sea level meant that most of Britain was underwater. This warm tropical
sea favoured limestone deposition but the pattern of sedimentation also depended on the shape
of the drowned landscape beneath. Some high areas remained as islands and did not receive
sediment. The deeper waters of flooded valleys and basins tended to accumulate beds of muddy
limestone and shale. Only in the really shallow and warmest regions of the sea was pure
grey/white limestone formed. These can be seen in the Northern Pennines (Ingleton &
Sedgwick Trail) and Peak District today. The presence of colonial coral and algal reefs shows
that tropical conditions existed during this time.
During the Upper Carboniferous deep-water sediments such as shales continued to be
deposited across southern England but, over the rest of the country, conditions were changing.
A series of uplifts linked to the Variscan orogenesis reduced the depth of water lying across
central Britain. As a result many of these shallow areas became infilled with sediment
deposited by deltas. These advancing deltas produced an interbedded sequence of shales and
coarse arkose sandstones known as Millstone Grit.
Mid-way through the Upper Carboniferous, the deltas had become so extensive that only the
Rheic Ocean (across southern England) remained a truly marine area. Everywhere else apart
from isolated areas and northern Scotland had become a vast swampy deltaic area of river
channels, lakes and coastal mudflats. It was this environment which provided Britain‘s most
valuable deposit; the Coal Measures. Newly evolved species of land plants grew rapidly in the
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hot wet equatorial climate, and the swampy conditions allowed vegetation to sink and
accumulate rather than rot away. The periodic ―drowning‖ of deltas (due to local subsidence &
global sea level rises as a result of glacial and interglacial episodes in Gondwana) allowed
sediment to bury the layers of plant material preserving it to eventually form coal.
3. The surface distribution of rocks in Britain has been determined largely by tectonic
activity related to former lithospheric plate movements.
55ºN
45ºN
40ºN
d. Alpine Orogenesis –
Cenozoic ~ 40 Mya
c. Opening of North Atlantic –
Cenozoic ~ 65 Mya
30ºN
20ºN
10ºN
0º
b. Variscan Orogenesis –
Upper Palaeozoic ~ 300
Mya
a. Caledonian Orogenesis –
Lower Palaeozoic ~ 400 Mya
30ºS
60ºS
253
Rocks from all the major subdivisions of geological time (i.e. Precambrian, Palaeozoic, Mesozoic
and Cenozoic) occur in Britain. A number of orogenic events have affected the British Isles;
these include the Caledonian, Variscan and Alpine Orogenic Belts. The study of these orogenic
belts enables a reconstruction of the plate tectonic regimes in which they developed. The
Tertiary Igneous Province of NW Britain provides evidence of the early history of the opening
of the North Atlantic Ocean, with associated basaltic volcanicity.
CALEDONIAN OROGENESIS – LOWER PALAEOZOIC ~ 400 MYA
Northwest Highlands
Europe's oldest rocks, 3 billion years old, are found in a narrow strip of the Northwest
Highlands. These ancient rocks are known as the Lewisian, after the Isle of Lewis in the Outer
Hebrides. These Lewisian rocks occur to the west of the Moine Thrust zone, which represents
the western margin of the Caledonian mountain belt. Therefore the old rocks of the Northwest
254
Highlands are referred to as belonging to the foreland area of the collision zone. This is a
stable segment of the crust that was largely unaffected by the later events of the Caledonian
orogeny. The Lewisian rocks are Precambrian gneisses (3300-1750 mya), banded in white, pink
and black, which have been folded and intensely metamorphosed deep within the crust. These
rocks are known as basement rocks and represent a relatively small fragment of a much more
extensive region of Precambrian continental crust in the North Atlantic, usually referred to as
Laurentia, which includes the Canadian Shield and Greenland.
Northern Highlands
The Northern Highlands lie between the Moine Thrust and the Great Glen Fault. The
metamorphic rocks of the Northern Highlands are known as the Moine rocks. They are silvery
grey, well banded coarse-grained schists, with abundant flakes of muscovite and biotite mica
which glisten on the flat surfaces. The growth of the platy mica crystals (recrystallisation) in
parallel layers forms foliation and is due to heat and pressure from regional metamorphism. The
Moine schists were originally deposited as a thick sequence of sediments deposited in a shallow
sea 1000-870 million years ago. This sea lay on the eastern margin of Laurentia. In total, about
12km thickness of sand, silt and mud was laid down and subsequently went through two
metamorphic events:

at 870-800 million years ago, when nappe folds formed, during the Grenville orogeny,
which was one of a series of continental-collision events that gave rise to the large
supercontinent of Rondinia;

and again at 450 million years ago in the main phase of the Caledonian orogeny.
The Moine Thrust zone varies in width from 0 to 12 km wide and runs down the northwest coast
from Loch Eriboll to the Isle of Skye, a distance of 200 km. Thrusting took place at about 400
mya, in a final episode of the Caledonian orogeny. The Moine schists of the Northern Highlands
were carried many kilometres to the northwest, possibly up to 150 km, over the top of old
basement rocks. During these movements the rocks on either side of the thrust were strongly
sheared and streaked out. This makes a new type of rock, formed by deformation, called
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mylonite. In places these strongly sheared rocks are over 100m thick. All the mineral grains in
the mylonite are streaked out and can become aligned, needle-like, in the direction of shearing.
By measuring the alignment
direction of grains geologists
can
establish
in
which
direction the Moine thrust
moved. The answer is towards
WNW - and this direction is
remarkably consistent all the
way down the Moine Thrust
zone. The thrust zone dips at
about 8-12º to the east as a
result of more recent tilting.
The photograph above shows
the Moine Thrust and the bed
of
mylonite
formed
by
dynamic metamorphism at the
Stac of Glencoul in NW
Scotland.
In the Assynt district of NW Scotland, it is possible to see several of these Caledonian
thrusts, with the Sole Thrust being the lowest (at Loch Assynt), followed upwards by the Ben
More Thrust, Glencoul Thrust and Moine Thrust. Between these thrusts, the underlying rocks
have been folded into anticlines and broken by many small-scale reverse faults and minor
thrusts, to create incredibly complex packets of deformed rocks. The photograph on the next
page shows the view across Loch Glencoul to the Aird da Loch peninsula. It is one of the most
photographed in World geology. For in that view there is the Glencoul Thrust, one of the major
structures of Assynt and one of the first thrusts in the world to be recognised as such.
Thrusts are the building blocks of mountain belts. In essence they are faults that stack up
rocks. They can occur on all scales, stacking up an individual bed time and again, or on a crustal
scale, repeating the crust.
Thrusts are structures that
repeat rock units. If strata
were originally in their normal
sequence (oldest at the
bottom, youngest on top) then
thrusts carry older rocks on
top of younger. In this way
they act to thicken units up which
is
why
they're
important
for
mountain
building - thrusts thicken the
crust. This might sound
straightforward but in detail
256
thrusts can produce very complex structures. First, they are rarely planar (straight) faults so
that rocks that have been moved on thrusts are generally folded. Second, thrusts very rarely
happen individually - they come in clusters. These are called imbricate stacks.
The structures in the Moine relate to crustal shortening. The lowest ductile thrust of note is
the Moine Thrust itself.
Collectively
the
Moine
structures give tell us how
the middle parts of the
continental crust deform
during mountain building.
The timing of deformation
in the Moine is highly
controversial, relying on
sparse radiometric dating.
The
ductile
thrusting
broadly
represents
Caledonian
deformation.
New information published in 2003 suggests that ductile thrusting in Sutherland is Scandian in
age 430 million years ago - based on the ages of various granitic bodies that were emplaced
during shearing and isotopic dating of recrystallised micas within the mylonite belt.
Grampian Highlands
The Grampian Highlands lie between the Great Glen Fault and the Highland Boundary Fault.
This vast region is dominated by huge thicknesses of repeatedly folded and metamorphosed
Dalradian slate, phyllite, schist, metaquartzite, marble and metavolcanic rocks, all formed
during the Caledonian orogeny. The Dalradian rocks represent sediments deposited on a
continental shelf that lay on the edge of a basin that opened to the east, probably as a result
of rifting of the Laurentian margin related to an early stage in the opening of the Iapetus
Ocean. Radioactive dating has shown these original Dalradian sediments to be late Precambrian
to early Ordovician (750-480 mya).
The Ballachulish Slate is an important member of the group. These black slates (regionally
metamorphosed muds) are up to 400m thick and indicate a deep water environment of
deposition. As the Dalradian basin became deeper over time fine-grained muds and sands were
deposited from rapidly flowing turbidity currents. These dense mixtures of sediment (mud, silt
& sand) and sea water, dislodged from the continental shelf by earthquakes, flowed rapidly
down slope to the ocean floor, where the suspended sediment settled out, forming a turbidite.
Regional metamorphism later converted these sediments into slates and greywackes, such as
the Easdale Slates.
During the Caledonian orogeny, the Dalradian rocks were folded 3 or 4 times in what are
referred to as episodes of deformation. The earliest major folds run NE to SW, and these
257
were later re-folded to produce complex overfolds or nappes, the most important being the
Tay Nappe. During the formation of the nappe folds, the Dalradian rocks were at the peak of
regional metamorphism, and was when most of the recrystallisation and growth of metamorphic
minerals took place. The type of metamorphic mineral varied depending on the metamorphic
grade experienced by the rocks (i.e. the pressure and temperature). Low grade metamorphism
would form chlorite, medium grade metamorphism would form biotite, then garnet and
staurolite, eventually leading to high grade metamorphism which formed kyanite and sillimanite.
Igneous rocks are abundant in the Dalradian of the Grampian Highlands. In the Buchan area of
northeast Scotland there is an important concentration of large flat-lying igneous bodies that
were intruded into Dalradian rocks. These intrusions include layered ultrabasic to basic bodies
of peridotite and gabbro, and were folded and disrupted during deformation, after they were
intruded. These igneous rocks are thought to represent the deep root zones of a volcanic arc
that formed on the northern margin of the Iapetus Ocean during the Ordovician period (490
mya).
As in other mountain chains around the world, the Caledonian belt is characterized by abundant
granite intrusions, which rose as a result of melting at the base of the continental crust, which
had been thickened because of folding and the piling up of thrust sheets during continental
collision. These post-collision granites form when the country rock is buried to a depth of over
20 to 25 km, raising the temperatures to 650-700ºC, causing the continental crust to partially
melt. Most of the granites, such as the Cairngorms, are about 405 million years old and are
258
relatively undeformed which suggests they represent the final stages in the closing of the
Iapetus Ocean as the continental crust thickened in the collision zone.
Midland Valley
The Midland Valley lies between the Highland Boundary Fault and the Southern Uplands Fault.
Just north of the Southern Uplands Fault lies the world-renowned Ballantrae Complex. This
consists of fragments of the ancient Iapetus Ocean floor - pillow lavas, gabbro, serpentinized
peridotite (a soft dark-green ultrabasic rock which has been altered by the effects of hot
water-rich fluids on olivine and pyroxene), and black shales - and it formed in the early
Ordovician period 500-475 mya. This fragment of oceanic crust was then folded, faulted and
pushed by thrust faulting and carried onto the continental margin when the Iapetus was closing
and Laurentia (Scotland) and Avalonia (England) were colliding together towards the end of the
Caledonian orogeny. Oceanic crust like this that is thrust over and interleaved with folded
continental crust in a mountain belt is known as an ophiolite, and is formed by the process of
obduction. Obduction of the Iapetus Oceanic crust began at about 480 mya, and by 470 mya
the ophiolite complex was finally welded onto the Midland Valley.
The Ballantrae area helps to piece together the evidence for the closure of the Iapetus Ocean
by subduction, followed by continental/continental collision in the final stages of the Caledonian
mountain-building episode. Oceanic crust was being destroyed at this time in a major subduction
zone that lay to the south, with a volcanic arc to the north, at the margin of a large continent,
Laurentia.
The oldest sedimentary rocks in the Midland Valley are Ordovician and Silurian sandstones,
mudstones and limestones laid down 460-430 mya in shallow tropical seas bordering the Iapetus
Ocean. Within these areas of rocks are a series of inliers (an area of older rock completely
surrounded by younger rocks) composed of andesitic and basaltic lavas, 420-410 mya. This
string of inliers probably represents a volcanic arc which formed along the edge of the
Laurentian continent when the Iapetus Ocean was closing as Avalonia approached from the
southeast. Volcanic rocks are very abundant in the Pentland Hills, near Edinburgh. These basalts
and andesites are 1800m thick in the Pentlands. In addition to lava flows, there are also beds of
tuff representing volcanic ash from explosive eruptions. Igneous activity associated with
subduction ceased in the early Devonian, 410 mya, as continental collision and suturing took
place.
By late Silurian to early Devonian time (420-400 mya), the closure of the Iapetus Ocean had
led to the formation of a high mountain chain. Also at this time, the Midland Valley began to
develop as a basin because of crustal tension that caused the Highland Boundary Fault and
Southern Uplands Fault to act as a pair of parallel normal faults, allowing the crust in between
the faults to sink and form a rift valley. Erosion of the high new mountains produced copious
amounts of sediment which was deposited in this basin. As sediment was brought into the
basins, faulting continued to allow the base to sink and accept more material. About 9km of Old
Red Sandstone sediments (conglomerates & sandstones) were deposited in the Midland Valley.
259
The decrease in grain size of these sediments over time indicates that the source area was
becoming less rugged, and the mountains were being worn away and fast-flowing torrents
carrying large boulders gave way to slower rivers that transported sand and small pebbles.
Southern Uplands
The Southern Uplands has rocks that were folded during the Caledonian Orogeny. Most of
these rocks were laid down on the ocean floor from 470 to 420 mya in the Ordovician and
Silurian periods. The region lies between the Southern Uplands Fault (NE to SW) and the
Iapetus Suture (a NE to SW structure that joins northern and southern Britain). Within this
75km-wide block, the rocks have a strike of NE to SW, parallel to which are many thrust and
reverse
faults, the
most important being
the Stinchar Valley
Fault, Orlock Bridge
Fault and Laurieston
Fault. These faults
divide the Southern
Uplands into many thin
slices
which
get
progressively
older
towards the NW.
Compared
to
other
areas in the Highlands
the rocks of the
Southern Uplands have
been
only
weakly
metamorphosed.
The
great majority of the
older rocks are coarse
greywackes
with
smaller
amounts
of
finer siltstone, mudstone, shale, volcanic ash and lavas. Metamorphism of the fine-grained rocks
(mudstone & shale) has produced slate. Greywacke, a grey, black, dark-green or deep-purplish
hard rock, is coarse-grained and poorly sorted. It contains angular fragments of quartz,
feldspar, ferromagnesian minerals and igneous and metamorphic rock fragments held together
in a dark mud or clay matrix. Low-grade metamorphism has recrystallised the cement to
produce a tough hard rock which often looks glassy and igneous in texture, but it is a
sedimentary rock laid down in deep water.
Fast-flowing currents transported sediment off the edge of the continental shelf down the
continental slope to the deep abyssal plain of the ocean floor. Mud, sand and water mixed
together into a dense slurry, and the currents were probably triggered off by earthquakes.
These flows, known as turbidity currents, gathered speed on the way down the slope, mixing
260
the components together. On reaching the flat ocean floor, the current lost energy and
suddenly slowed down. Coarse mixed sediment was deposited first to form thick sandstone and
greywacke layers, on top of which fine sand settled to form siltstone, followed gradually by the
much slower deposition of clay and mud to form shale and mudstone at the top of the sequence.
The pattern repeated itself many times over to produce a series of graded beds, from coarse
at the bottom to fine at the top. Sediments deposited from these currents of sand and mud
are known as turbidites. Modern turbidity currents have been observed in the Atlantic Ocean
and they are known to flow very fast, at speeds of over 60km/hour, with tremendous erosive
power so that they are able to excavate channels and grooves into the sea bed on their way
down slope. These erosional features at the base of the beds are known as sole marks and
include features such as flute casts and tool marks formed as stones were dragged along the
sea bed. These marks are very useful for showing the correct way-up of the sediments.
Fossils in the Southern Upland rocks tend to be restricted to the fine sediments, where dead
animal remains would have been preserved by slow accumulation of mud far below the oxygenrich upper waters. The most abundant fossils are the graptolites, a group of marine animals,
now extinct, that floated in the early oceans as colonies of tiny polyps attached by thread-like
structure. By way of contrast the greywackes are too coarse and were deposited in too violent
an environment to allow anything to be preserved, let alone fossilised.
The Southern Uplands greywackes are divided into 3 main zones, separated by faults that run
NE-SW, parallel to the strike of the beds. The northern belt has Ordovician beds (470-460
mya), the central belt contains early Silurian sediments (460-430 mya) and the southern belt
has mid-Silurian rocks (430-420 mya), that is, the rocks become progressively younger towards
the south east. Intense compression has produced a series of tight anticlines and synclines in
which the limbs of the folds are nearly parallel. The many NE-SW trending faults in the
Southern Uplands all have their downthrow side on the SE. Within each of the 30 or so fault
blocks, the oldest beds are on the SE side, but the blocks get progressively older towards the
NW.
When all the field evidence is considered together, the Southern Uplands area is interpreted
as having formed at the Laurentian margin of the Iapetus Ocean when the ocean floor was
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destroyed by subduction northwards beneath the plate carrying the continent. Sediments lying
above the subduction zone were folded, sliced by thrusts and reverse faults, and stacked up in
a pile against Laurentia. Finally, this pile of greywackes was thrust towards the south when
Scotland and England collided. It is possible that the Iapetus Ocean was over 5000km wide and,
if the present 75km wide belt of the Southern Uplands could be unfolded and the effects of
faulting removed, then the sediments would stretch out to at least 1500km. The term given to
this pile of sedimentary rocks, scraped off the ocean floor as Avalonia and Laurentia collided, is
an accretionary prism.
In this process the greywackes were folded and weakly metamorphosed as they became
physically attached to Laurentia in the Caledonian orogeny. During the collision, the sediments
were scraped off the ocean floor while oceanic crust was being destroyed by being forced down
the subduction zone. Sediment piles were transported along thrust faults and successive
thrusts pushed from underneath, causing the overlying material to be folded and pushed up into
a much steeper orientation. Eventually, the Southern Uplands ended up being thrust right over
continental crust as the ocean closed and was finally destroyed at the end of the Silurian, 418
mya.
Lake District
1. Ordovician Skiddaw Group rocks are the oldest in the Lake District, forming almost one
third of the mountain core, and outcropping in a wide belt in the northern and western fells.
There are also inliers of Skiddaw Group rocks such as around Haweswater and Ullswater and,
further east. These rocks are very difficult to interpret. They are mainly greywackes,
sandstones, siltstones and mudstones (some now metamorphosed to slates). They were
deposited by turbidity currents in fairly deep water on the continental slopes of the former
Iapetus Ocean. They are more than 3,000 m thick. They are monotonous and uniform grey in
colour. They are difficult to interpret as evidence of sedimentary structures are difficult to
find, and fossils are rare. Structurally they are also difficult to interpret, as they are very
complex. Besides having undergone many phases of folding, they are also highly cleaved (hence
the old name 'Skiddaw Slates') and they have been altered by thermal metamorphism and
mineralisation.
2. Ordovician Volcanic Rocks came after the Skiddaw Group, indicating new environmental
conditions. This was a direct result of subduction of the Iapetus Oceanic crust. The Iapetus
was gradually becoming narrower as the Laurentian (northern Britain) and Avalonian (southern
Britain) tectonic plates drifted towards one another during the Ordovician period. During the
early Ordovician, Britain - like many of the Earth's continents - was situated in the southern
hemisphere. Northern Britain was part of the Laurentia (North American) continent which lay
about 15ºS and southern Britain was part of a small continent (Avalonia) at about 60ºS, much of
which was covered by a shallow sea. In between these fragments there lay about 4,000 km of
Iapetus Ocean! Over the next 100 Ma the Iapetus gradually closed as the rate of subduction
overtook that of seafloor spreading. By the Devonian (418 Ma ago) the ocean had closed
completely.
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This igneous activity is a direct result of this subduction and it indicates that subduction took
place beneath the southern fragment of the British Isles. Early igneous activity was extrusive
and produced large volumes of andesites, ashes, ignimbrites and other volcanic material. Such
volcanism is typical of volcanic island arcs, or convergent plate boundaries over subduction
zones. (The marine fossils associated with nearby sediments also indicate an island arc
setting.)The oldest eruptions are recorded in the Eycott Volcanic Group. These were partly
erupted under water, as some of the last Skiddaw Group sediments are interbedded with them.
They differ chemically from the later groups - being mainly basalt and basaltic-andesite lavas.
The rocks of the Borrowdale Volcanic Group (BVG) come after the Eycott Volcanic Group.
These represent the main volcanic episode. They are the real heart of Lakeland - extending
from Wasdale and the Duddon Valley in the west through Scafell to Helvellyn, High Street and
Haweswater in the east. Over 6,000 m of these rocks were erupted in only 10 million years
during the mid-Ordovician. Today they cover an area of 800 km².
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The volcanic scenario forming them would have resembled that of Mount St Helen's (Cascade
Ranges, USA). The evidence shows that the volcanoes erupted above water, on land, but some
of the material collected in short-lived freshwater lakes. Streams then eroded fresh piles of
volcanic material, transporting it and reworking the debris. Two major phases can be
distinguished in the BVG. The first phase in which 2,500 m of rock accumulated, was dominated
by lavas flows with only occasional explosive eruptions. There were several vents or centres of
activity, but these are notoriously difficult to locate. The lavas, usually andesitic, flowed across
the landscape. These flows were often fairly localised, their extent and thickness depending on
the precise composition of the magma and the local relief. There were occasional explosive
eruptions forming air-fall tuffs. The second phase marked an abrupt change to explosive
activity with ignimbrite sheets, pyroclastic surge and fall deposits. Enormous volumes of
material were erupted very quickly, and with violent eruptions of andesite, dacite and rhyolite,
much of which fell back to the surface as air-fall tuffs. Much has now been removed by erosion.
Large parts of the volcano broke away, tilted and sank back into the molten volcanic pile 'Caldera collapse'.
3. Windermere Supergroup rocks lie on top of the volcanic rocks and include more than 5,000
m of very varied sedimentary material. The succession ranges in age from the late Ordovician
to the end Silurian. These rocks form the third major belt of the old core, extending across
the southern Lake District (the Duddon Estuary to the Howgill Fells). Towards the end of the
Ordovician the Iapetus Ocean had almost closed, the Skiddaw Group rocks had been folded and
the Borrowdale Volcanic Group uplifted. This is indicated by the lowest unit of the Windermere
Supergroup, the Dent Group (e.g. the Coniston Limestones) which lies above an unconformity an erosion surface cutting down into the BVG and the Skiddaw rocks. These two groups must
have been folded and uplifted prior to the deposition of the Dent Group. It outcrops between
the Duddon Estuary and Shap. Besides the thin Coniston limestones which suggest shallow
marine conditions, this group contains some volcanic materials from the last dying throes of the
Ordovician volcanoes. During the Silurian there was continuous sedimentation and the seas
deepened once more as subsidence occurred, perhaps due to the formation of a back-arc basin
behind the main volcanic arc. Thick turbidite sequences of sandstones, greywackes, siltstones
and mudstones were deposited, and can be seen on the present A6 near Shap. These have also
been folded subsequently due to the final closure of the Iapetus Ocean.
4. Granite Intrusions are the fourth and very distinctive group of rocks in the Old core. They
range in size from substantial plutonic intrusions such as the Ennerdale and Eskdale Granites to
small plugs, dykes and sills. These intrusions all join up deep underground to form one large
batholith. The Shap Granite is a well-known outcrop on Shap Fell and has been dated
radiometrically to 393 Ma (+/- 4Ma). These intrusions are post-collision granites formed by
partial melting of continental crust thickened during continental collision. Most will have been
emplaced at 5 km depth or greater and are now exposed due to erosion and uplift.
North Wales
The Cambrian rocks of North Wales are represented by a thick sequence of marine sandstones
and shales (the shales have since been metamorphosed into slates). Many of the sandstones,
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such the Rhinog Grits, are greywackes formed from sediments deposited in a turbidity current.
Above each greywacke layer the sediments grade into fine-grained shales and now form the
Llanberis Slates. The succession continues into the volcanic rocks of the Ordovician. These
rocks form an arcuate outcrop from Snowdonia around east of the Harlech Dome and towards
Cadair Idris. These volcanic rocks are formed from a volcanic island arc and contain acidic
rhyolites and intermediate andesites. In some areas many of the volcanic rocks are welded ash
flows called ignimbrites. These were laid down by high-temperature pyroclastic flows which
raced down the volcano flanks at high speeds. Subsequently all these sediments have been
folded up into large upright anticlines and synclines due to tectonic activity. In the Harlech
Dome area the rocks have been upfolded with the shape of an upturned dish, hence the term
dome.
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VARISCAN OROGENESIS – UPPER PALAEOZOIC ~ 300 MYA
SW England
In the British Isles, deformation and metamorphism associated with the Variscan Orogeny was
most strongly felt in south-west England (Devon and Cornwall), south Wales (Milford Haven to
the southern edge of Glamorgan) and the south of Ireland (Counties Cork and Kerry). The
Variscan Orogenic Belt consists primarily of Devonian and Carboniferous strata, with older
metamorphic rocks at Lizard Point and Start Point. The Rheic Ocean separated the British
Isles and northern mainland Europe (once Avalonia but now joined with Laurentia) from the rest
of Europe during most of the Palaeozoic. Towards the end of this Era, the Rheic Ocean closed,
with collision of Gondwana and Laurentia. Continental collision resulted in deformation and
metamorphism, followed by the emplacement of post-orogenic granites. What is the evidence
for this?
1. The evidence for the existence of an ancient ocean to the south of the British Isles during
the Palaeozoic can be found near Lizard Point. This is an Ophiolite complex (radiometrically
dated to ~ 400mya, Early Devonian); representing a section of obducted oceanic crust (basaltsgabbros) and upper mantle (Peridotite).
2. The large granite intrusions in south-west England could either be the product of subduction
associated with ocean closure or formed by partial melting of collisionally-thickened crust. The
Land's End to Bodmin Moor granites are emplaced in Lower to Upper Devonian rocks, whereas
the Dartmoor granite cuts across boundaries in the Upper Carboniferous rocks. The granites
must be younger than the sediments that they intrude. This gives a maximum age of Late
Devonian for the Land's End to Bodmin Moor granites and the Late Carboniferous for the
Dartmoor granite. As no contact relationship is visible between the granites and younger rocks
a minimum age of emplacement cannot be obtained in this way. A more accurate age of
emplacement of ~280-270 mya (Early Permian) has been obtained for all the granites by U-Pb
radiometric dating. Furthermore, geophysical evidence (gravity anomalies) reveals that below
the surface, the separate granites are actually one continuous batholith. Petrographic,
structural and field-based studies carried out on these granites show they are undeformed.
These facts imply that the granites were intruded after the surrounding Upper Palaeozoic
(Devonian & Carboniferous) sediments had been deformed by the Variscan Orogeny. In other
words, the granites are post-orogenic formed by partial melting of collisionally-thickened crust.
Using this age-related evidence, the final stage of Variscan deformation must therefore have
occurred after the Upper Carboniferous sediments were deposited, but before the Early
Permian granites were intruded.
3. The intense folding of Devonian and Carboniferous rocks in the region is also indicative of
collision. The Variscan folding in each of these areas is as follows:

St Govan's Head: WNW-ESE trending fold axes with steep dips (as shown by narrow
outcrop widths), several tight anticlines and synclines.
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
South Wales & Gower Peninsula: the south Wales coalfield is a broad syncline, the
visible closure at the eastern end indicating a basin-like structure. Steeper dips on
the southern limb of this syncline reflect the northwards-directed Variscan
movements. The Gower Peninsula shows folding similar to south-west Wales, but the
slightly broader outcrop widths are indicative of less steep dips.

South-west England: a very broad syncline is present with Devonian outcrops in north
Devon and Cornwall separated by a central belt of Carboniferous. The outcrop strikes
indicate a general E-W trend to the axis of the syncline.
From these structural analyses it is evident that Variscan deformation was caused by N-S
crustal shortening. This produced a series of asymmetrical folds with E-W trending fold axes
across the southern British Isles.
4. The Lizard Thrust indicates the existence of a thrust belt as one continental crust collided
with another.
5. Low grade regional metamorphism of the fine-grained muds deposited in the Devonian and
Carboniferous into slates such as can be found at Delabole also indicate continental collision as
an ocean closes.
3. Intense Folding
e.g. St Govan‘s Head
5. Regional Metamorphism
e.g. Delabole Slate
1. Ophiolites
e.g. Lizard Ophiolite Complex
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2. Batholiths
e.g. Dartmoor &
Bodmin Moor
4. Thrust Faulting
e.g. Lizard Thrust Fault
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SOUTH
WALES
Gentle open folds, Broad Haven,
South Wales
• Gentle folds
• 2% shortening of crust
• No cleavage
SOMERSET
DEVON
Less intense, more irregular
folds, Bude, Devon
• Complex folding & thrusting.
• 50% shortening of crust
• Good cleavage
CORNWALL
Recumbent folds, Millhook Haven,
north Cornwall
• Very complex folding & extensive
thrusting.
• 65% shortening of crust
• Very strong cleavage
OPENING OF THE ATLANTIC OCEAN – CENOZOIC ~ 65 MYA
NORTH ATLANTIC IGNEOUS PROVINCE (NAIP)
By the beginning of the Tertiary (65 Mya) a mantle plume (the Iceland mantle plume) intruded
into the base of the continental crust in the area between Britain and Greenland (both part of
the Laurasian continent). By 63 Mya the mantle plume initiated a huge area of the Laurasian
continental crust to be thermally uplifted, over 2000 km across. The plume produced extensive
igneous activity from NW Britain to the west coast of Greenland and up to 2 km of uplift at its
centre. The first pulse of igneous activity occurred 60 Mya with eruptions from fissures onto
the Laurasian continental crust of extensive sheets of basaltic lava, known as flood basalts.
This was then followed 58 Mya by large volcanoes erupting more lavas onto the land. The roots
of these volcanoes can be seen in Skye, Mull & Ardnamurchan as outcrops of granite & gabbro.
Igneous material was also injected into the country rocks as dykes, often they were so
numerous as to occur in what are called dyke swarms. 55 Mya continental rifting over the hot
spot above the Iceland mantle plume eventually caused the Laurasian continent to rift and seafloor spreading was initiated and the North American plate started to drift away from the
Eurasian plate. As ocean-floor spreading started the North Atlantic opened and by 50 Mya a
constructive plate margin had fully developed and can now be seen as the Mid-Atlantic Ridge.
The mantle plume hot spot continues to be active (although at a reduced rate) to the present
day and is responsible for the continued existence of the volcanic island of Iceland (formed 25
Mya).
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The igneous activity was localised at a number of centres (Skye, Rhum, Ardnamurchan, Mull,
Arran and Antrim) falls into three main types: extrusive igneous activity, intrusive igneous
activity and tensional dyke swarms.
In many places the roots of the old volcanoes and magma chambers have been exposed by
erosion and have allowed classic work to be carried out on the details of volcanic and igneous
processes. This is particularly true on Skye (see diagram over page), Mull, Rhum and
Ardnamurchan.
1. Extrusive Igneous Activity:
Extensive upwelling of flood basalts in NE Ireland (Antrim) & NW Scotland (northern Isle of
Skye). Low viscosity flood basalts flowed substantial distances to form horizontal plateaux &
distinctive flat topped hills. On northern & western Skye the skyline dominated by the stepped
landscape created by the lava flows. Each lava flow from 10 to 15m thick (4 - 8m in Skye) and
overall reached thicknesses of 2000m (although most has now been eroded off). In Antrim
1500 square miles of lava still preserved. The lavas are found on the sea bed across to
Northern Ireland, where they form the spectacular Giant's Causeway with its hexagonal basalt
columns. These are also found at Fingal's Cave on the Isle of Staffa. The 6-sided columns are
the result of the perfect cooling of a basalt lava. As it cooled it hardened, cracks/joints
formed in a perfect mathematical pattern. The tops of the lavas often have many gas bubble
holes (vesicles) filled in with minerals (amygdales). Weathering of lava flows in the hot, humid
climate of the early Tertiary Period caused red laterite (iron-rich soils) to form.
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These lavas erupted on land from fissures similar to those found in Iceland today. These
fissures formed as the continental land mass of Pangaea was uplifted and domed over a 2000km
area due to the thermal uplift of the Iceland mantle plume that had been initiated due to
thermal instabilities on the core-mantle boundary some 60 mya. The initial surge of material in
a mantle plume develops into a large spherical head that rises through the mantle, followed by a
much narrower tail. This plume head eventually flattens out to form a disc of hot material
below the lithosphere. It has been suggested that flood basalt areas result from the larger
volume plume head whilst later lower levels of extrusive activity are the product of the more
continuous but less voluminous plume tail. The heat from the plume causes the thermal uplift of
the continental land mass above. This causes extra erosion and along with stretching and
extension of the lithosphere associated with the plume causes the lithosphere to thin. Thinning
of the lithosphere allows the ductile asthenosphere to rise nearer to the surface where
pressures are reduced. As pressures reduce decompression melting occurs and the hot
peridotite material partially melts to form a "basaltic" magma.
2. Intrusive Igneous Activity:
The central igneous intrusive complexes (Skye, Mull, Arran, Ardnamurchan, Rhum) are
dominated by large, coarse-grained plutonic masses of gabbro, ultra-basic rocks and granites,
surrounded by arcuate (arc-shaped) suites of smaller intrusions, ring dykes, cone sheets, etc.
The ultra-basic rocks, usually peridotites, and the gabbros often show remarkable layered
structures. This layering is caused by crystals of different melting points and densities settling
out by gravity during crystallisation of the magma. This remarkable feature gives the igneous
rock an almost sedimentary appearance. The gabbros and ultra-basic rocks represent magma
chambers formed up to 4km down and since exposed by erosion. They are associated with
volcanic vents and collapsed calderas representing surface volcanicity which occurred early in
the development of the centre. This early volcanic activity was followed by major intrusions
some of which were granitic in composition. These granites formed due to partial melting of the
basement rocks and fractional crystallisation of basic magmas. On Skye there is the well-known
contrast between the dark gabbro of the Cuillin Hills and the lighter granites of the eastern
Red Hills. The gabbros and ultra-basic rocks are older than the granite intrusions, which in
places have a roof of earlier plateau basalt. Many of the gabbros show a layered structure, and
the whole mass of gabbro is made up of a number of intrusions forming rounded outcrops. The
Red Hills are composed of a number of granite intrusions, again often resulting in rounded
outcrops. The gabbros are really part of a deep magma chamber exposed by erosion and
scattered about the Cuillins and the Red Hills are the remains of old volcanic vents with ashes
and lavas.
3. Tensional Igneous Activity:
Rising plutons of basic magma caused localised uplift and consequently tension and crustal
extension which formed cracks and fractures in the rocks. Magma has later risen up these
fractures to produce dykes. In the Cuillins, many of these dykes occur as cone sheets. These
are sloping sheets which have the form of huge cones several kilometres in diameter dipping
inwards towards the deep core of the volcano. The localised rising of magma causing updoming,
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stretching and fracture of the country rocks can also allow the central block to sink, a process
known as cauldron subsidence. Magma then rises up the fractures to feed volcanoes at the
surface forming concentric intrusions known as ring dykes. Both these features can be seen on
the peninsula of Ardnamurchan in NW Scotland (see map below).
As rifting continued the crust was stretched and thinned allowing dyke swarms to form. These
are basalt and granitic dykes occur all over the Tertiary igneous province of western Scotland,
northern England and North Wales. They have a general NW to SE trend, reflecting a NE to
SW extension of the crust associated with the early opening of the North Atlantic. On Arran
300 dykes are exposed in a 3km coastal section on the south of the island, indicating a 10%
crustal extension. The Cleveland Dyke in north-east England may be derived from a magma
chamber on Mull and was emplaced as a single pulse in a matter of days. It can be traced 400km
from Mull through Cumberland to Durham and is dated at 58 Ma.
How does a mantle plume form?
Laboratory experiments suggest that a starting plume initiates as a thermal instability at the
core-mantle boundary. With time, hot material at the boundary layer begins to upwell and rise
through the mantle. This initial surge of material develops into a large spherical head that rises
through the mantle, followed by a much narrower tail. The plume head is initially a few hundred
kilometres across but expands by picking up surrounding mantle material and eventually flattens
out to form a disc of hot material
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By the beginning of the Tertiary 65 Mya the
Laurentian continent between Britain and Greenland
was undergoing a period of rifting and extension due
to the initiation of a mantle plume (an upwelling of hot
material form deep within the mantle). The
emplacement of a large plume head beneath the
lithosphere will produce uplift on a large scale due to
the additional buoyancy provided by heat from the
plume. This uplift causes extension and thinning of the
lithosphere, and subsequently large volumes of basaltic
lavas are erupted due to decompression melting of the
asthenosphere.
A current popular theory is that flood basalts (vast outpourings of basaltic lava) are related to
the impact of mantle plumes. All modern-day examples of plume-related magmatism such as in
Hawaii are the surface manifestation of mantle plumes that have a long eruption history of
volcanism produced at a more or less constant rate. For example, eruption rates along the
Hawaiian chain have remained between 0.1 and 0.2 km3/year for the last 10 million years.
However, mantle plumes do not last forever – they have to begin at some time and eventually
die; and, while the complete history of a mantle plume has not been recorded in detail, they
probably last for 100 million years or more.
How does a mantle plume begin? Laboratory experiments using suitable materials suggest that a
starting plume initiates as a thermal instability at the core-mantle boundary. With time, hot
material at the boundary layer begins to upwell and rise through the mantle. The initial surge of
material develops into a large spherical head that rises through the mantle, followed by a much
narrower tail. The plume head is initially a few hundred kilometres across but expands by
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entrainment of surrounding mantle material and eventually flattens out to form a disc of hot
material below the lithosphere. The overall profile is like a growing mushroom and the dynamics
are similar to those that operate in the atmosphere
during the eruption of large volcanic plumes or the
mushroom clouds of nuclear explosions. The initial
plume head is followed by a plume tail, a narrow jet
of hot material. It has been suggested that
continental flood basalts result from the large
volume plume head while oceanic islands such as
Hawaii today are the product of the more continuous
but less voluminous plume tail.
the tank.
This photograph shows a laboratory analogue of the
initiation of a mantle plume. The original laboratory
experiment involved heating the basal layers of a
tank of two viscous fluids. The lower layer is slightly
denser than the upper layer but when heated it
expands and its density decreases. Eventually, its
density drops below that of the upper layer and a
large mass of material starts to rise to the top of
ALPINE OROGENESIS – CENOZOIC ~ 40 MYA
The fold mountains of the Alps were formed during the Tertiary as Africa closed with Europe,
and the Tethys Ocean crust was subducted back into the mantle. It is often assumed that
ripple effects from this Alpine orogenesis caused the Tertiary folding of southern England,
which appears to have occurred 20 mya. However, although the main compression in the
southern Alps was at about this time, 20 mya, the eastern Alps were mainly formed about 40
mya. Also it is not certain that crustal movements as far north as Britain could have resulted
from Alpine compression. The structures formed over southern England could reflect crustal
readjustments following the initiation of sea-floor spreading in the North Atlantic, particularly
with the movement of Greenland away from Europe and the formation of the Bay of Biscay.
Tertiary folding in southern England is particularly well known in south Dorset, including the
Purbeck area, where dips are often near to vertical, and this fold belt can be traced eastwards
into the Isle of Wight. In the Purbeck area and south Dorset, rocks trend in an east-west
direction and have been tightly folded to form the Chaldon and Poxwell Periclines and the
Purbeck Monocline. A pericline is a steep anticline in which the dips are inclined in all directions
away from the centre, a bit like an upturned boat. A monocline is a single fold with a very steep
dip in an otherwise uniformly dipping sequence. In the Purbeck area and south Dorset, a number
of faults trend east-west and are the same age as the folding.
A number of other Tertiary folds and faults, all trending east-west, affect the Hampshire
Basin area, and the Chalk areas further north. These include the Portsdown Anticline, the Hog‘s
Back Anticline, the mere Fault, and the Pewsey-Kingsclere Anticline. Many of these fold
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structures and faults involve steep monoclines, and are associated with faults in the underlying
Variscan basement.
The main Weald Anticline is a Tertiary fold structure, as are the associated synclines to the
north and south, the London and Hampshire Basins. The Weald Anticline is really an inversion
structure where a subsiding basin in the Jurassic, Cretaceous and early Tertiary changed to be
an uplifted dome during the late Tertiary.
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GEOLOGICAL MAPWORK
The AS GL2 Geology unit introduced interpretation of subsurface geological structure using
simplified geological maps. In this sub-unit, map interpretation skills are developed further by
investigating outcrop patterns on real geological maps.
This sub-unit aims to develop:
Extended skills in interpreting structural information contained in geological maps;
Skills in extracting information from geological maps related to a range of geological
applications;
An awareness of the environmental and sustainable issues involved with geological applications.
Key Idea 1: Outcrop patterns on geological maps can be used to identify and interpret
structural elements
Outcrop patterns of dipping strata and faults in relation to topography:
Direction of closure of V-shaped outcrops in valleys as an indication of dip direction;
Close parallelism of geological boundaries and topographic contours as a sign of near horizontal
dip;
Linear geological boundaries crossing topographic relief as an indication of steep dip.
Key Idea 2: Geological maps contain information relevant to a wide range of geological
applications
Geological maps provide an essential database of detailed information about the distribution of
rocks at the surface that can be used to interpret or predict subsurface geological conditions,
and for geological applications:
Design of construction projects;
Identification of geological hazards;
Location of resources - groundwater, fossil fuels; alternative energy sources;
Identification of environmental issues from extraction of these resources;
Assessment of suitability for sustainable waste disposal.
INTRODUCTION TO BRITISH GEOLOGICAL MAPPING
The Geological Map of Britain and Ireland opposite has been compiled by reduction and
simplification of data from many maps covering much smaller areas in great detail. The key
shows the sedimentary rocks are divided into four main groups – Cenozoic, Mesozoic,
Palaeozoic and Upper Proterozoic – which correspond to particular spans of time in the
stratigraphic column. The first three groups are called Eras of geological time and together
constitute the Phanerozoic Eon (visible life). The time before this is known as the Cryptozoic
Eon (hidden life) or alternatively the Precambrian.
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The stratigraphic column is one of the most important features of the key on any geological
map; it is always arranged in stratigraphic order, with the oldest rocks at the bottom. However,
there are two departures from this; the metamorphic rocks are divided into two main groups
which do not appear to relate to periods in the stratigraphic column, and the igneous rocks are
not classified by age, but are divided into intrusive rocks and volcanic rocks. The metamorphic
rocks are divided on the basis of the age of the metamorphic event which affected them, and
these events do not necessarily correspond exactly to particular periods. Igneous rocks are
divided into only two categories to simplify the map, and not because their ages are known.
Remember that although each geological period is represented on this simplified geological map
of Britain by a single colour, it may include a variety of rock types. For example, the Silurian
(lilac) includes shales, mudstones and limestones, and the Jurassic (olive-green) includes
limestones and clays. This simplification is necessary to be able to show the geology of Britain
on such a small map. For the A2 level exam you need to familiarise yourself with the names of
the eons, eras and periods of the stratigraphic column shown here, although you do not need to
remember any of the numerical dates!
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Cartography is the science and art of drawing maps. Hills, valleys and other undulations of the
land surface, are known as relief, and maps showing these details are called topographic maps.
In 1791, the Board of Ordnance was founded under the control of the British Army. The work
of the Board of Ordnance introduced a systematic approach to mapping: a detailed survey of
the country was carried out and heights of prominent hills were related to sea-level. The most
important step of all in topographic mapping soon followed: the representation of relief on maps
by contours drawn at regular intervals. Each contour on the map represents the intersection
between the topography and an imaginary horizontal plane at a specific height above mean sealevel. Mean sea-level is used as the reference plane, and is known as ordnance datum (OD).
Later advances in cartography included the standardisation of map scales, and the
establishment of the National Grid system of reference for Britain that enables particular
features to be located easily. Eventually the Board of Ordnance was renamed the Ordnance
Survey (OS) and a parallel organisation, the Geological Survey of Great Britain was established
to plot the geology on the new ordnance maps. In 1985 the Geological Survey of Great Britain
was renamed the British Geological Survey (BGS).
Any point in the UK can be located on a map by means of its grid reference. To be able to
answer any question in the GL4 Mapwork section you need to be able to use grid references.
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The National Grid is in effect a grid overlain on a map of Britain. This grid is divided into
100km grid squares, with each grid square identified by two letters. For example, in the UK
National Grid map above, the grid square that encompasses the area around Land‘s End in
Cornwall is SW, while the grid square that encompasses Edinburgh is NT. From the map you may
have noticed that the country appears to be split in two with most of the 100km grid squares in
Scotland and northern England start with the letter N (except for Orkney and Shetland which
begin with H), while those in southern England with S.
The east-west running lines of the grid are referred to as northings while the north-south
running lines are called eastings. The 100km grid squares are further divided into smaller
squares by grid lines at a 10km spacing, each numbered 0 to 9 from the south-west corner in an
easterly (left to right) and northerly (upwards) direction. You probably remember the saying
―along the corridor and up the stairs‖ when you were taught this at school. The diagram below
shows a full example of how this system works using the 100km grid square TL. Using this
system, you can identify a 10km square grid by giving first the two-letter code for the 100km
grid square followed by the easting and then the northing on the 10km grid. So for example, in
diagram below identifies the 10km grid square TL63.
On 1:50,000 OS maps, you can find the two grid letters on the key or on the corner of the
maps, and at this scale the grid has been further divided into 1km intervals, as shown in (c) in
the diagram below. As with the 10km grid, we begin in the south-west corner and quote first
the eastings then the northings. The example in (c) on the diagram on the previous page
identifies the 1km grid square TL 6432. By estimating the distance between the 1km grid lines,
you can specify a position to within 100 metres, so for example the 100 metre grid reference
for the point shown in (c) above is TL 643327.
The geological maps we use at A2 level are all published by the BGS. These are printed in colour
and represent interpretations of where the solid rock, the bedrock, would occur on the Earth‘s
surface were all the soil, vegetation and buildings to be removed. The colours on these BGS
maps are separated from each other by sharp boundaries. These geological boundaries are lines
showing where surfaces separating different kinds of rocks in the ground intersect the
topographic surface. The scale which the BGS has made most use of in Britain for published
maps is 2cm to 1km (1:50,000). The 1:50,000 Moreton-in-Marsh Sheet we will be using is an
example. The BGS has also produces 1:25,000 geological sheets such as the Cheddar Sheet
which we will be using.
THE TEN MILE MAP
As an introduction into geological mapping we will use the BGS Ten Mile Map, so named because
its scale of 1:625,000 is approximately 10 miles to 1 inch. The whole of Britain is covered in two
sheets, north (N) and south (S), showing the geology of most of Northern Ireland as well as
Britain. The maps are described as showing the solid geology. This indicates that the map does
not show any superficial deposits which includes all unconsolidated material deposited during
the Quaternary Ice Ages and all other materials deposited since, such as river gravels and
alluvium, peat and beach sands.
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Around the map itself lots of information is given. The sedimentary rocks are set out in
stratigraphic order. Eras and periods are marked down the right-hand side of the key. Each
sedimentary rock has a number from the oldest (Precambrian, 60) at the bottom to the
youngest (Pleistocene, 115) at the top. The individual coloured boxes in the key up to and
including the Carboniferous Period represent time subdivisions of the periods. Thus Caradoc
(70) represents a time interval during part of the Ordovician, and Namurian (81) a time interval
during the Carboniferous. Neither name gives any clue to the types of rock to be found on the
ground.
After the Carboniferous, the system used on the key changes and now represents different
types of rock strata, rather than only periods of time, so it is called a lithostratigraphic (rock
stratigraphic) column. It can be divided into rock units for which there is a separate hierarchy
of terms. The smallest unit (too small to show on the Ten Mile Map) is an individual bed of a
particular rock type which is separated from those above and below by bedding planes. Beds
may be grouped together into units called members and members can be grouped into
formations. Formations may be amalgamated to form groups that have broadly similar
characteristics. In the example given below, the Bracklesham Group is made up of three
formations, the Camberley Sand Formation, the Windlesham Formation and the Bagshot
Formation, with a number of members in each formation.
The lithostratigraphic system is not used for the Carboniferous and earlier periods because
well-defined rock units of a single type are not readily traced for long distances. In different
areas, different rocks were generally being deposited at the same time so that strata of the
same age are represented by different kinds of rock in different parts of the country.
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The term used to describe the general characteristics (colour, texture, mineralogy, etc.) of a
sedimentary rock is lithology. Rock units occurring at different localities, which may be of
different lithologies, are said to correlate if they can be shown to be of the same age.
Correlation of different rock outcrops between different areas of the country cannot
therefore rely on recognising a distinctive rock type. The complex mixed terminology in the key
arose because until recently there were no generally agreed principles for assigning names to
rock units for stratigraphic purposes.
The key to igneous rocks is considerably simpler than that for sedimentary rocks. The igneous
rocks are grouped into intrusive and extrusive (volcanic) types. The extrusive rocks are broadly
grouped according to the time during which they were erupted, and the coloured boxes are
labelled according to the type of igneous rock represented. Thus, you cannot tell from this map
whether an outcrop of basalt coloured pink and labelled 42 is of Silurian, Ordovician or
Cambrian age. In other words, this part of the key is not a stratigraphic column in the strict
sense. The intrusive igneous rocks are not divided by period at all, and are coloured and
numbered by rock type. Finally the metamorphic rocks are also labelled descriptively. The last
items on the key are major faults and major thrusts.
Turning to the North Sheet of the Ten Mile Map, you will see that the key reflects the
differences in geology between the north and the south of Britain. There are fewer
sedimentary formations although they cover a similar time span. The range of igneous rocks is
essentially the same as the South Sheet but with the important difference that extrusive
igneous rocks of the Tertiary Period are present in Scotland particularly on the islands of Skye
and Mull. There are no areas shown as Cambrian or Precambrian extrusive rocks. This is because
in this area rocks of these ages have been metamorphosed after the Cambrian Period and so
are now classified as metamorphic rocks. The key to metamorphic rocks is much more detailed
than the South Sheet, reflecting the fact that the greater part of the Highlands of Scotland
are composed chiefly of these rocks along with major igneous intrusions. Another prominent
feature of the North Sheet is the more or less regular pattern of red lines of basaltic rock
which occur in the whole of western Scotland from northern Skye to the south-eastern
boundary of the Dalradian Complex. These igneous rocks have been intruded into the rocks of
the Highlands and Islands along steep or vertical cracks, and have solidified to form thin
curtains of rock known as dykes.
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A2 MAPWORK RULES
1. If a geological
boundary is a
straight line across
the contours, then
the boundary is
vertical or nearly
vertical (e.g. dyke
or fault).
2. If a geological
boundary follows
the contours, the
beds are
horizontal.
3. The steeper the
dip of the beds,
the narrower the
outcrop of the
formation.
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4. Beds dip
towards younger
rocks.
6. In an anticline
the oldest beds
occur in the centre
(core) of the fold.
7. In a syncline the
youngest beds
occur in centre
(core) of the fold.
8. In a valley, the
geological
boundaries ―V‖ in
the direction of
dip of the beds.
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9. In a valley, the
longer the ―V‖, the
shallower the dip
of the beds.
10. In an angular
unconformity the
beds above have
different strike
and dip directions,
and different dip
angles, and cut
across the
structures below.
11. If a fold shown
on a map has a
closure or nose,
the fold is
plunging. In an
anticline, the fold
closure or nose
points in the
direction of plunge.
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12. True dip is
shown by dip
arrows. Beds drawn
at 90° to the true
dip appear
horizontal.
Apparent dip is
always less than
true dip.
13. Outliers are
areas of younger
rock surrounded by
older rock. Inliers
are areas of older
rock surrounded by
younger rock.
14. Normal faults
dip steeply in the
direction of the
downfaulted block
and have a straight
outcrop pattern on
the surface.
15. Reverse faults
dip at around 45º
and have a straight
outcrop pattern on
the surface.
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16. Thrust faults
are nearly
horizontal and may
have a curved or
undulating outcrop
pattern on the
surface.
17. Strike slip or
tear faults are
vertical.
18. In dip faults
(normal, reverse &
thrust), where two
beds are in contact
across the fault,
the younger bed is
on the downthrow
side.
19. In dextral tear
faults (strike slip)
rocks move to the
right relative to
each other.
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20. In sinistral
tear faults (strike
slip) rocks move to
the left relative to
each other.
21. The throw of a
fault is the
vertical
displacement
between beds, and
can be measured
using the vertical
scale.
22. The oldest
rocks are usually
on the upthrow
side of a fault
separating rocks of
different ages.
23. Two parallel
fault lines with
their downthrow
sides on the same
side indicate a rift
valley (graben).
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24. Two parallel
fault lines with
their downthrow
sides on opposite
sides indicate a
horst (upthrown
block).
25. An anticline
that plunges at
both ends is known
as a dome. A
syncline which
plunges at both
ends is known as
basin. These
structures both
produce circular or
elliptical outcrop
patterns.
26. Boundaries of
plutonic intrusions
(batholiths) are
near vertical and
may undulate, and
are discordant
with the
surrounding
country rocks.
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27. Sills have an
outcrop pattern
which is curved,
wide, and
concordant and
follows the
topography.
28. Dykes have an
outcrop pattern
which is linear,
narrow, and
discordant, and cut
across topography.
29. A drift geology
map shows the
loose,
unconsolidated
geological
materials, such as
sand, peat, alluvium
and glacial
material, which
overlie bedrock.
These are known as
drift or superficial
deposits.
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30. Maps that
exclude drift or
superficial
deposits
altogether are
known as solid
geology maps. Solid
geology maps show
the consolidated
and lithified
materials i.e. rocks.
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