Download CHAPTERS FROM GEOLOGY

CHAPTERS FROM GEOLOGY
Edited by Zoltán Püspöki
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rocks
Solidification
Ortometamorphism
Ultrametamorphism
Olivine
Pyroxenes
Anorhite-Bytownite
Magma
(indifferentiated)
Magma
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Labradorite
Pyroxenes
Andesine
Amphiboles
Oligoclase
Biotite
Albite
Quartz
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14.0 - 14.4 A
Chapters from Geology
Chapters from Geology
for
Geographers, Geography teachers and Biologists
Edited by Zoltán Püspöki
University of Debrecen
Authors
McIntosh, Richard William
Püspöki, Zoltán PhD
Rózsa, Péter PhD
Revised by
Csorba, Péter PhD
Müller, Pál DSc
English text revised by
Semsei-Szekeres, Edit
Native reviewer
McIntosh, Richard Duncan
Supported by
Ministry of Education
World Language Programme
Oktatási
Minisztérium
Tempus Public Foundation
TEMPUS KÖZALAPÍTVÁNY
Terra-Mina Ltd.
CONTENTS
1. INTRODUCTION (TO GEOLOGY?).....................................................7
1.1 CLASSICAL GEOLOGICAL BRANCHES .............................................................................. 7
1.2 APPLIED GEOLOGICAL SUB-DISCIPLINES ......................................................................... 8
1.3 THE ROCK CYCLE AND ITS CONSEQUENCES .................................................................... 9
2. IGNEOUS ROCKS (PÉTER RÓZSA) .........................................................12
2.1 THE NATURE OF MAGMA .............................................................................................. 12
2.2 OCCURRENCE OF IGNEOUS ROCKS ................................................................................ 13
2.2.1 Types of intrusions............................................................................................... 13
2.2.2 Types of extrusions .............................................................................................. 15
2.3 STRUCTURE AND CRYSTALLISATION OF ROCK-FORMING SILICATE MINERALS .............. 17
2.4 PRINCIPAL TEXTURES OF IGNEOUS ROCKS .................................................................... 22
2.4.1 Principal textures of volcanic rocks .................................................................... 23
2.4.2 Principal textures of hypabyssal and plutonic rocks ........................................... 23
2.5 CLASSIFICATION OF IGNEOUS ROCKS ............................................................................ 24
2.6 CLASSIFICATION OF VOLCANICLASTIC ROCKS .............................................................. 28
2.7 SHORT DESCRIPTION OF THE MAIN IGNEOUS ROCK TYPES ............................................. 29
2.7.1 Plutonic rocks (figure 12).................................................................................... 29
2.7.2 Volcanic rocks ..................................................................................................... 30
3. PLATE TECTONICS (RICHARD WILLIAM MCINTOSH) ..........................32
3.1 THE CONCEPT OF PLATE TECTONICS ............................................................................. 32
3.2 DIVERGING MARGINS: RIFTS, MID OCEANIC RIDGES AND OCEANIC BASINS ................... 34
3.2.1 Formation of the rift systems ............................................................................... 34
3.2.2 Magmatic processes ............................................................................................ 34
3.2.3 Structural and global tectonic consequences ...................................................... 36
3.3 CONVERGING MARGINS: SUBDUCTION OF THE OCEANIC CRUST; ISLAND ARC SYSTEMS 37
3.3.1 Development of the converging margins ............................................................. 37
3.3.2 Magmatic processes ............................................................................................ 38
3.3.3 Forearc Region.................................................................................................... 39
3.3.4 Island arc............................................................................................................. 39
3.3.5 Back-arc Region .................................................................................................. 40
3.4 COLLISION, OROGENE REGIONS .................................................................................... 40
3.4.1 Development of orogene regions ......................................................................... 40
3.4.2 Magmatic and metamorphic processes................................................................ 41
3.5 TRANSLATIONAL PLATES ............................................................................................. 42
3.6 EVOLUTION TRENDS OF THE LITHOSPHERE ................................................................... 42
4 SEDIMENTARY ROCKS (ZOLTÁN PÜSPÖKI) ........................................45
4.1 ORIGIN, CLASSIFICATION AND GENERAL DESCRIPTION OF SEDIMENTARY ROCKS ......... 45
4.1.1 Formation of sediments and sedimentary rocks .................................................. 45
4.1.2 Classification and general characteristics of sedimentary rocks ........................ 46
4.1.3 The main sedimentary environments (sedimentary basins) ................................. 46
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4.1.4 Bedding................................................................................................................ 47
4.2 DESCRIPTIVE CHARACTERISTICS OF SILICICLASTIC ROCKS ........................................... 49
4.2.1 Determination and interpretation of grain size distribution................................ 49
4.2.2 Analysis and interpretation of grain morphology................................................ 55
4.2.3 Mineralogical composition of siliciclastic sediments .......................................... 57
4.3 WEATHERING OF SILICATE STRUCTURES ...................................................................... 58
4.3.1 The concept of weathering................................................................................... 58
4.3.2 Way and control of the alteration of silicate structures....................................... 60
4.3.3 Place and products of weathering ....................................................................... 61
4.4 CARBONATE ROCKS AND CARBONATE DEPOSITIONAL ENVIRONMENTS ........................ 64
4.4.1 Chemical and textural components of carbonate rocks....................................... 64
4.4.2 Classification of carbonate rocks ........................................................................ 65
4.4.3 The main environments of carbonate deposition ................................................. 67
4.5 ORGANIC MATERIALS COAL AND HYDROCARBON DEPOSITS ......................................... 71
4.5.1 Classification and lithology of coals ................................................................... 71
4.5.2 Palaeoenvironmental conditions of the coal deposites........................................ 72
4.5.3 Formation and migration of hydrocarbons ......................................................... 73
4.6 EVAPORITES ................................................................................................................. 75
5 METAMORPHIC ROCKS (ZOLTÁN PÜSPÖKI).......................................77
5.1 PROCESS AND CLASSIFICATION OF METAMORPHISM ..................................................... 77
5.2 THERMAL METAMORPHISM .......................................................................................... 78
5.2.1 Terminology of the thermal metamorphism......................................................... 78
5.2.2 Processes and mineral associations in the thermal metamorphism .................... 78
5.3 DYNAMOTHERMAL METAMORPHISM ............................................................................ 79
5.3.1 Classification and processes of the dinamothermal metamorphism – mineral
facies............................................................................................................................. 79
5.3.2 Stages of the dynamothermal metamorphism ...................................................... 80
5.4 THE UPPER BOUNDARY OF THE METAMORPHISM - ULTRAMETAMORPHISM ................... 83
6
1. INTRODUCTION (TO GEOLOGY?)
(Don’t skip over; it has been difficult to write!)
Student A: What on Earth is a Geologist?
Student B: Maybe something like a Theologian.
Student A: You mean he works on the phenomena of the soul?
Student B: And creates contact between the Earth and Heaven.
Teacher: Boys, the difference between a Geologist and a Theologian is like that between the
Earth and Heaven.
Geology tries to summarise the material and energy transport
processes of the Earth, especially of the earth crust, appearing in the
geological past and leading to the present face of our globe. In this relation
geology has to be able to describe the materials of the earth, the processes
producing these materials, and the history framing these processes.
To get a clear view of the most important fields of geology it is worth
giving the list of the main geological sub-disciplines with short explanations
on their main purposes.
1.1 Classical geological branches
Elementary fields (describing any organisation level of the material of the
earth crust):
• Crystallography
describes the chemical and physical-chemical
characteristics of the crystals and so the minerals
• Mineralogy
describes chemical variations, forming and
appearance of the different minerals
• Petrology
describes the mineral assemblages (rocks)
considering their appearance (petrography) and
formation (petrology)
• Sedimentology
as a “further-developed” part of petrology focusing
on sedimentary rocks it also touches upon their
formation and the description of palaeogeographic
conditions
• Structural geology analyses the structural elements and processes of
the earth crust on micro-, meso- and macro(global) scales
• Paleontology
in close connection with some fields of biology
(e.g. taxonomy, comparative anatomy) it describes
the living creatures of the geological past
7
Synthesizing fields (creating reconstructions for the development of the earth
crust)
• Stratigraphy
defines the relative or real age of the different
materials of the earth crust. Depending on its
method it may be e.g. relative stratigraphy (based
on the relative position of different formations),
biostratigraphy (based on the evolution lineages),
radiometric chronology (depending on physical
characteristics – half-period – of radioactive
elements)
• Historical geology compiles the results of elementary fields and
stratigraphy giving complex model for the
development of a certain part of the earth crust.
There are some very important applied fields of geology not only to
help the development of the economy, but also inspiriting the development of
the classical fields and so promoting us to understand our relationship with
the natural circumstances.
1.2 Applied geological sub-disciplines
• Economic geology
examines the formation of raw materials (ore and
non-ore deposits) to determine the economic and
environmental risks of mining
• Engineering geology determines the stability of the earth crust and
shallow geological formations (soils) to predict the
technological risks of building and mining
operations
• Hydrogeology
describes the amount, quality and movements of
the subsurface waters to determine the opportunity
and limitations of water mining
• Environmental geology determines the geological hazards on human
activities and anthropogenic (human) risks on the
natural environments – e.g. geological aspects of
waste management
• Agro geology
describing the mineralogical characteristics of soils
to determine the optimal (natural) ways to increase
soil fertility
So we can see that in order to understand the main aspects of geology
we need a great deal of courses like in the case of biology from taxonomy
through molecular biology to ethology (only to mention some distant fields
of this very important discipline).
8
1.3 The rock cycle and its consequences
In this short course we have no opportunity to give a general outline
of geology, we would like to clear only some general fields of it that can help
us to understand the most important processes of the earth crust.
To understand the materials of the earth crust we have to see that the
main chemical elements of the earth crust are oxygen, silica and aluminium,
together with iron, magnesium, potassium, sodium and calcium. These
elements form the different crystals of silicate structures, which are the most
important minerals of the earth crust. So if we would like to know the most
important minerals of the Earth, we can ignore diamond and ruby and have to
focus on silicate structures.
To summarize the most common processes of the earth crust we have
to know and understand the rock cycle (in the frame of plate tectonics)
(figure 1). We have to see that the material in the rock cycle is more or less
the same throughout the entire cycle. We can see the alteration of the silicate
structures according to the continuously changing physical factors (especially
pressure, temperature, and chemical content – referred to as P, T, C
conditions).
The “principal constituent of the magma is silica (SiO2) ranging from
35 to 75 %” (see later) so with the cooling and solidification (crystallization)
of the magma, rock forming silicate structures will be formed and will form
minerals assemblages called igneous rocks.
Under superficial conditions the rock forming minerals of igneous
rocks will be unstable phases and will be destroyed by the processes of
weathering. The previous silicate structures will be altered into different ones
becoming stable under superficial conditions. Sedimentary rocks (except
some special forms like limestones, evaporates and organic materials) can be
regarded as assemblages of these newly formed weathered silicate structures.
Under increasing pressure and temperature the silicate structures
become unstable phases again. They “loose” their natural water content, and
“change” their crystal structure becoming stable phase under the changed PT
conditions. Their alteration (called metamorphism if it takes place in solid
phase) leads to the formation of metamorphic rocks as assemblages of
metamorphic minerals.
There are some short circuits in the rock cycle (can be seen in figure
1) which have no important roles from the aspect of this introduction.
9
10
Surface, uppermost crust
Crust
Solidification
-crystallization
-formation of
amorphous
phase
Sediments
Magma
(indifferentiated)
Magmatic
rocks
Weathering
-disintegration
-chemical
decomposition
-transportation
-deposition
Parametamorphism
-recrystallization
Ultrametamorphism
-partial melting
-ion exchange
Metamorphic
rocks
Sedimentary
rocks
Magma
(differentiated)
Orthometamorphism
-recrystallization
Weathering
Resedimentation
Diagenesis
-compaction
-cementation
-loss of porosity
Figure 1 Stages, places and processes of the rock cycle
In conclusion we can state that the main materials of the earth crust
are the silicate structures, these minerals form mineral assemblages (rocks)
and the processes of the earth crust lead to the alteration of silicate
structures and so to the alteration of rocks. In this relation the rock cycle can
be regarded as the cyclic life of silicate structures and the history of the earth
crust forming rock cycles can be regarded as the history of the silicate
structures in the earth crust.
From this aspect – we hope – everybody can accept that to give a
short introduction to geology (materials and processes of the earth crust) we
had to choose the level of petrology. Since the rocks can be regarded as
mineral assemblages we can state follows:
• Rocks enclose minerals and not randomly but keeping the law of
associations that within the same association we can see members
tolerating similar conditions (PTC conditions in geology)
• The mineral associations (rocks) in their pattern and minerals content
can perfectly reflect the results of their forming processes.
To discuss the stages of these alterations we chose the order of the
rock cycle so the chapters will be about
•
•
•
Igneous rocks (followed by the plate tectonic model)
Sedimentary rocks
Metamorphic rocks
So finally we wish you a satisfying excursion into the “fields” of
geology and if you become attached to geology our work has not been in
vain.
the authors
11
2. IGNEOUS ROCKS
2.1 The nature of magma
Igneous rocks are formed by the solidification of molten or partly
molten mobile material termed magma. The nature of magma cannot be
directly observed because it originates from the partial melting of the lower
crust and upper mantle of the Earth, usually at depths between 50 and 200
km below the surface. Our understanding of the magma is based on
observations of volcanic activity, analyses and studies of different igneous
rock types, and on laboratory experiments made on synthetic magma-like
melts.
Most magmas are not entirely liquid but a complex mixture of liquid,
solid and gaseous materials. Magmas may contain crystals in a large portion
and this mixture moves slowly and sluggishly. Dissolved gases (mainly
water and carbon dioxide) usually constitute a small percentage of the
magma, however, they may reach as much as 14 % of the volume. These
volatiles are important because they strongly determine the viscosity and the
explosive characteristics of volcanic eruptions. Volatiles tend to decrease
the viscosity (hence to increase the fluidity) of the magma. Magmas rich in
volatiles tend to erupt more violently.
The principal elements of the magma are O, Si, Al, Ca, Na, K, Fe,
and Mg, however, the principal constituent of the magma is silica (SiO2)
ranging from 35 to 75 %. The silica content also strongly influences the
viscosity of the magma: the higher the silica content is, the greater the
viscosity will be. As a consequence, basaltic magmas containing about 50 %
SiO2 are characteristically fluid, whereas granitic magmas, which contain
60–70% SiO2, are thick and viscous.
Changes of principal elements (mainly silica and alkalies) in igneous
rocks show more or less distinct trends. A group of rock types that reflects
similar chemical tendency is called rock suite or rock series.
- Alkaline rocks consequently show a high concentration of alkalies
(Na, K) with respect to Ca and silica. The most common alkaline
rocks belong to the olivine basalt group.
- Calc-alkaline rocks Na+K exceed Ca only at the silicic end of the
series (SiO2 higher than 60%). This series comprises the common
volcanic basalt-andesite-dacite-rhyolite and plutonic gabbrodiorite-granodiorite-granite associations.
- Tholeiitic rocks are characterised by considerable iron enrichment.
There are some correlations between rock series and plate tectonic
settings. The general pattern is shown in table 1.
12
Table 1 General relationship between plate tectonic settings and igneous
rock series
Plate tectonic settings
Oceanic rift
Continental rift
Subduction zone
Intraplate
Igneous rock series
Mainly tholeiitic, marginally alkaline
Tholeiitic, alkaline
Calc-alkaline, tholeiitic
Alkaline, tholeiitic
2.2 Occurrence of igneous rocks
Igneous rocks may form below and/or at ground level. Intrusive
igneous rocks solidify in pre-existing rocks beneath the surface in the form
of so-called intrusions, whereas extrusive igneous rocks are formed by flow
of lava or deposition of volcanic clasts on the surface. Intrusive rock bodies
are usually divided into two further groups: shallow-seated intrusive rocks
are called hypabyssal, and deep-seated intrusive rocks are referred to as
plutonic.
2.2.1 Types of intrusions
There are several types of intrusions, and transitional ones may also
occur (figure 2a).
- Sills are tabular bodies of approximately uniform thickness and
are relatively thin compared to their lateral extent because they
are emplaced essentially parallel to the bedding or the foliation of
the country rock.
- Laccoliths are commonly mushroom-shaped intrusions that dome
up the overlying rocks; they occur in relatively undisturbed
sediments at shallow depths.
- Lopoliths are large, lenticular, centrally sunk basins or funnelshaped intrusive masses that are generally found in unfolded or
gently folded regions; their thickness is usually much less than
their width (one-tenth to one-twentieth).
- Phacoliths are associated with folded rocks. Following the crest
of an anticline and the depression of a syncline, they have the
shape of a double convex lens in cross section.
- Dykes are tabular intrusions that cut across the foliation or
bedding of adjacent rocks. They may occur alone or in swarms,
and are typically emplaced into pre-existing joint systems. Dykes
may also be associated with other shallow intrusions or volcanic
necks.
13
- Batholiths are large intrusive bodies, typically lacking any known
floor. Batholiths are usually composed of silicic rocks, and the
size of their outcrops ranges from one hundred to several
thousand square kilometres.
Figure 2 Forms of intrusive and extrusive magmatism
A Forms of intrusions
phacolith
laccolith
sill
dyke
lopolith
B Forms of extrusions
shield volcano
composite volcano
or
stratovolcano
caldera
dome
14
2.2.2 Types of extrusions
Extrusive rocks represent a wide variety of forms, depending on the
nature of the erupted material. Magma of low gas content and of relatively
low viscosity produces lava flows, while that of high gas content and of
relatively high viscosity tends to produce pyroclasts. The basic types of lava
flows are as follows (figure 3).
- Pahoehoe lavas have glassy, smooth, and occasionally rope-like
surfaces. The shape of the upper surface often resembles irregular
fold in cloth.
- Aa lavas have rough, fragmented surfaces. The fragments are
extremely rough and very irregular in shape.
- Block lavas are composed of relatively smooth fragments. The
surface of a block lava flow is quite irregular.
- Pillow lavas consist of ellipsoidal or pillow-shaped bodies. Each
pillow has a glass crust and shows radial jointing. This type of lava
is the most abundant of the types of flows because it is the most
common one on the sea floor.
Lava may emerge from either a long narrow fissure or a central
eruption. Fissure eruptions may form lava plateaus (in the case of fluid
basaltic magmas) or chains of volcanoes (in the case of less fluid lava).
Eruptions from a central vent may build a variety of volcanic forms,
depending on the fluidity of the magma and on the amount of the pyroclastic
material.
Fluid basaltic lavas may form a broad cone around the vent called a
shield volcano (figure 2b). These volcanoes have a wide base (often more
than 100 km in diameter) and gentle slopes. Their internal structure consists
of innumerable thin basalt flows with little ash.
The higher the viscosity of the magmas is, the more explosive and
violent the volcanic activity will be. Consequently, the more silicic the
magma is, the greater the amount of pyroclastic material will be. The
alternating layers of tephra (accumulated volcanic ash) and lava produce a
composite volcano, or stratovolcano, characterised by a high, steep-sided
cone. The summit area sometimes collapses and a large, basin-shaped
depression called a caldera may form.
Extremely silicic magma is so viscous that it forms massive plugs or
bulbous domes over the volcanic vent.
15
Figure 3 Some types of appearance of extrusive rocks
pahoehoe flow
volcanoes.usgs.gov/Products/
Pglossary/pahoehoe.html
aa flow
volcanoes.usgs.gov/Products/
Pglossary/aa.html
block lava
pillow lava
library.thinkquest.org/17457/ volcanoes/
hazards.lavaflow.php
volcanoes.usgs.gov/Products/
Pglossary/PillowLava.html
16
2.3 Structure and crystallisation of rock-forming silicate minerals
For igneous rocks the basic rock-forming minerals are the silicate
ones. Beside this, these minerals are by far the most abundant group of
minerals in the Earth’s crust. The basic building block of all silicate
minerals is the so-called silicon tetrahedron (figure 4), in which the
relatively small Si4+ cation fits in tetrahedral coordination with four O2anions, i.e. oxygen anions occupy the corners of a tetrahedron with a Si4+ in
the centre. Each O2- has one of its two valence charges satisfied by bonding
with the Si4+ at the centre, therefore an isolated silica tetrahedron has –4
charges. This pattern, however, has another consequence: the silicon
tetrahedra may polymerise because the remaining –1 charge of the oxygen
anions is available to bond with a Si4+ in the centre of another tetrahedron.
Figure 4 Structure and polymerization of the silicon tetrahedron
silicon tetrahedron:
O
connection with other elements:
4-
O
SiO4
Si
O
O
O
O
O
Mg
O
Si
O
Si
O
O
depleted oxygen
polymerization:
O
O
O
O
O
Si
O
Si
O
connection
with
common
oxygen
O
O
The degree of polymerisation (that is the degree to which the oxygen
anions are shared between tetrahedra) provides the basic structural
classification of the silicate minerals (table 2).
17
Table 2 Silicate Classification
Silicate Class
Orthosilicates
Disilicates
Ring silicates
Chain silicates
Single chain
Double chain
Sheet silicates
Framework
silicates
Number of O2Shared per
Tetrahedron
0
1
2
Si:O
Ratio
1:4
2:7
1:3
Isolated tetrahedra
Double tetrahedra
Rings of tetrahedra
2
2 or 3
1:3
4:11
3
4
2:5
1:2
Chains of tetrahedra
Double chains of
tetrahedra
Sheets of tetrahedra
Framework of tetrahedra
Structural Configuration
The orthosilicates (nesosilicates) show no polymerisation. The
silicon tetrahedra form isolated structural entities, and their negative charge
is balanced by bonding with other cations such as Mg2+, Fe2+, and Ca2+
(figure 5a). The most common igneous rock-forming orthosilicate mineral is
olivine ([Fe,Mg]2SiO4).
The disilicates (sorosilicates) show the least degree of
polymerisation in which a single O2- anion is shared between two silicon
tertahedra (figure 5b).
In the ring silicates (cyclosilicates), each tetrahedron shares two O2anions, and they form rings with three, four or six members. Ring silicates
of six members are the most frequent (figure 5c).
The chain silicates (inosilicates) may be divided into two subgroups:
in the single chain group two O2- per tetrahedra are shared; in the double
chain group half of the tetrahedra share two O2-, the other half of them share
three (figure 5d and e). The most common igneous rock-forming chain
silicate groups of minerals are pyroxenes (single chain) and amphiboles
(double chain).
In the sheet silicates (phyllosilicates), three O2- per tetrahedra are
shared to form continuous sheets (figure 5f). The most important igneous
rock-forming sheet silicate minerals are biotite (K[Mg,Fe]3[AlSi3O10][OH]2)
and muscovite (KAl2[AlSi3O10][OH]2).
The framework silicates (tectosilicates) show the highest degree of
polymerisation: all O2- anions of each tetrahedron are shared to form a threedimensional framework (figure 5g).
18
Figure 5 Classification of silicate structures
a
b
SiO4
c
Si2 O7
Ortho- (neso-) silicates
Si6 O18
Di- (soro-) silicates
Group (soro) silicates
Si4 O14
Si4 O12
Si4 O16
d
e
Explanation:
oxygen
Si4 O11
oxygen with silicon below
two oxygens and two silicons
at the same axe
Si2 O6
silicon-oxygen tetrahedron
Chain (ino) silicates
Si4 O12
elementary cell
Double chain silicates
Si4 O11
f
g
SiO2
Si2 O5
Sheet (phyllo) silicates
Si4 O10
Framework (tecto) silicates
Si4 O8
19
However, except for quartz (SiO2) and its polymorphs, Al may
occupy some tetrahedral centres in the structure. These tetrahedra have a net
negative charge, therefore, cations such as Na, K, and Ca are required to
balance charges. The most common igneous rock-forming framework
silicate minerals are quartz and its polymorphs (SiO2), potassium feldspars
(KAlSi3O8) and plagioclase feldspars (continuous solid solution series of
albite [NaAlSi3O8] and anorthite [CaAl2Si2O8]) (figure 6).
Figure 6 Geochemical system of feldspars
Sa
nid
ine
Alk
ali
fel
ds
pa
rs
Potassium feldspar
Orthoclase
K[AlSi3O8]
Anorthoclase
Albite
Albite
Na[AlSi3O8 ]
Oligoclase
Andesine
Labradorite
Plagioclase feldspars
Bytownite
Anorthite
Anorthite
Ca[Al2Si2O8]
The process by which a single homogeneous magma is able to
produce a variety of different igneous rocks is called differentiation. The
most important process of differentiation is fractional crystallisation.
During this process the crystals are completely or partially prevented from
reacting with the melt. The mechanism of magmatic differentiation by
crystal fractionation is summarised in Bowen’s reaction series. The
American petrologist N.L. Bowen was the first who, as a result of laboratory
experiments and field studies, emphasised that there was a single parent
magma of basaltic composition from which all the magma types evolved.
Although it is known that all igneous rocks are not derived by differentiation
from a basaltic magma, the processes suggested by Bowen are of great
20
significance in relating the crystallisation order of rock-forming silicate
minerals and the origin of igneous rock types.
Bowen’s reaction series consists of two branches (figure 7). On the
right side the plagioclases form a so-called continuous reaction series, as
plagioclases may grade into each other. During crystallisation the crystals
react continuously with the melt, changing their composition toward the
albite (NaAlSi3O8) end member. Minerals of the left branch are
compositionally and structurally distinct and form a so-called discontinuous
reaction series. Reaction between crystals and melt occurs only during
certain portions of the cooling sequence. There cannot be a solid solution
between olivine and pyroxene, pyroxene and amphibole, amphibole and
biotite.
Minerals at the upper part of both branches of the series, such as
olivine and calcic plagioclase, crystallise at high temperature and are
characteristic of basalts. Minerals at the lower part of the series, such as
pyroxene, amphibole, calcic-alkalic or alkalic-calcic plagioclase, crystallise
at lower temperature and are characteristic of intermediate (andesitic) rocks.
Minerals at the bottom of the series, such as biotite, alkalic plagioclase,
potassium feldspars, quartz, crystallise at the lowest temperature and are
characteristic of granitoids.
Figure 7 Bowen’s reaction series
Olivine
Pyroxenes
decreasing crystallization
temperature
increasing Si content
Pyroxenes
Anorhite-Bytownite
Labradorite
Andesine
Amphiboles
Oligoclase
Biotite
Albite
Quartz
(Zeolites)
21
2.4 Principal textures of igneous rocks
Texture is the geometrical aspect of the component particles of a
rock, including size, shape and arrangement. Besides the chemical
composition the cooling rate of magma is the principle factor that
determines the texture of the igneous rocks.
Textures of igneous rocks are basically controlled by the following
factors:
- the chemical composition of the magma;
- the cooling rate of the magma;
- the mechanism of magma emplacement;
- the thermal differences between the magma and its surroundings
(upper mantle, crust, open air, water);
- the nature and relative proportions of the volatile constituents.
These variables may act simultaneously during crystallisation and
important variations in texture may exist within the same magmatic body.
Perhaps the most important factor of crystallisation is the cooling rate. Slow
cooling of a melt may provide sufficient time for complete crystallisation.
However, when cooling rate is extremely fast the melt may solidify without
crystallisation and this results in the formation of glass.
Differences in the cooling rate of the various portions of the magma
produce differences in grain size, too. Changes in cooling rate during
crystallisation produce a rock having crystals of different size. Magma starts
to cool slowly at greater depths and then it is extruded (or intruded at a
higher level). There it cools more rapidly and the remaining liquid solidifies
as a fine-grained aggregate or as glass. The larger crystals are called
phenocrysts and the rock is called porphyritic (figure 8).
Figure 8 Some textures of igneous rocks
22
The first-formed crystals are often able to develop well-formed
crystal faces. Such crystals are described as euhedral. As crystallisation
proceeds, the crystals may have only a partial development of faces (called
subhedral) or no faces (called anhedral).
2.4.1 Principal textures of volcanic rocks
Volcanic textures that consist of more than 90 % of amorphous or recrystallised glass:
- vitreous texture has essentially no crystalline structure, it can be
regarded as amorphous material; sometimes it may contain
concentrated cracks (so called perlitic structure);
- vitro-clastic texture contains X or L shaped glass splinters and
often has innumerable cavities.
Textures of volcanic rocks having less than 90 % of amorphous or recrystallised glass:
- microlitic texture composed of phenocrysts and microlites (crystal
laths of 10-20 µm in lengths) included in an amorphous
groundmass; microlites may show flowing pattern;
- trachytic texture is characterised by less than 10 % of glass in
which oriented microlites or feldspar phenocrysts can be found;
- felsitic texture composed of phenocrysts and microlites included in
a groundmass of fine grains of re-crystallised glass;
- spherulitic texture may occur in microlitic rocks; it contains
spherules (radiating fibres) which may form due to the recrystallisation of glass or to the rapid late-magmatic crystallisation
of acicular microcrystals;
2.4.2 Principal textures of hypabyssal and plutonic rocks
- Granophyric texture is characterised by fine-grained intergrowth of
quartz and alkali feldspar.
- Microgranulitic texture is composed of phenocrysts included in
granular groundmass.
- Microholocrystalline texture consists of fine-grained crytallised
mineral constituents of approximately the same size.
- Granitic or granular texture refers to coarse- and medium-grained,
granular rocks in which nearly all of the mineral constituents are
anhedral (xenomorphic) or subhedral and of approximately the
same size.
23
2.5 Classification of igneous rocks
Three principal discriminant features are used in the classification of
igneous rocks: grain-size characteristics, modal parameters and chemical
composition. The basic classification systems were established by the
International Union of Geological Sciences (IUGS) Subcommission on the
Systematics of Igneous Rocks.
On the basis of the grain-size of the common igneous rocks, the
IUGS classification distinguishes phaneritic and aphanitic rocks. Grains in
phaneritic rocks are sufficiently coarse to be individually distinguishable
and these rocks are classified as plutonic. Grains in aphanitic rocks are too
small to be individually distinguishable and these rocks are classified as
volcanic. Within these categories the rocks are named on the basis of modal
(mineral) and/or chemical composition.
To classify igneous rocks based on modal composition the following
parameters can be determined:
Q = quartz, trydimite, cristobalite;
A = alkali feldspars, including orthoclase, microcline, perthite,
anorthoclase, sanidine and sodic albite (anorthite is less than 5
%);
P = plagioclase (anorthite is more than 5%);
F = feldspathoids (foids) including nepheline, leucite, sodalite, etc;
M = all other minerals (mainly mafic minerals) such as micas,
amphiboles, pyroxenes, olivine, opaque minerals, accessory
minerals, primary carbonates, etc.
The sum of Q, A, P, F and M must be 100 %, however, minerals
belonging to Q and F are mutually exclusive. The common rocks are best
shown in a triangular arrangement. The triangle permits classification of
igneous rocks containing mafic constituents less than 90 %. For each rock,
QAP or APF must be recalculated so that their sum is 100 %. For example, a
rock with Q = 10%, A = 30%, P = 40% and M = 20% would give
recalculated values as follows:
Q = 100 x 10/80 = 12.5%
A = 100 x 30/80 = 37.5%
P = 100 x 40/80 = 50.0%
These values can be plotted onto the QAP diagram. The plutonic and
the volcanic rocks are named on the basis of triangular arrangements similar
to each other (figure 9a and 9b). However, distinction between basalt and
andesite is made mainly on the basis of silica content: a rock containing
more than 52 % SiO2 is andesite, and a rock with less than 52 % SiO2
content is basalt.
24
A
10
lkali
feldspar
20
60
foidgabbroic-rock
Feldspatoid
peridotite
Mafic>90
foid-dioritic-rock
foidolitic
-rock
foid-syenitic
-rock
anorthositic-rock
gabbroic-rock
dioritic-rock
65
granitic-rock
quartz-richcoarse-grained
-rock
90
plutonic rocks
Quartz
syenitic-rock
60
a
feldspar
Plagioclase
A
lkali
feldspar
10
20
60
90
65
Feldspatoid
ultramafite
Mafic>90
basalticrock
tephritic-rock
foiditicrock
phonoliticrock
trachytic-rock
andesiticrock
daciticrock
volcanic rocks
Quartz
rhyolitic-rock
60
b
feldspar
Plagioclase
Figure 9 Classification of plutonic and volcanic rocks according to their
mineral content (Streckeisen 1976, 1979)
25
Rocks with more than 90 % mafic minerals are regarded as
ultramafic rocks and these rocks are classified according to their mafic
mineral content. Two diagrams can be used for classification, one for rocks
that contain pyroxene and olivine (figure 10a), and another for rocks that
contain pyroxene, olivine and hornblende (figure 10b).
Figure 10 Classification of igneous rocks according to their mafic mineral
content (IUGS Subcommission on the Systematics of Igneous Rocks)
Olivine
a
dunite
e
ha
lherzolite
ine
rth
op
yro
xe
nit
oliv
40
ite
erl
wh
rzb
urg
ite
90
no
cly
10
oliv
in
eo
ite
en
rox
py
olivine websterite
orthopyroxenite
clynopyroxenite
websterite
Orthopyroxene
Clynopyroxene
Olivine
b
dunite
te
ot i
rid
pyroxenehornblende
peridotite
pe
ep
de
len
py
rox
en
rnb
ho
eri
do
tite
90
ine
oliv
hornblende pyroxenite
olivinepyroxene
hornblendite
pyroxene hornblendite
ite
nd
26
olivinehornblende
pyroxenite
le
rnb
ho
Pyroxene
ine
10
pyroxenite
oliv
py
rox
en
ite
40
hornblendite
Hornblende
As igneous rocks crystallise from melts, the magma is becoming
oversaturated with a particular mineral. This concept is mainly used for
silica. Thus a rock may be described as silica-oversaturated if it contains
quartz (or an equivalent SiO2), or as silica-undersaturated if it contains
silica deficient minerals that cannot exist in the presence of quartz
(feldspathoids, olivine). Another approach comes from the idea that igneous
rocks are the crystallisation products of silicic acids. Thus a rock can be
called an acid rock (with silica content higher than 63%), an intermediate
rock (with silica content between 52 and 63%), a basic rock (with silica
content between 45 and 52%), or an ultrabasic rock (with silica content
lower than 45%).
The chemical classification of igneous rocks is based on the socalled TAS (total alkali and silica) diagram recommended by the IUGS
(figure 11). This classification requires only the values of Na2O + K2O and
SiO2. Analyses of rocks that weathered, altered, metasomatised,
metamorphosed should be used with caution. It can be regarded as a general
rule that only analyses with H2O+ less than 2% and CO2 less than 0.5%
should be used. Before using the TAS classification all analyses must be
recalculated to 100% on an H2O- and CO2-free basis.
Figure 11 Total Alkali Silica (TAS) diagram (Le Bas et al 1986)
16
Phonolite
14
12
Tephriphonolite
Trachyte
Phonotephrite
s
8
de
an
hy
c
a
Tr
Fo
idi
te
Na2O + K2O w%
10
Rhyolite
Tephrite
6
Trachybasalt
Dacite
4
2
0
e
sit
Basalt
Picrobasalt
37
41
45
49
Basaltic
Andesite
53
Andesite
57
61
65
69
73
77
SiO2 w%
27
2.6 Classification of volcaniclastic rocks
The general term ‘volcaniclastic’ was introduced in the early 1960s
and redefined in the early 1990s. This term includes the entire spectrum of
clastic materials composed in part or entirely of volcanic fragments and
formed by pyroclastic, hydroclastic, epiclastic, autoclastic, etc. mechanisms.
It is suggested that a rock classified as volcaniclastic must have more than
10 % of volcanic debris.
There are two basic types of volcaniclastic fragments: pyroclastic
fragments and epiclastic fragments. Fragments that are formed by volcanic
activity and perhaps reworked by sedimentary processes can be considered
as ‘pyroclastic fragments’. Fragments whose origin, as fragments, is a direct
result of sedimentary processes should be considered as ‘epiclastic
fragments’. Rocks that contain less than 25 % pyroclastic fragments are
called volcaniclastic sedimentary rocks and sediments; rocks containing
pyroclastic fragments of 25 to 75 % are named as tuffites; rocks with more
than 75 % of pyroclastic fragments are pyroclastic rocks and sediments.
The classification of pyroclasts includes three basic variables: (1)
their mode of origin, (2) their size, and (3) the composition of the
component fragments. Regarding the mode of origin the following types of
pyroclastic fragments can be distinguished:
- pyroclasts (crystals, crystal fragments, glass and rock fragments)
are formed directly from magma;
- hydroclasts are chilled particles formed by magma-water
interactions during subaqueous or subglacial extrusion;
- autoclasts are formed by mechanical friction of moving lava flow;
- cognate fragments are formed during earlier volcanic activity, but
are ejected with other pyroclastic debris;
- accidental fragments are generally derived from the subvolcanic
basement.
Considering their size pyroclastic fragments are distinguished in the
following categories (table 3):
- bombs: generally rounded fragments with a mean diameter of more
than 64 mm;
- blocks: angular or subangular fragments with a mean diameter of
more than 64 mm;
- lapilli: fragments of any shape with a mean diameter of 2 to 64
mm;
- ash: fragments with a mean diameter of less than 2 mm.
The composition of the component fragments can be characterised
by their mineralogical and/or chemical composition. In this way terms such
as basaltic, andesitic, rhyolitic, etc. agglomerate (or tuff) can be used.
28
Epiclasts are fragments (crystals, crystal fragments, glass and rock
fragments) in volcaniclastic rocks that are liberated from consolidated
volcanic or non-volcanic rock by weathering or erosion and transported by
gravity, air, water or ice.
Volcaniclastic rocks and sediments which consist of pyroclastic
fragments ranging from 25 to 75% are called tuffite. The prefix ‘tuffaceous’
should be used with the standard root names for sediments and sedimentary
rocks, e.g. tuffaceous-sand, tuffaceous-mud, tuffaceous-sandstone, etc.
Table 3 Classification and nomenclature of pyroclastic fragments and wellsorted pyroclastic sediments and rocks (after Schmidt, 1981)
Fragment
size
in mm
64
2
Dominant
pyroclastic fragment
Pyroclastic
sediments
Pyroclastic rocks
bomb, block
bomb-tephra
block-tephra
agglomerate
pyroclastic-breccia
lapillus
lapilli-tephra
lapillistone
coarse ash grain
coarse ash
coarse tuff
fine ash grain
fine ash
fine tuff
0.032
2.7 Short description of the main igneous rock types
2.7.1 Plutonic rocks (figure 12)
Ultramafic rocks. Ultrabasic igneous rocks composed of mafic
mineral constituents. Dunites have olivine as the essential mineral.
Peridotites have olivine as the dominant constituent with some other mafic
minerals (mainly pyroxene and hornblende). Pyroxenites and hornblendites
are composed essentially of pyroxene and hornblende, respectively.
Gabbros. Coarse or medium grained basic igneous rocks composed
of Ca-rich plagioclase feldspar and mafic constituents (pyroxene, olivine,
hornblende, biotite) of a considerable percentage. If the Ca-rich plagioclase
content exceeds 90 % the rock is called anorthosite. The texture of gabbros
is usually subhedral-granular.
Diorites. Medium to coarse grained intermediate igneous rocks
composed of Ca-bearing plagioclase of intermediate composition. Diorites
have a considerable percentage of mafic mineral constituents mainly biotite,
hornblende and pyroxenes. The typical texture of syenites is subhedralgranular. The hypabyssal equivalents of diorites are called microdiorites.
29
Syenitic rocks. Medium to coarse grained intermediate igneous
rocks composed of alkali feldspars with or without subordinate Ca-bearing
plagioclase and with some mafic constituents (usually biotite and
hornblende). Peralkali syenite has only alkali feldspars, and syenite has both
alkali feldspars and intermediate Ca-bearing plagioclase. In most cases the
texture of syenites is subhedral-granular.
Granitic rocks. Holocrystalline, medium to coarse grained silicic
igneous rocks that are composed essentially of quartz and one or more
feldspars, with micas (biotite and muscovite) and amphibole (hornblende).
Potassium feldspars are usually orthoclase or microcline; plagioclase is rich
in albite. These minerals are usually anhedral. Besides quartz and some
mafic constituents, alkali granites consist of alkali feldspars only. Granites
have both alkali feldspars and calcium-bearing plagioclase. In granodiorites
Ca-bearing plagioclase is the dominant mineral constituent, the alkali
feldspars are subordinate. The hypabyssal equivalents of granits are termed
microgranites.
2.7.2 Volcanic rocks
Basalts. Basalts are the volcanic equivalents of gabbros. Basalts are
usually very fine grained. The characteristic minerals are Ca-rich
plagioclase, pyroxene, iron ore and olivine. Tholeiitic basalt contains
pyroxene as the main mafic constituent, olivine is absent or present only in a
small amount. However, alkali basalt has olivine as an essential constituent.
Calc-alkali basalt has Ca-rich plagioclase as the most prominent mineral
constituent. The hypabyssal equivalent of basalt is termed dolerite which is
coarser grained than basalt.
Andesites. Andesites are volcanic equivalents of diorites. The main
mineral constituent is plagioclase feldspar that is found as phenocrysts as
well as laths in the groundmass. Composition of plagioclase feldspar may
show considerable variation. The characteristic mafic minerals of andesites
are pyroxene, amphibole, and biotite. The more basic basaltic andesites may
have olivine, too.
Trachytes. Trachytes are volcanic equivalents of syenites. The most
important phenocryst is sanidine. The other common phenocryst is biotite.
Hornblende is less frequent. Trachytes often have a so-called trachytic
texture characterised by feldspar laths oriented more or less parallel to each
other.
Silicic volcanic rocks. These very fine grained or glassy igneous
rocks are the volcanic equivalents of granitic rocks. Rhyolites have a
groundmass that is more or less crystalline and composed of feldspar and
silica minerals (quartz, tridymite, crystobalite). The most common
phenocryst is feldspar (mainly sanidine). Phenocrystal quartz may also be
30
present. The mafic minerals are subordinate, the most important is biotite;
hornblende is rarer. If the cooling rate of silicic magma is very rapid it will
solidify as glass, with or without some phenocrysts, and the rock is termed a
rhyolite-obsidian. Rhyolite-pitchstones are also glassy silicic volcanic rocks
but they have a considerable proportion of water (up to a maximum of 10
percent). Dacites are volcanic equivalents of granodiorites. Consequently,
their dominant mineral constituent is albite-rich plagioclase feldspar; quartz
and sanidine are present in a lower proportion. The main mafic constituents
are biotite and amphibole and pyroxene may also occur.
Selected titles for further reading
BARTH T. F. W.: Theoretical Petrology. John Wiley and Sons Inc. New York
1962.
EHLERS, E. G. & BLATT, H.: Petrology. Freeman and Co., San Francisco,
1980.
Dictionary of Geological Terms. Anchor Press, New York, 1975.
HAMBLIN, W.K.: The Earth Dynamic Systems. Burgess Publishing,
Minneapolis, 1985.
MACKENZIE, W.S., DONALDSON, C.H., GUILFORD, C.: Atlas of igneous rocks
and their textures. Longman, 1984.
NESSE, W. D.: Introduction to Mineralogy. Oxford University Press, New
York–Oxford, 2000.
NOCKOLDS, S.R., KNOX, R.W.O’B., CHINNER, G.A.: Petrology for students.
Cambridge University Press, 1978.
31
3. PLATE TECTONICS
3.1 The concept of plate tectonics
Scientists with wide knowledge suggested the idea that the
continents were once connected to each other way back in the 17th century.
However, the concept of Continental Drift was first published and
introduced into geology by Alfred Wegener, an Austrian meteorologist in
the year of 1912. The theory of the way continents could “drift” relative to
each other was proposed by Arthur Holmes, a Scottish geologist in the
1930s suggesting that convectional currents in the Earth’s mantle “drift” the
oceanic crust. Two American geologists, namely Harry H. Hess and Robert
Dietz, who proposed the idea of sea floor spreading in 1960 and 1961
respectively, took the third and final step towards a global plate tectonic
concept. Finally the concept of plate tectonics was accepted by the scientific
public at a conference on “The history of the Earth’s crust” held in New
York in 1966 and the term plate tectonics was first used by W. J. Morgan in
1968.
The concept of plate tectonics means that first the Earth’s lithosphere
is divided into 7 major and several smaller lithospheric plates and then these
lithospheric plates are in constant motion relative to each other. This motion
can be imagined as a rotation around a common rotational axis that is the
rotational axis of the Earth itself. Figure 12 outlines the 7 major lithospheric
plates namely the Eurasian, African, Indo-Australian, Pacific, North
American, South American and Antarctic and also some of the smaller
plates (Philippine, Fiji, Cocos, Nazca and Caribbean, Arabic).
The crust and the upper, solid part of the mantle form the so-called
litosphere, which is 50-100 km thick under continents and reaches 35 km in
thickness under oceans. The crust of the litospheric plates may be basic and
denser (3 g/cm3) oceanic crust, which is usually thin (6-10 km) or may
consist of a lower and denser basic and an acid (granitic) less dense (2,7
g/cm3) upper part called as continental crust. The latter migth be 30-35 km
thick. Some plates have no continental crust (Pacific) and consist of only
oceanic crust. The shelf region, which is covered by water is the part of the
granitic continental crust.
The heat differences in the mantle induce energy and material flow
producing convectional currents under the lithosphere. The lithospheric
plates float in the semi-molten material of the mantle driven by the energy
of the convectional currents. These currents build up so called convectional
cells in the mantle. The material is moved upward as it is heated in the
ascending part of the cell and subsides when cooled in the descending part
of the cell.
32
Pacific
Nazca
Cocos
South American
Caribbean
North American
Antarctic
African
Arabian
Eurasian
Indo-Australian
Philippine
Fiji
Pacific
Figure 12 Major and small plates of the Earth’s lithosphere
33
The geologically most active zones of the lithospheric plates
especially from the aspect of the magmatic processes are the plate margins,
which – due to the movements induced by the convection currents – can be
grouped as follows:
1. Diverging margins
2. Converging margins
3. Translational margins
3.2 Diverging margins: rifts, mid oceanic ridges and oceanic basins
3.2.1 Formation of the rift systems
Above the ascending part of the convection cells heat energy is
transported and accumulated under the continental crust. This heat slowly
burns through the crust and results in the formation of a trench called rift
valley.
As the rift widens the newly separated lithospheric plates are moved
away from each other by the convection currents. In this dilatational stress
field numerous tension joints develop parallel with the main rift valley and
create a step-like appearance on both sides of the trench (figure 13).
3.2.2 Magmatic processes
The heat flux is measured to be very high along the rift systems
suggesting that hot mantle material is rising into the space developed
between the diverging plates. This ultrabasic, so called tholeiitic mantle
material arriving from a depth of around 35 km gives the source of the
magma that accumulates in several deeper magma chambers, where it
undergoes relatively high pressure fractional differentiation. In the course
of this differentiation a series of ultra basic and basic rocks is produced that
is called the ophiolite series.
At the bottom of the ophiolite series i.e. in the deeper parts of the rift
the least differentiated magma material crystallises into peridotites with
very high ultra basic mineral and metal content. Then as the SiO2 content of
the magma increases and basic silicate minerals like olivine, clinopyroxene,
orthopyroxene and basic plagioclases are produced gabbros lherzolites
and harzburgites are formed. Then the magma is trapped in shallow magma
chambers at a depth of around 2-5 km where it undergoes low-pressure
crystal fractionation resulting in the formation of olivine phyric rocks.
Other magma chambers receive periodic infusions of magma from
greater depths, or primitive magma directly from the mantle resulting in the
mixing of magmas. The gabbros are overlain by dolerites. Finally basalts
are produced representing the top of the ophiolite series. The basalt lava
entering the surface due to its low viscosity flows all over the surroundings
34
of the ridges to produce floodbasalts. However, the drop-like solidification
of the basalt rocks producing pillow basalts (see figure 3) is also very
frequent. Basalt lava may flow into and solidify in the tension joints filling
them producing and basalt dykes. Therefore as oceanic crust is formed
along the rift valleys it is built up primarily of intrusive gabbros and
extrusive basalts.
Occasionally, the magma intruding into the tension joints may be
trapped in small pockets and undergoes extreme fractionation and SiO2
accumulation resulting in the production of granite rocks (A type granite).
Figure 13 Structure and magmatism of a rift system
Pillow lava+
veins
Sheeted dykes
Isotropic gabbro
magma
Gabbro cummulate
mafic cum
mulate
ultramafic cumm
ulate
harzburgite, restite
melt+residue
35
3.2.3 Structural and global tectonic consequences
As magma continuously ascends and oceanic crust is formed, the rift
valleys are in a restless widening process of becoming mid oceanic ridges.
So all mid oceanic ridges are former rift valleys representing the border of
diverging lithospheric plates today. The complete length of the mid oceanic
ridges on the Earth is 75000 km. These areas are very active regarding both
volcanism and shallow seated earthquakes (hypocentre is less than 30 km
deep) as oceanic crust is continuously being formed here.
Pillow basalts may form several hundred metres high and 500-1000
m wide basalt ridges. This is one of the reasons why mid oceanic ridges rise
2000-3000 m above the ocean floor. The other reason is thermal expansion.
It is relatively rare, however, that mid oceanic ridges reach the surface of the
ocean as in the case of Iceland.
The explanation for the extreme uplift of the northern part of the Mid
Atlantic Ridge is found in global tectonics. As was mentioned before, mid
oceanic ridges separate two diverging lithospheric plates but in the case of
Iceland, global tectonic trends are reversed and Eurasia is moving to the
Northwest as Africa is pushing it, while North America remains static as the
Pacific plate is not consumed on its eastern side. Thus the northern part of
the Mid Atlantic Ridge is under compressional pressure that uplifts it above
the surface of the Atlantic.
The volcanism along mid oceanic ridges is limited to the central
trench and to the closest joints. The basic igneous rocks are “stuck” to the
edge of the diverging plates. This process is called accretion. The rate of the
accretion seems to be about the same as the velocity of divergence. This
divergence, accretion and the resulting increase of the oceanic lithosphere
are the process called sea floor spreading.
The primary evidence of sea floor spreading was found by
palaeomagnetic investigations. As oceanic crust is formed along mid
oceanic ridges the magnetic minerals turn towards the magnetic poles of the
Earth as they are oriented at the time of the crystallisation of the rocks.
Based on this, belts of rocks with different magnetic orientations can be
observed parallel with the main trench of the mid oceanic ridges on their
both sides. The magnetic orientation can be normal when oriented as it is
today and reverse when oriented the opposite way. Scientists determined
the age of every polarity change detected in the stripes of the oceanic basin.
Therefore the rate of sea floor spreading and thus the rate of divergence
could be reconstructed.
It was discovered that the velocity of divergence has been variable
both in space and time. However, its average rate is between 2 and 20
cm/year. As the divergence of two lithospheric plates can be regarded as
36
rotation with opposite directions the break of the mid oceanic ridges is
inevitable. This break occurs as transform faults along which significant
tectonic activity can be observed. Based on palaeomagnetic, biostratigraphic
and radiometric age data the age of the oldest oceanic crust is found to be
not older than 170 million years.
3.3 Converging margins: subduction of the oceanic crust; island arc
systems
3.3.1 Development of the converging margins
The reason why oceanic crust older than 170 million years cannot be
found on the Earth is that the oceanic crust is consumed at several places
namely where the lithospheric plates are converging. As a lithospheric plate
is normally composed of a continental and an oceanic crust, usually it is the
oceanic crust that is broken when two such plates converge. The exact place
where the oceanic crust breaks depends on the blocking effect of the
continental crust. The greater the continental crust is, the further away from
the continent the oceanic crust will be broken.
Once the oceanic crust is broken the denser will be pushed
underneath the other one into the upper part of the mantle. This part of the
oceanic crust is called the descending slab. As the ascending slab is pushed
deeper and deeper into the mantle, its initial heat and gravity anomalies
vanish and it takes up the physical properties of the mantle. This
consummation of the ascending slab is called subduction. The plane along
which subduction takes place is called the Benioff zone. This zone holds the
vast majority of the hypocentres of the earthquakes produced by the
subduction (figure 14).
The angle at which the subduction mainly happens depends on the
speed of the subduction i.e. the relative velocity of the converging plates.
The angle of the Benioff zone is low when the velocity of the subduction is
high. This is the case when the converging plates move in opposite
directions. And the angle of the subduction is high when the plates move in
the same direction but one of them moves faster so they are still considered
as converging plates. However, usually the angle of the Benioff zone is
detected to be around 45°.
37
Deep sea trench
Accretionary prism
Be
ni
of
fz
on
e
Su
b
du
ct
in
g
sla
b
Forearc basin
Island arc
Backarc basin
Continent
Figure 14 Structure and magmatism of an island arc system
3.3.2 Magmatic processes
The maximum depth to which the subducting slab is pushed is
observed to be 700 km. Beyond this depth the subducting oceanic crust
cannot be detected.
There are several reasons for the subducting oceanic crust to become
molten:
¾ the Astenosphere is much hotter than the oceanic crust
¾ frictional heat is produced while the oceanic crust descends
¾ the altered basalt in the top layer of the oceanic crust contains a
significant amount of water that decreases the melting point of the
minerals that make up the subducting crust
The molten material of the subducted crust gives the source of the
magma material formed at the place of consummation. The magma is
accumulated in magma chambers where it undergoes fractional
differentiation for the second time in its tectonic life. However, in this
case the parent material of the magma is not ultra basic mantle material but
the basic rocks of the oceanic crust. Therefore differentiation will produce
melts dominated by neutral silicate minerals such as amphiboles and neutral
plagioclases (labradorite and andesine) that together with pyroxenes and
basic plagioclases make up neutral intrusive and extrusive igneous rocks as
diorite and andesite respectively.
38
The volcanoes produced by this volcanism build up alongside the
edge of the continents in a curved shape forming the island arcs when
reaching the surface of the ocean.
The island arc system is composed of three parts, namely:
- Forearc region
- Island arc
- Backarc region
3.3.3 Forearc Region
The forearc region is situated between the deep sea trench and the
island arc. The deep sea trench is usually an elongated, slightly curved, very
narrow asymmetric depression that represents the deepest parts of the
Earth's crust. As the subducting slab descends the majority of the pelagic
deep sea sediments on the top is scraped off and thrust under the edge of the
opposing lithospheric plate. Due to this underthrusting the sediments will
form strongly deformed imbricated nappe systems that are stuck to the
edge of the other plate. This process is called accretion and the resulting
strongly deformed fan like wedge of deep sea sediments is called the
accretionary prism. Due to the underthrusting the vergence of the
imbricated structure of the accretionary prism points towards the island arc.
The oldest sediments are found in the top of the accretionary prism and their
angle of dip is the greatest.
The forearc basin is found between the accretionary prism and the
island arc. This is normally characterised by coarse unmatured terrigenous
sediments eroded from the volcanic arc. The form of the basin is greatly
dependant on the sedimentation rate and the deepening of the basin. Fast
sediment accumulation may fill up the basin in which case it forms a wide
shelf.
The rocks that build up the forearc basin are situated in belts. High
pressure low temperature blueschist facies metamorphic rocks and
melange characterise the oceanward side while metamorphosed volcanics
and batholits are found on the continentward edge.
3.3.4 Island arc
The island arc is always situated right above the Benioff zone on the
continentward side of the deep sea trench to which it is parallel. The line of
island arcs may be several thousand km long but their width is around only
200 km while the particular belt that is characterised by active volcanism
may be only 50 km wide. The position of the igneous arc is determined by
the outline of the Benioff zone. Considering a subducting slab that descends
along a plane that has an angle of 45° the igneous arc may be as far as 200
km from the deep sea trench.
39
The island arc is built up mainly of the neutral igneous rocks that
are produced by the differentiation taking place under the arc. However,
acid and basic igneous rocks can also be found in the arcs. The amount of
acid and basic igneous rocks is determined by the material of the crust upon
which the arc is developed. As island arcs may form on either continental or
oceanic crusts that contaminate the ascending magma, acid-neutral rocks
can be formed like dacites and granodiorites. Usually the island arcs are
characterised by strato volcanoes that are built up by subsequent strata of
lava rocks and pyroclasts.
3.3.5 Back-arc Region
There are different back-arc regions on the different sides of the
Pacific plate. On the western side of South America the back-arc region
consists of former basins that formed in the course of formerly ceased
subductions. In contrast, one of the most characteristic features of the
eastern Pacific is the complex system of back-arc or marginal basins. The
development of back-arc basins is associated with local rifting. The cause of
this local rifting is that the subducting slab induces local disturbances in the
asthenosphere under the back-arc basin therefore local convection currents
will develop. Thus back-arc basins normally have oceanic crust.
3.4 Collision, orogene regions
3.4.1 Development of orogene regions
When the converging movement of two lithospheric plates lasts for a
long time and the oceanic crust between the two continental crusts is nearly
completely consumed there can be several parallel island arc systems
developed. And when the oceanic crusts are consumed between the island
arcs these arcs will be compiled onto each other due to the compressional
stress, to form highly deformed nappe structures with a vergence pointing
opposite to the direction of the principal stress axis i.e. where the maximal
force is coming from (figure 15).
The final phase of the convergence is when two continental crusts
collide. In this case neither of them will be pushed down into the
asthenosphere due to the great difference in density. Instead the two
continents will move upwards, while the strongly broken and folded
structures will form interfingering with each other. Some parts of the
denser crust, however, will be underthrust the less dense one thus uplifting it
even higher.
The results of this process are the strongly folded mountain ridges
like the Alps, the Carpathians and the Himalayas that contain the highest
points on the Earth. Where the enclosure of the oceanic crust that is between
the colliding continents happens, some parts of the oceanic crust is not
40
subducted but obducted. These obducted basic ophiolite rocks represent the
line of the so called suture along which the two continents collide. This
collision zone between two continental crusts with strongly folded and
jointed island arc remnants, obducted ophiolites and fracturing and
elevating continental crust materials forming high mountain ridges is called
the orogene region.
3.4.2 Magmatic and metamorphic processes
When two continents collide the material of the continents and the
material of the enclosed former island arcs is re-melted at a relatively
shallow depth, 8-12 km deep in the root of the orogene due to the frictional
heat. When the rocks are only partially re-melted the process is called
anatexis and when they are completely re-melted it is called palingenesis.
This re-melting of the orogene material results in the forming of shallow
seated magma chambers.
Figure 15 Structure and magmatism of an orogene system
suture
nappe system
foreland basin molasse sediments
In these magma chambers the magma material undergoes fractional
differentiation for the third time in its tectonic life producing acid magma
material. This is also mixed with the material of the surrounding rocks
forming extra acid igneous rocks named granitoid rocks. When the magma
is the result of the re-melting of mainly igneous rocks I-type granites are
formed while the re-melting of mainly sedimentary rocks produces S-type
granites. The material of the continental crust is formed in these orogene
regions.
41
Due to the high compressional pressure acting when two continents
collide dynamo metamorphic rocks also form. And due to the re-melting of
both igneous and sedimentary rocks ortho- and parametamorphic rocks are
formed respectively (see metamorphism).
3.5 Translational plates
There are lithospheric plates that move parallel to each other. Their
movement takes place along a strike-slip fault. This movement results in
neither the development of deformed structures of folded mountain ridges
nor the formation of magma chambers and resultant volcanoes. However,
strong tectonic activity is indicated by heavy earthquakes as in the case of
the San Andreas Fault in California.
3.6 Evolutionary trends of the lithosphere
Summarising the chapter we have seen that an ultra basic mantle
material goes through a global tectonic evolution in the course of plate
tectonics involving a series of fractional differentiations. In this relation
the evolution means the increasing of silicon content of the material (figure
16).
The global tectonic evolution starts with the rift valleys where the
ultra basic material of the mantle through the first fractional
differentiation forms the basic ophiolitic series of the ocean floor which –
among the lithospheric plates – can be regarded as the less differentiated one
(“undifferentiated”).
In the island arc systems the basalt and gabbro of the subducting
oceanic plate due to the second fractional differentiation turn into neutral
andesites and diorites. So the island arc systems can be regarded as more
differentiated parts of the lithosphere if they are not independent plates.
Because of the strong selectivity of the weathering (see later at the “concept
of weathering”) sedimentary processes also promote the geochemical
changes of the material and since these changes also tend to increase the
silicon content, they can be regarded as the “sedimentary differentiation”
of the material.
In the root of the orogene regions neutral rocks mixing with
sediments and metamorphic rocks re-melt and become acid-neutral dacites
and granodiorites and acid rhyolites and especially granites (see
“ultrametamorphism”). These acid igneous and metamorphic rocks form the
acid parts of the continental lithosphere that – in this relation – can be
regarded as the end product of the global tectonic evolution.
42
Rock forming
processes
Geodynamic
Position
and
Pyroxene
(Ca-rich)
Mafic
Basalt
Gabbro
Basic
Oceanic crust
45
melting of the
mantle
Rift
Olivine
Ultramafic
Colour
Mineral
Composition
Komatiite
Peridotite
Ultrabasic
Extrusive
Intrusive
Composition
Silica % 30
fractional
differentiation
Island arc
63
Orogene
80
Craton
Muscovite
(Na-rich)
Quartz
Orthoclase
Felsic
Rhyolite
Granite
Acid
melting of the "crust"
(ultrametamorphism)
Biotite
Plagioclase
Intermediate
Andesite
Diorite
Intermediate
Amphibole
52
Figure 16 Relationship between the igneous rocks, rock forming processes
and global tectonic events
43
Selected titles for further reading
PARK, R. G.: Geological Structures and Moving Plates. Blackie USA
Chapman and Hall, New York 1988.
RAMSAY, J. G. & HUBER M. I.: The Techniques of Modern Structural
Geology I-II. Academic Press Limited, London 1989.
SUPPE, J.: Principles of Structural Geology. Prentice-Hall, Inc. Englewood
Cliffs, New Jersey 1985.
44
4 SEDIMENTARY ROCKS
4.1 Origin, classification and general description of sedimentary rocks
4.1.1 Formation of sediments and sedimentary rocks
Formation of sedimentary rocks is directly related to the superficial
processes of the earth crust. The most important stages in their development
are disintegration (decomposition), transportation, accumulation
(sedimentation) and diagenesis.
Disintegration includes several processes of the physical and
chemical decomposition.
Physical processes can be as follows
-insolation can be related to the frequent changes of the near-surface
temperature in the course of which dilating stress may develop
between the differently expanding minerals (e.g. silicate structures)
-frost shattering means that the infiltration of fluid water into the
small joints and cavities of the rocks which in the course of the
solidification (freezing) fragment the enclosing material by the
expansion
-weathering summarizes the chemical processes (hydrataion and
hydrolyses) of the superficial alteration
-biogenic disintegration means that the growing plants mainly with
their roots can crack the rocks
-salt crystallization means that the pressure of the growing crystals
can also lead to the disintegration of the rocks
Transportation is dominantly induced by superficial vertical
unevenness due to tectonic events like elevation and subsiding and
determined by gravitation. It can be characterised mainly on the basis of the
transporting agents, which can be gravitation itself (slope debris, flysch) or
air (wind) resulting in so-called wind-blown sediments, water (rivers,
marine currents) resulting in fluvial and marine sediments and ice (glaciers
and continental ice sheets) resulting in tillites.
Accumulation of the sediments is determined by the
(palaeo)environment, mainly by its physical, chemical and occasionally by
its biological effects. The most important environments of the accumulation
can be grouped as marine and terrestrial ones, with a great number of special
environments within them.
Diagenesis includes the processes of compaction, desiccation and
cementation.
45
4.1.2 Classification and general characteristics of sedimentary rocks
The classification of sedimentary rocks is based on the main forming
processes and consequently on their main geochemical characteristics. In
this way we use the terms of siliciclastic, carbonate, biogenic rocks,
evaporites and argillaceous rocks as the result of the weathering of silicate
structures.
4.1.3 The main sedimentary environments (sedimentary basins)
In the course of transportation sediments are eroded, transported and
deposited in deeper parts of the earth surface, which are therefore frequently
called sedimentary basins and frequently but not necessarily filled with
water. Based on the actual superficial conditions the sedimentary basins can
be grouped into marine and terrestrial basins.
The most important ecological environments of the marine region
are the deep oceanic basin (pelagic conditions), the continental slope
(hemipelagic conditions), the shelf and the shore zone. The main
sedimentary systems in the marine region are the siliciclastic and the
carbonate systems. The combination of the two aspects can lead to the most
important marine sedimentation systems as the siliciclastic or carbonate
shore zones, shelves, continental slopes and deep basins.
The ecological environments in the terrestrial region can be
grouped into wet and dry terrains such as rivers, lakes and marshes
contrary to soil systems. Both the wet and dry terrestrial systems are
relatively unstable places of the sediment accumulation, since due to
subsequent elevation the temporally deposited sediments will be redeposited
(reworked) towards a deeper local sub-basin of the area till they reach the
marine region. (Note that dry terrains can not be the place of the stable and
important sediment accumulation since due to erosion this area is just the
source area of reworked sediments.)
Within these ecological environments and sedimentation systems
smaller, so-called facies appear with well defined environmental factors
(water depth, sedimentation rate, current conditions, biogenic activity etc.).
Some examples from the marine region are the delta complex of a river
within a siliciclastic system, or a reef complex within a carbonate shelf zone
etc. In a given fluvial system of the dry terrain a facies may be the point bar
or the channel bar. These facies mean the elementary units of the
sedimentation where the defined environmental, especially ecological
factors directly determinate the lithological characteristics of the forming
sediments. Therefore the recognition of facies has essential importance in
the (palaeo)ecological reconstructions.
The main purpose of sedimentology is just to describe these facies
and define the relationship between the environmental factors and the
46
lithological characteristics of the forming sediment and/or sedimentary rock.
From this aspect the sedimentary rocks can be regarded as the essential tools
in the reconstruction of the geological past in the surface of the earth crust.
4.1.4 Bedding
Possibly the most common nature of the sedimentary rocks is that
they are usually bedded. One bed represents and so describes one complete
“chapter” i.e. one elementary period of the development under permanent
palaeoconditions. The changing of the environmental conditions (change of
water depth, sedimentation rate, current etc.) generates changes of the
sedimentation, initiating the development of a subsequent bed on the surface
of the previous one.
In this relation a given bed can be characterised by its underlying
and overlying associates, separated from them by its lower and upper
bedding surfaces. The morphology of the bedding surfaces reflects
essential data on the energy of the palaeoenvironment. The bedding surface
can be e.g. straight and wavy, reflecting undisturbed or disturbed conditions.
Figure 17 shows the most important types of bedding surfaces. Some
extremely various forms can be regarded as the result of compaction
simultaneous with or subsequent to the sedimentation (convolute bedding)
or in the course of or after the diagenesis (stylolitic structures).
Figure 17 Main types of bedding (after Campbell 1967 and Allen 1977)
horizontal bedding
convolute bedding
ripple lamination
irregular bedding
trough bedding
stylolitic bed surface
The next informative character of the bedding is the thickness of the
beds. The thickness is determined by the intensity of the sedimentation and
by the stability of the (palaeo)conditions. In case of extremely low
sedimentation rate (e.g. deep marine environments only a few mm per 1000
years) or in case of extremely frequent changes of sedimentation conditions
like the seasonal changing in lacustrine environment laminar bedding can
47
be produced with only a few millimetres of bed thickness. Contrary to this
in case of the shelf environments due to the relatively high rate of
sedimentation within relatively stable sedimentary conditions rather thick
layers of sedimentary rock can be formed. They are called well or thick
bedded rocks. There are some processes that can destroy the bedding like
bioturbation which refers to the mixing activity of the animals living and
feeding within the sediments (benthic living creatures).
A special form of bedding is the cross-lamination or cross-bedding.
The definition refers to beddings where the bedding surfaces are connected
with secondary surfaces called foresets. The foresets are positioned at sharp
angle to the bedding surface, their morphology can be sinuous or straight.
The formation of cross-lamination is related to the accumulation in a current
(wind or water) and the dip of foresets is one of the most important tools in
determining the existence, the strength and also the orientation of the
palaeocurrents. Presumably this is the main reason why the terminology of
cross lamination, the basis of which was established by ALLEN (1963), is
rather diverse and frequently used. Some important forms of the cross
lamination can be seen in figure 18.
Figure 18 Some examples for cross stratification (after Allen 1963)
According to the shape of the foresets
angular
tangential
sigm oidal
According to the shape of the foreset sets
tabular
wedge-shaped
trough
According to the 3D m odel of the stratification
48
4.2 Descriptive characteristics of siliciclastic rocks
The main descriptive characteristics of the siliciclastic sedimentary
rocks are grain size distribution, grain size morphology and mineral
composition. Below we show the most important aspects of determination
and interpretation of them only to illustrate how much information on rock
forming (palaeo)environments can acquired by the simple analysis of the
simultaneously formed sedimentary rocks.
4.2.1 Determination and interpretation of grain size distribution
The grain size distribution of a sediment or a sedimentary rock is
basically related to the conditions of the transportation, especially to the
viscosity and current characteristics of the transporting material. Large
blocks together with extremely small particles can be transported by ice (e.g.
huge erratic blocks were transported by continental ice sheets in North
America and North Europe in the Pleistocene) while the wind-blown
sediments can not be coarser than 0,2 mm. So the grain size distribution of
the sediment was found to be one of the most important descriptive
characteristics of the texture of the sedimentary rocks.
The terminology in sedimentology is related to the most important
categories of the particles. Based on ATTERBERG 1903, UDDEN 1914,
WENTWORTH 1922 we use special terms for these categories like gravel
(commonly above 2 mm), sand (to 0,02 or 1/8 possibly 1/16 mm), silt (to
0,002 or 1/128 or possibly 1/256 mm) and clay (under 0,002 or 1/128 or
possibly 1/256 mm).
In the course of designating a given sedimentary rock we use the
name of the dominant category (gravel, sand, silt or clay), together with the
secondary and even with tertiary fraction as indicative terms. In case of sand
with silt as subdominant fraction we use the term silty sand. The ratio
between these main categories became one of the most important descriptive
factors and determined the terminology of the sediments and sedimentary
rocks. The mixing of three adjacent fractions (gravel-sand-silt or sand-siltclay) can be regarded as a common, natural case since a perfect selection of
similar fractions requires special conditions. This is the main reason why in
sedimentology the comparative analysis of sediments is frequently made by
using ternary diagrams, some examples of which (e.g. SHEPARD 1954, FOLK
1968) can be seen in figure 19.
49
Figure 19 Ternary diagrams for indicating the ratio of sand, silt and clay
(according to Shepard 1954 and Folk 1968)
sand
y
cla
ty
sil
sa
nd
yc
lay
clay
fine
clayey silty
sand sand
silty
sand
sand
silty
clayey
sand
silty sand
sandy
clayey
silt
ilt
ys
ye
cla
cla
ye
ys
an
d
sandy
silty clay
sandy silt
sandy
clay
silt
clay
sandy
fine silt
fine silt
sandy
silt
silt
If a detailed (descriptive or comparative) analysis of the sedimentary
rock is required for example to differentiate similar but slightly different
plaeoenvironments or to give an adequate description of the texture (e.g. the
rate of the effective porosity), a detailed statistic analysis of the grain size
distribution has to be produced. The gravel and sand fractions can be
fractioned into sub-populations using sieve-series, the silt and clay ones can
be fractioned and measured by sedimentation methods based on Stock’s
law. The separated fractions can be measured by analytical pair of scales
and represented in grain size distribution diagrams.
A common property of these diagrams (figure 20) is that the number
of particles in a given fraction is expressed by weight percent. The grain size
is indicated on the abscissa in millimetres or on phi (φ) scale. (Phi is the
negative logarithmic value to base 2 of the millimetre value, so 1/2
millimetre is 1, 1/4 millimetre is 2, 1/8 millimetre is 3 while 1/16 millimetre
is 4 in phi etc.) The millimetre scale must be logarithmic while the phi scale
can be equidistal because of its logarithmic base. The percent of frequency
is indicated on the ordinate axis.
The indication of categories can be made by histograms, by
frequency and by cumulative curves (figure 20). The first represents the
real results of the analysis, the second is suitable for quick statistic
interpretation of the given distribution, while the third allows the quick and
easy determination of the grain size category for given frequency percent
values (e.g. d5, d10, d15, d50, d60, d85, d90, d95) after the interpolation of
the data. Since these values have great importance for the graphic statistic
50
analysis of the grain size distribution diagrams the latter method became the
most widely used in sedimentology in the 20th century.
Figure 20 Grain size distribution histogram and curves in mm and phi scale
1
2
3
4
5
6
7
8
9
phi
10
100
%
75
50
25
0
1
0,
5
0,
2
0,
1
0,
06
0,
02
0,
01
0,
00
5
0,
00
2
0,
00
mm
1
In a statistical interpretation of the grain size distribution curves
(figure 21) the most important parameters are modality, mean, median,
standard deviation, and skewness of the distribution.
The grain size distribution of the sediment may be uni-, bi- or
oligomodal and polymodal. The dominance of one grain size population
refers to a strongly selective (palaeo)condition which initiated the
accumulation of one favoured particle type. The existence of two grain size
populations may refer to the two different ways of sediment transport
considering same transporting material (e.g. suspended silt in the water mass
and saltated sand in the bed surface of a river) while the two and oligomodal
curves can refer to a reworking and mixing of two or more originally
different sediments. Polymodal curves refer to sediments without any
separation as in the case of the ice transported tillites.
The mean and the median refer to the dominant strength of the
transporting material regarding that its real position in the x axis indicates
the dominant grain size population in the sediment (clay, silt, fine or coarse
sand or even pebble). Nowadays the importance of the mode value is greater
than that of the mean.
51
52
modality
standard deviation
and curtosis
relationship between
median and mode
high value of the
standard deviation
and the kurtosis
low value of the
standard deviation
and the kurtosis
symmetric
unimodal
asymmetric
(with positive or negative skewness)
bimodal
polymodal
Figure 21 Some important distributions in sedimentology
The standard deviation of the distribution refers to the selective
effects and in this way primarily to the viscosity of the transporting material
and secondarily to the current conditions of the transporting processes. So
the value of the standard deviation is rather low in sand-blown sediments
and extremely high in the ice transported materials depending on the low
and high viscosity of the air and the ice. At the same time it is relatively low
in well sorted shallow marine sediments transported by strong marine
currents; while relatively high in deep marine sediments where the selective
effect of the marine currents can not be recognised.
The existence of skewness refers to the fact that the mode and the
mean of the distribution is not the same point on the x axis. From the aspect
of the sedimentary processes it means the extreme concentration of the fine
or coarse grain size population in the sediment. Based on (MASON AND FOLK
1958) this can be seen in the case of the shallow marine sandy sediments to
the advantage of the coarse part of the distribution and in the open marine
sediments to the fine portion. Other investigations pointed out that at the
background of the skewness, a concealed bimodality can be detected. So the
importance of the skewness is behind that of the modality, the mean, the
median and the standard deviation.
Based on the determination of statistic parameters of sufficient
amount of sediment samples taken from recent marine and terrestrial
environments a great series of variation diagrams and associated
sedimentary models were developed to separate the samples from different
sedimentary (palaeo)environments. Some examples (FOLK 1968, PASSEGA
1964) are shown in figure 22a and 22b. The detailed description of the
diagrams can not be given in the frame of this course but is easily traceable
in the significant work of BALOGH (1991) on sedimentology. The usefulness
of these models is unquestionable, but the utilization of the diagrams
requires great caution since the applicability of them is limited mainly to
similar environments. So the correct interpretation of sedimentary rock
should be based on simultaneous analysis of the bedding, the grain size
distribution and the undermentioned grain morphology and mineralogical
composition.
53
Figure 22 Variation diagrams for sedimentological analyses (according to
Folk 1968 (a) and Passega 1964 (b))
a
sha
llow
ma
rin
e
2
σ
stal
coa
1
silt
sand
pebble
0
-4
clay
8
4
Mz(φ)
12
b
M
(med)
01
0,
02
0,
0,
04
06 08 1
0, 0, 0,
2
0,
4
0,
6 8
0, 0,
rounded particles and suspended
material together
IX
III
4
II
I
V
pelagic suspension
IV
0,2
0,02
54
0,04
=
0,1
0,08
0,06
C
subdominantly
sorted
suspension
M
VIII
VI
2
1
0,8
sorted
suspension
VII
10
8
6
0,6
0,4
C
(> 1%)
4.2.2 Analysis and interpretation of grain morphology
The size of the particles of the siliciclastic sediments can be
investigated from two aspects: sphericity and roundness (figure 23). The
first can indicate the mechanism of transportation and the second can refer
to the scale of it.
Figure 23 Determination of morphology of clastic sediments (according to
Sloss 1963)
sphericity
0,9
0,7
0,5
0,3
0,1
0,3
0,5
0,7
0,9
roundness
The sphericity refers to the similarity of the particle to an ideal
sphere or to an ideal ellipsoid with two or even three axes. There are
published observations that the rounded pebbles of the tidal environments
have ellipsoid forms with two axes in contrast to the fluvial pebbles which
have ellipsoid forms with three axes. The reason for the differences may be
the different orientation of the currents and so the different mechanisms of
the transportation. In the first case the pebble is moved back and forth by the
tidal currents turning round the same axis. Contrary to this, in the latter case
the pebble is transported on the bed surface of the river, its longest axis is
parallel with the river axis, the second is horizontal perpendicularly with the
main current, and the shortest one is vertical. In this case the pebble is
polished by the suspended material, which maintains the original shape of
the pebble.
55
Roundness means the superficial micromorphology of the pebble. In
this case the surface of the pebble is investigated and the ratio of the
concave convex and plane parts is determined. There is a negative
correlation between the ratio of the concave parts and the way of
transportation since in the course of the transportation the ratio of the plane
and convex parts increases.
The scanning electronmicroscopic (SEM) investigation of the
surfaces of small (sand) particles may be also very useful in case of
distinguishing tidal, fluvial or even aeolian circumstances. In the case of
tidal sands well rounded, well polished particles are common contrary to the
regularly splintered particles of the fluvial sands (figure 24). In the case of
the aeolian sands the well rounded and polished particles are regular
similarly to the tidal ones, but small “impact craters” on the surface of the
polished particles initiated by frequent collisions of the saltated particles are
also characteristic.
Figure 24 SEM photographs of sand particles - beach at left, fluvial at right
and aeolian below (photos were made by Borsy)
56
4.2.3 Mineralogical composition of siliciclastic sediments
The mineralogical composition of the siliciclastic rocks is
determined by microscopic analysis in the course of the
micromineralogical investigation. The first stage of the analysis is the
separation of the light and heavy minerals and the determination of the
ratio between the two fractions. The definition is based on the density of the
rock forming minerals, the separation is made by any kind of heavy liquids
such as bromoform (CHBr3, its density is 2,9 g/cm3). The floating light
minerals can be collected on the surface of the liquid while the sinking
heavy minerals can be collected by filtering.
Light minerals have a dominant role in the formation of siliciclastic
rocks. Their ratio is regularly more than 90 percent. The most important
light minerals are quartz (the dominant), feldspars and muscovite. Some
characteristic minerals from the heavy fractions are the tourmaline, garnets,
zircon, hornblende, pyroxenes, sphene, rutile, apatite.
The dominant mineral in sandstones is the quartz because of its great
resistance to weathering (see later). If the feldspar becomes an important
mineral in the light fraction the term arkosic sandstone is used.
Considering that the feldspars cannot resist the long-term weathering these
materials can be regarded as non-matured sediments with short
transportation and fast covering. If rock fragments (connecting groups of not
less than three specimens of minerals) can be seen in the sandstone the term
litarenite is used proving also short transportation (figure 25).
Figure 25 Ternary diagrams showing mineral composition of sandstones
(according to Folk 1968 and Krynine 1948)
quartz
quartz, chert
silica sand
arkos
rock fr e with
agmen
ts
ark
os
e
feldspar
kaolin
uw
gra
rock fragment
mixed
arkose
r
pa
ds
fel
in
or
ar
sp
po
eld
ke
nf
ac
hi
uw
ric
ke
gra
ac
e
nit
are
lith
feldspar
silica
sand
sublitharenite
atic
feldsp ite
n
lithare
ark
os
e
subarkose
clay, mica,
chlorite
57
4.3 Weathering of silicate structures
4.3.1 The concept of weathering
The earth crust is dominantly built of magmatic and metamorphic
rocks, the main components of which are different groups of silicate
structures ortho, chain, double-chain, sheet and framework silicate structures
depending on the silicon content of the given material (rock or mineral
association) (See chapter 2). These silicate structures were formed under
high pressure (several thousand atmospheres) and high temperature
(regularly higher than 700 °C, which is the lower boundary of the main
crystallization phase). Therefore it is easy to realize that they may not be
thermodynamically stable phases under surface conditions the most
important physical factors of which are the extremely low pressure (one
atmosphere), low temperature (0-20 °C) and a relatively great concentration
of water. The weathering of the silicate structures means a geochemical
decomposition of these rock forming minerals and at the same time the
formation of phases i.e. silicate structures which can be stable under surface
conditions.
The stability of the silicate structures is greatly determined by the
level of polymerization. Ortho and chain or double-chain silicate structures
prove to be essentially unstable structures similarly to feldspars and
feldspatoids. In contrast the sheet silicate structures and especially the quartz
can successfully resist even the long-term weathering. (This phenomenon is
the main reason why the dominant mineral among the light minerals in
sandstones is the quartz.) The main causes of this resistance are the
relatively compact structure of the sheet and ideal framework silicate
structures and the stability of the quartz structure under low temperature and
pressure (see the late phase of the hydrothermal crystallization).
The generic stable silicate phase (i.e. structure) of the surface
environments is a special group of the sheet silicate structure, the special
physical characteristic of which is that it can build in a great amount of
water in the crystal lattice without any decomposition or even any
disorganisation. This group of the silicate structures is composed of different
forms of clay minerals, which are in this way as dominant representatives
of the surface environments and sedimentary rocks as the rock forming
minerals of the Bowen series in the formation of magmatic rocks. The clay
minerals are not only stable phases under the superficial conditions (low
pressure and temperature) but their existence and even their formation is
mainly limited to these circumstances, while under increasing pressure and
temperature they alter into mica structures (see the metamorphism).
58
In the family of clay minerals (figure 26) three- and two-layered
forms are also known. The more organized three layered sheet silicate
structure near to the mica structure is represented by the structure of the
illite. The more disordered three layered sheet silicate structure with greater
water content and at the same time with much greater flexibility is the
montmorillonite. The latter can be expanded by enclosing a significant
amount of water. The two layered sheet silicate structure is represented by
the kaolinite. The size of the clay mineral crystals is regularly lower than
0,002 mm so their investigation can be made primarily by X-ray
diffractometer (XRD).
Figure 26 Structure and classification of sheet silicate structures (after
Weaver 1989)
K+
7.1 - 7.3 A
9.1 - 9.4 A
Kaolinite-Serpentine
9.6 - 10.1 A
Talc-Pyrophyllite
Mica
14.0 - 14.4 A
14.4 - 15.6 A
Smectite
(Montmorillonite)
oxygen
hydroxyl group
oxygen + hydroxyl group
(in projection)
Chlorite
octahedral cation (Al, Fe, Mg)
tetrahedral cation (dominantly Si)
59
4.3.2 Way and control of the alteration of silicate structures
In the case of the dominance of quartz the weathering processes
become blocked at the initial stage of the alteration but in other cases a
decomposition of the silicate structures proceeds towards the formation of
clay minerals as stable end stages of the silicate weathering. The
environmental factors of the weathering can determine the effective final
product of the alteration.
If the way of the weathering is interpreted as a special example of the
successive developments, the three above mentioned representatives of the
clay mineral group should be regarded as three main stages of the alteration.
The less altered sediments are regularly composed of the original silicate
structures together with illite (and with some other clay minerals of the
primitive weathering). The second stage of the alteration is represented by
the significant formation of montmorillonite (with remnants of illite) while
the strong alteration is indicated by the appearance of kaolinite. Since the
dissolution of the silicon-oxide is extremely poor under acidic pH
conditions, in the way of clay mineral formation slightly acidic
circumstances should be postulated. This stage of the alteration is called
siallitic weathering referring to the connected appearance of silicon and
aluminium in the same geochemical systems (i.e. crystal lattices). (Note that
due to the aluminium substitution of the silicon in the tetrahedral positions
the connected appearance of the silicon and aluminium is one of the most
common phenomena in the silicate structures.)
Under extreme conditions like the slightly alkaline pH conditions of
the tropical weathering crusts the complete decomposition of the silicate
structures has also a rather great importance. Under such conditions the
silicon-oxide has high solubility while the aluminium-oxide precipitates at
the place of the decomposition of the silicate structure (figure 27). Because
of the separation of the two regularly connected elements this way of the
alteration is called allitic weathering referring to the accumulation of the
aluminium at the place of the weathering (i.e. silicate decomposition).
60
Figure 27 Relationship between the pH conditions and solubility of the main
silicate forming elements (Si, Al)
Al2O3
10
9
8
millimol/l
7
6
5
4
3
2
1
SiO2
1
2
3
4
5
6
7
8
9
10 pH
4.3.3 Place and products of weathering
The environments of the alteration may be marine and terrestrial
conditions, and based on the rock cycle the source rocks of the weathering
may be volcanic (such as rhyolite tuffs), metamorphic (e.g. granitoid rocks)
or even sedimentary rocks. The initiative factor of the alteration may be the
superficial position directly, but it might be catalysed by other geological
processes (e.g. volcanic hydrothermal effects). The developing clay minerals
can form a nearly homogenous (monomineralic) sedimentary rock, the
name of which can be illite (with the dominance of illite), bentonite (with
that of montmorillonite) and kaolin (with kaolinite). Clay mineral
depositions of submarine or volcanic (hydrothermal) origin are also well
known even in the area of North Hungary.
The most important places of silicate weathering are the exposed
surfaces of dry terrains, which are at the same time the effective place of
soil formation. Silicate weathering can be regarded as the dominant process
in the development of soils. The forming clay minerals as the end products
of the weathering are the most important soil forming minerals. The nonorganic part of soils is mainly made of clay minerals, the colloidal
appearance, the water accumulation and the ion exchange capacity of which
basically determine the same characteristics of the forming soil.
61
Taiga
Temperate Mediterranean
and sub-tropical
climate
climate
Arid climate
Savannah
Equatorial
climate
62
ex
t
r
a
e
llit
me
ic
ly
w
e
z w
ath
o
ne eath
eri
er
ng
ed
be
d
ro
ck
9.6 - 10.1 A
9.1 - 9.4 A
smectite structure (montm.)
sli
wea ghtly
the
zon red
e
14.4 - 15.6 A
kaolinite structure
gly
on ered
r
t
s th
a
e
we zon
7.1 - 7.3 A
amorphous silicon content
Figure 28 Climatic control on formation of soil-forming (clay) minerals
(after Strakhov 1967)
illite structure
mica structure
Tundra
The successive characteristic of the weathering is well traceable in
the formation of soils of different climatic belts, since the clay mineral
association in a given climate related soil can be regarded as an analogy of
the plant association in a climate related biocoenosis. This regularity was
magnificently recognized and described by STRAKHOV (1967), who proved
the connection between the forming soil and the predominating climate
(figure 28).
Considering the most common trends we can conclude that the soils
of cold areas are determined by remnant fragments of the original minerals
with subordinate appearance of illite. Under cold temperate climate the
importance of illite increases significantly. In the temperate climates illite
and montmorillonite are the dominant soil forming minerals, while in the
sub-tropical soils the appearance of kaolinite refers to the more intensive
weathering. In tropical conditions in the area of deserts the intensity of
weathering is extremely low so the formation of clay minerals is a
subordinate process. In the area of the savannahs the soil forming clay
mineral is again the montmorillonite, while in the area of tropical forests the
total decomposition of the silicate structures can be seen within the frames
of allitic weathering. This latter case is how the tropical laterite soils form
with strong concentration of the residual aluminium and remnant bulks of
kaolinite. The diagenesis of laterite (at the place of weathering or after any
way of reworking) can lead to the formation of bauxite.
63
4.4 Carbonate rocks and carbonate depositional environments
4.4.1 Chemical and textural components of carbonate rocks
The main chemical components of carbonate rocks are the different
combinations of the [CO32-] divalent anion, so the most important rock
forming minerals are calcite and aragonite (CaCO3) and occasionally
dolomite ((Ca, Mg)CO3). The origin of the CaCO3 can be biogenic or
chemical (precipitation) and regularly both components can be traceable in
the same rock so the carbonate rocks (like limestones) can be regarded as
chemical-biogenic sedimentary rocks.
The texture of a carbonate rock contains two main parts of the rock
forming components, the micrite and the patite, the appearance of which
has an essential role in the classification of the carbonate rocks.
• Micrite is the dominant, enclosing matrix of the limestones, the
material of which is CaCO3, the particle size is regularly smaller than
0,005 mm (carbonate silt). Based on its origin it might have direct
biogenic origin (biomicrite) or might have been reworked
(allomicrite).
• The patite is the crystallized cement of the material, which regarding
its chemical components might be calcite or dolomite (dolopatite) and
based on the size of the crystals, which does not exceed 0,02 mm,
might be coarse or micropatite.
Other, frequently macroscopic components (i.e. grains) of the texture
might be e.g. bioclasts, ooliths, pisoliths, oncoids, intra- and extraclasts.
• Bioclasts are the definable reworked remnants of living creatures
(gastropods, molluscs, formaninifers etc.) enclosed into the fine
micrite. The ratio of these fossils can change from a few percent to the
dominant ratio as one of the most important aspect of the
classification.
• Ooliths are spherical or subspherical rock particles which have grown
by accretion around a nucleus, so the inner structure of the grain is
regularly concentric. Their size is between 0,2-2 mm.
• Pisoliths are similar to the ooliths but their size is bigger than 2 mm.
• Oncoids are spherical and concentric components as well, but the form
of the concentric spheres is undulatory. Their size ranges from
millimetre to centimetre; their origin is in connection with the biogenic
activity of algae.
• Intraclasts are the fragments of the given material which have been
torn up from the bottom by any sudden occasion (e.g. a storm) and
64
•
then reworked into the same material, so its age is nearly the same as
the enclosing matrix.
The extraclasts are significantly older than the enclosing material and
have been reworked into the given carbonate sediment from an outer
space (extrabasinal origin).
4.4.2 Classification of carbonate rocks
The petrographic classification of carbonate rocks is based on the
investigation of the textural particles identifiable under optical microscope.
There are several systems to describe the basic and intermediate types of the
textures (e.g. FOLK 1959, DUNHAM 1962) (figure 29).
Folk’s system is based on the ratio of the matrix and cement
distinguishing the micrite (dominantly with matrix) and patite (dominantly
with cement) types of the limestones as basic categories. The basis of the
further classification is the type of the enclosed frequently macroscopic
particles (grains) such as ooids, ooliths etc. So the name of a given
limestone in Folk’s system may be e.g. oomicrite referring to the ooliths
enclosed into micrite or oopatite referring to ooliths cemented by patite.
Similarly the bioclastic patite or extraclastic micrite can also be seen.
Dunham’s system refers to the relationship between the
macroscopic components (grains as bioclasts, ooliths, oncoids etc.) and the
enclosing matrix (micrite) and/or cement (patite) of the texture but contrary
to that of Folk it basically neglects the genetic type of the grains. If the
grains are in contact with each other and the inter-grain space is filled up
with patite, the texture is called grainstone while if the infilling material is
micrite, the name of the texture is packstone. If the grains are not in contact
with each other but their ratio exceeds 10 percent the name of the texture is
wackestone, if the ratio of the grains is less than 10 percent the name is
mudstone.
Comparing the two systems it can be seen that while Dunham’s
system focuses on the ratio and distribution of the grains, Folk’s system
emphasizes their genetic type. Nowadays the facies studies combine the two
aspects rather frequently. So in the special literature we can see the
combined nomenclature such as “bioclastic patite – grainstone” referring to
a grainstone after Dunham with bioclasts as grains in it or such as
“extraclastic micrite – wackestone” referring to a texture with extraclasts in
it, the ratio of which is more than 10 percent according to Dunham.
65
66
ooids
bioclasts
biopatite
oopatite
dominantly with cement
(patite)
biomicrite
oomicrite
dominantly with matrix
(micrite)
according to the enclosing material
Folk
packstone
grainstone
wackestone
grains > 10 %
mudstone
grains < 10 %
grains are not in contact with each other
filled with micrite
filled with patite
grains are in contact with each other
Dunham
Figure 29 Denomination of carbonate textures after Folk 1959 and Dunham
1962
according to the enclosed grains
4.4.3 The main environments of carbonate deposition
The main environment types of the carbonate formation can be
grouped as terrestrial and marine environments. The latter has much more
significant extension under the recent environmental circumstances and
similarly had much greater importance in the geological past.
Based on comparison of recent carbonate depositional environments
and fossil carbonate facies a generic sedimentological model has been
developed to summarize the most important marine belts with special
carbonate development (WILSON 1975). The so called standard facies belts
of Wilson are numbered from I to IX and range from the deep oceanic basin
to the tidal zone extending over the continental slope and shelf region as
well. Below we give a short list of the belts partly to show the most
important characteristics of the model but primarily to describe the main
characteristics of the sediment processes in the marine region (figure 30).
1. Wilson’s I facies zone, deep oceanic basin. Within these conditions the
basement is frequently deeper than the depth of the so-called
carbonate compensation depth (CCD). The CCD is formed by the
phenomena that the deep oceanic water is unsaturated with carbonate,
so the sinking carbonate skeletons of the living creatures are dissolved
and cannot reach the bottom. The only depositing sediments are the
sinking remnants of the planktonic organisms of siliceous skeleton
(e.g. radiolarians, diatoms) and the finest sediments coming from
terrestrial regions, the terrigenous clay minerals which can be
floated from the tidal region to the open ocean. The sedimentation
rate is very low (a few mm per 1000 years) so the forming
sedimentary rocks (red shales, siliceous shales, radiolarites) are
laminated.
2. Wilson’s II facies zone, the lower part of the continental slope (open
shelf - periplatform). The continental slope connects the nearly
uniform basement of the ocean with the elevated and dissected
continental margins, the submarine region of which is called shelf
zone. The angle of dip along the continental slope is about 5 degrees,
which allows the development of mass movements in the
oversaturated and unconsolidated fine submarine sediments. The
mixed material of water and solid particles is transported by
gravitation as slurry towards the lower parts of the slope. These
materials contain upward fining beds reflecting the sudden
transportation and the subsequent slow sedimentation of the material
from time to time. The repeated appearance of upward fining beds on
each other is called grading while the sediments with gradation are
named turbidites referring to the current circumstances in the course
67
of the sudden transportation. Turbidites might be proximal (coarse,
near the source) and distal (fine, far from the source) and are rather
general on the submarine slopes like on the continental slopes as well.
Graded sediments related to orogenic events are frequently called
flysch and are rather frequent in the Carpathian region.
In the lower part of the continental slope the distal turbidites are
common frequently changing with carbonate sediments. Here the CCD
can appear only temporarily so sediments with high and low carbonate
content can form a varied series with changing beds of limestones,
marls and shales. The limestones are typical pelagic limestones,
mudstones and wackestones with bioclasts of planktonic organisms.
3. Wilson’s III facies zone, the middle part of the continental slope (open
shelf - periplatform). Similar to the previous one, but it is continuously
above the CCD so the formation of the pelagic limestones is general
and permanent as well. In the material of the bioclastic mudstones and
wackestones the traces of slide can be seen also frequently reflecting
the gravitational instability of the material on the submarine slope.
4. Wilson’s IV facies zone, the upper part of the continental slope (the
front of the shelf - periplatform). This zone is situated above the socalled aragonite compensation depth (ACD). The ACD is formed by
the fact that the aragonite is a less stable form of CaCO3 than the
calcite, so the remnants of living creatures with aragonite skeleton are
more quickly dissolved than those of the species with calcite skeleton.
This way the appearance of fossils of organisms with aragonite
skeleton (e.g. ammonites) can be regarded as a sensitive depth
indicator and can provide a good tool to distinguish the upper part of
the continental slope from the deeper ones.
In this region of the continental slope a great amount of reworked
fragments coming from the eroded reef can be deposited so the
limestones are frequently grainstones and packstones with fragments
of reef-forming species (corals, crinoids etc.) properly enclosing the
maintained fossils of open marine species (e.g. ammonites). Because
of the uneven morphology mass movements of the unconsolidated
material may also appear reflected by intraclastic structures.
5. Wilson’s V facies zone, the reef of the shelf margin. On the outer
margin of the shelf, extended reefs built from reef-forming organisms
are frequent. These reefs are rather complex systems with encrusting
forms (e.g. red algae) towards the open sea because of the strong swell
of the sea and with patch reefs towards the inner belt of the shelf. The
in situ cemented material of the reef-forming organisms forms the
special facies type of the limestone the name of which is boundstone
in the facies system of Dunham.
68
6. Wilson’s VI facies zone, calcarenite piles of the shelf margin. A part of
the strongly fragmented material of the reef is frequently deposited
directly at or near its basement. Because of the short transportation the
dominant grain size of the material is greater than the common size of
the carbonate silt and it is equal to that of sand (arenite), so frequently
called calcarenite.
7. Wilson’s VII facies zone, the open shelf lagoon. The main belt of the
carbonate accumulation is the shelf zone. Contrary to the oceanic
basin, the sedimentation rate is very high, it reaches or occasionally
even exceeds 100 cm per 1000 years. The main source of the
carbonate formation and deposition is the biogenic activity of the
plankton algae vegetation in the upper zone of the sea water, which
can lead to the formation of floating “carbonate silt clouds”. The main
carbonate fraction depositing at the bottom of the lagoon is micrite, so
the main facies type is the mudstone and the wackestone. Due to the
intensive accumulation and the associated relatively stable ecological
conditions the forming sedimentary rocks are well and thickly bedded.
The enclosed grains (e.g. bioclasts) prove the extremely diverse form
of the marine fauna and flora. The common limestone types are
frequently named after the enclosed index fossil (foraminifers, shells,
gastropods, algae etc.). Because of the occasional occurrence of strong
bottom currents, intraclastic textures can also appear.
8. Wilson’s VIII facies zone, the closed lagoon and tidal plain. Towards
the beach zone the lagoon becomes shallower and anoxic conditions
can appear frequently. Due to the organic matter content the colour of
the forming limestone is frequently dark grey, blackish grey or black.
The material is well bedded patite or micrite frequently with oncoids.
Micrites with algae laminate can also occur. Because of the strong
effect of small sea-level changes short-scale cyclic series of limestone
types frequently with interbedding palaeosoil horizons can be seen.
9. Wilson’s IX facies zone, sabkha – salina. Near the beach zone the
terrigenous effect of the dry terrain become more and more dominant.
Because of the strong evaporation the geochemical characteristic of
the forming carbonates is frequently dolomitic and typical evaporates
(salt, anhydrite) can also appear in the limestone or between the
limestone beds. The limestone is frequently poorly bedded micrite
because of the strong effect of bioturbation of benthic species (e.g.
crustacean forms), which can destroy the original bedding. Special
biogenic forms of the tidal zones are the stromatolite structures, which
are special patch reefs formed by cyan bacteria, while the special
sedimentary structures are the small scale cross-beddings after the
ripple marks formed within the tidal current conditions.
69
70
VIII.
continental crust
IX.
inner shelf - platform
VII.
VI.
V.
IV.
III.
oceanic basin
oceanic crust
older sediments
I.
calcite compensation depth (CCD)
II.
aragonite compensation depth (ACD)
continental slope - outer shelf
- periplatform
open shelf
basin
II
I
III deep shelf margin
IV foreslope
IX sabkha - salina
VIII closed lagoon, tidal flat
VII open lagoon with currents
VI calcarenite
V bioherms (reef)
standard facies zones
Figure 30 Facies belts of the marine carbonate depositional environments
4.5 Organic materials coal and hydrocarbon deposits
4.5.1 Classification and lithology of coals
Coals are solid rocks formed by the alteration of plant remnants.
The starting point for coal formation is usually partially decayed vegetation
matter. By process of compaction and slight heating during burial, the
material is converted into coal. So the main environmental factors
determining carbonization are the pressure, the temperature and the time
of the alteration.
The different levels of the carbonization (coalification) are peat,
lignite (brown coal), sub-bituminous, bituminous coal (black coal), subanthracite, anthracite with gradually increasing carbon content. Peat and
lignite are light brown coal forms. They are rich in cellulose and well
identifiable plant remnants (xilite) can be seen in their texture. The colour of
brown coal is brownish black and well identified plant remnants are rare.
The material is frequently fragmented by desiccating cracks. The surface is
lustreless its streak colour is brown. The colour of the black coals is black,
and plant remnants can be seen in the texture only after special preparation.
Conchoidal fracture is typical. Anthracite has black colour and conchoidal
fracture as well.
The three important factors (temperature, pressure, time) of
coalification determine the way and intensity of the alteration together.
However, time has an essential role, since although the development may be
fast under extreme temperature or pressure, the geologic time can lead
surely to an alteration of the organic matter. This is the main reason why the
old coal deposits are regularly black coals (e.g. Carbon – West Germany,
East USA, Lower Jurassic – Mecsek Mts.) the relatively old deposits are
usually brown coals (Eocene – Trans Danubian Mountain Range, Miocene –
Nógrád Basin, Borsod Basin) while the young deposits are frequently lignite
(Pannonian – southern foreland of the Bükk and Mátra Mts.).
The individual chemical components of coals are macerals, the three
main groups of which are vitrinite (gelous form of humine), exinite
(aliphous hydrocarbons) and inertinite (oxidized form of humine). The
form and amount of the macerals can determine the luminosity of the coal
(vitrinite reflection), which is the base of the identification of coals and so
the determination of the level of coalification (figure 31).
71
Figure 31 Levels of coalification according to the change of the gas content
and vitrinite reflection (after Wallacher 1993)
Level of
LOM coalification
Gas
%
(Suggate 1959)
2
Vitrinite
reflection
(International
anthracological
scale)
Brown
coals
4
6
8
10
Black
coals
Semianthracite
1,0
1,5
2,0
10
2,5
16
18
0,5
40
35
30
25
20
15
12
14
45
Anthracite
5
LOM: Level of Organic Metamorphism (Hood et al 1975)
4.5.2 Palaeoenvironmental conditions of the coal deposits
The most important ecological conditions of coal formation are the
marsh environments with rich vegetation and great sediment accumulation
rate. The first is the source of the organic material, the second initiates the
occurrence of reductive geochemical conditions. Marsh environments may
be in connection with lakes (lacustrine) or with marine lagoons along a
beach zone (paralic).
72
Sedimentary basins with marshy conditions can be formed primarily
by tectonic events frequently combined with relative or general (eustatic)
sea-level changes. The deepening of the continental area together with the
relative increase of the sea-level may lead to the shifting of tidal and
shallow marine belts towards the inner part of the former dry terrain
(transgression). The newly flooded area may turn to marshes with
intensively spreading vegetation. Because of the proceeding transgression
the marshy region develops to a shallow marine area and the deceased plant
remnants will be buried by fine marine sediments. In this way as a result of
the permanent transgression anaerobic conditions with relatively high
pressure and increasing temperature are provided for the progressive
processes of the coalification. Since the tectonic events and transgression
processes are usually periodical, the formation of coal deposits is usually a
rhythmic phenomenon with several repetitions of coal seams above each
other (figure 32).
The elevation of a terrestrial area or a decrease of the sea-level can
lead to the shift of the tidal zone towards the inner part of the basin
(regression). It can lead to the formation of marshy conditions, and marshy
deposits on the surface of the formerly deposited shallow marine sediments.
However, because of the lack of subsequent burying, this situation, contrary
to a transgression, does not lead to the formation of significant coal deposits.
beach forest marsh
inner lagoon
algae
outer
lagoon sea
floating
plant
sapropelite fragments
geological
processes
facies
shifting
Figure 32 Model for development of complex coal deposits
progradation
transgression 3
progradation
transgression 2
progradation
transgression 1
progradation
4.5.3 Formation and migration of hydrocarbons
The naturally occurring liquid hydrocarbons are named crude oil
(or petroleum) the gaseous forms are called natural gases. The two forms
of the hydrocarbons are frequently associated with salt water and sometimes
with solid hydrocarbon forms (bitumen, asphalt) in the same deposit.
73
The formation of hydrocarbons is essentially related to deep marine
environments where, because of the lack of marine currents, the sea water
cannot be mixed. In this case the sinking decayed organic matter of the
deceased living creatures (dominantly plankton organisms) is accumulated
within anaerobic conditions. The so formed decaying fine, dominantly silty
sediment in the course of the diagenesis changes into oil shale, which can be
regarded as oil-source rock.
Because of the increasing pressure and temperature the organic
matter frequently segregates from the siliceous solid phase and as the fluid
or the gaseous phases emigrate from the oil-source rock along the effective
porosity of the enclosing rocks, usually horizontally or upwardly towards
the lower pressure. The migration of hydrocarbons proceeds until the front
of the migrating material reaches an impermeable zone. Here the oil and
the gas can be trapped and accumulated forming hydrocarbon
accumulation (figure 33). The impermeable zone may be a clayey
intercalation (cap rock) as overlying bed with anticlinal form
(stratigraphic trap), a tectonic fault where a thick sand body and clayey
material can be in contact laterally (structural trap) or frequently even the
combination of the two cases (combined trap). The term oil pool is used
for an area with oil bearing reservoir rocks below the cap rock. The
reservoir rocks are frequently sandstone or limestone with high effective
porosity.
Figure 33 Some examples of petrolepum accumulation (oil traps)
Anticlinal
Unconformity
Facies
Tectonic
Explanation:
Impermeable rock
Other rock
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Gas bearing
Oil bearing
Water bearing
4.6 Evaporites
There are cases when the soluble matter in the water (marine, lake or
even groundwater) is precipitated due to the evaporation of the solution
beyond the saturation point of the dissolved salt. One of the most typical
examples can be a separated lagoon under arid climate.
During the evaporation the order of the precipitation is determined
by the solubility of the dissolved material. The crystallization starts with
carbonate compounds as the less soluble phases. Depending on the
concentration of cations calcite (CaCO3) or dolomite ((Ca,Mg)CO3) can
appear with carbonate (CO3-) anion at the bottom of the lagoon. The next
anion complex is the sulphate (SO42-) usually together with Ca as cation,
forming gypsum (CaSO4×2H2O) or in the case of loosing the water content
of the crystal lattice anhydrite (CaSO4). Then halite (NaCl) is precipitated
followed by potash salts such as sylvite (KCl) with magnesium (e.g.
carnallite – KCl×MgCl2×6H2O) at last. These evaporate sediments regularly
lie on each other depending on the order of the precipitation i.e. forming the
under- and overlying associates of each other (figure 34).
Figure 34 Distribution of the salt concentration and evaporation in a lagoon
with open circulation (after Krumbein and Sloss 1963)
dune
3
1 ,2
chl
ori
de
beach sand
0
3
3
m
g /c
1 ,1
5
m
g /c
1 ,1
0
m
g /c
3
1 ,0
5
m
g /c
open sea
dune
sulphate
carbonate
The ideal order of the precipitation can be disturbed by outside
processes. It can happen for example that the barrier between the lagoon and
the open marine is destroyed by storm events or because of transgression
(sea-level increasing) the sea water intrudes into the lagoon. In these cases
the actual stage of the evaporation can stop and the concentrated
(hypersaline) water of the lagoon will mix with less saline marine water.
This can lengthen the actual (e.g. sulphate or halite) precipitation or can
reject the evaporation to an earlier stage of the development (from the halite
stage back to the sulphate one). These disturbances can cause the formation
of extremely thick beds or sedimentary mixing or interfingerings of the
different evaporate materials.
75
Selected titles for further reading
BOUMA, A. H.: Methods for the Study of Sedimentary Structures. WileyInterscience, New York, 1969.
CARVER, R. E.: Procedures in Sedimentary Petrology. Wiley-Interscience,
New York, 1971.
GALLOWAY, W. E. & HOBDAY, D. K.: Terrigenous Clastic Depositional
Systems. Springer-Verlag, New York, 1983.
PETTIJOHN, F. J. & POTTER P. E.: Atlas and Glossary of Primary Sedimentary
structures. Springer-Verlag, New York, 1964.
TWENHOFEL W. H.: Principles of sedimentation. McGraw-Hill Book
Company Inc., New York 1950.
76
5 METAMORPHIC ROCKS
5.1 Process and classification of metamorphism
Mineral phases both siliceous and non-siliceous ones represent the
stable crystalline phase of their chemical composition under the pressure (P)
and temperature (T) conditions of their forming processes. Under changing
PT conditions these crystalline phases become unstable and thus take part in
the processes of metamorphism, which means a well defined recrystallization of the crystal lattice. Metamorphic processes take place in the
solid state of a rock, so its melting (even partial) and its solution cannot be
regarded as a part of the metamorphism. The end products of metamorphism
are the metamorphic rocks representing interaction between the
metamorphic agents and the parent (pre-metamorphic) rocks.
Geologists distinguish three main types of metamorphism from the
aspect of the dominant metamorphic agencies.
In the case of thermal metamorphism (thermometamorphism) the
main factor is the heat without significant pressure effects, usually
associated with igneous intrusions. In these cases a contact aureole can be
formed along the boundary of the igneous mass and the enclosing rocks, so
the term contact metamorphism is also frequently used.
The term dynamic metamorphism refers to the dominant effect of
pressure, when any kind of localised stress can lead to a significant change
of the material frequently under relatively low temperature conditions. This
type of metamorphism can characterize e.g. the elongated belts of the
obduction zones along the outer belts of island arcs.
Dynamothermal metamorphism means joint effects of increasing
pressure and temperature, producing a wide range of new minerals. This
form of metamorphism is usually connected with orogenic belts and its
effect can be traced along hundreds of kilometres wide and thousands of
kilometres long zones. The widespread character of this metamorphism is
indicated by its “regional-” term.
According to the characteristics of the parent rock we can talk about
orthometamorphism when the pre-metamorphic rock is igneous and about
parametamorphism when it is sedimentary. In the case of the renewed
metamorphism of a previously metamorphosed rock we talk about
polymetamorphic rocks.
Metamorphic processes are classified as isochemical when there are
no significant differences in the chemical composition of the parent and that
of the metamorphic rocks. If any important change in the chemical
composition can be observed the term allochemical metamorphism is used.
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5.2 Thermal metamorphism
5.2.1 Terminology of the thermal metamorphism
When the hot magma rises into the crust it forms a contact zone with
the rock of its immediate surroundings (wall rock). The contact zone in the
terminology of mining is called skarn (figure 35). Its width can range from
a few metres to 2-3000 m depending mainly on the extension of the
inclusion. The enclosing rocks regarding their main geochemical component
usually can be siliceous (e.g. other igneous rock, sandstone etc.) or
carbonaceous (e.g. limestone, dolomite). In the former case we talk about
silicate-silicate skarn while in the latter one about lime-silicate skarn. The
skarn and the metamorphic alterations are extended to both the igneous and
the wall rocks. The igneous belt of the contact aureole is called endoskarn,
the encrusting belt within the wall rocks is the exoskarn. In the zone of the
contact aureole with higher temperature the zone of pyroxene hornfels,
while in the zone of lower temperature the amphibole hornfels are formed.
Figure 35 Main forms of the contact metamorphism
ck
all ro
tic w
s
a
l
c
i
silic
exoskarn
silicate-silicate skarn
endoskarn
water emigration
(contact metasomatism)
k
roc
all
w
ate
on
b
r
ca
igneous intrusion
lime-silicate skarn
5.2.2 Processes and mineral associations in the thermal metamorphism
Because of the increasing heat decarbonisation and dehydration
can appear in the contact aureole as essential metamorphic processes. Some
of the characteristic forming metamorphic minerals in the hornfels are
pyroxenes (e.g. diopside, enstatite) and amphiboles, orthosilicate
structures (forsterite, garnets e.g. grossularite), cyclosilicates (cordierite)
78
sheet silicate structures (flogopite, biotite), and aluminium minerals
(corundum).
If the intruding magma has a significant water content, due to the
emigrating water solutions contact metasomatism can become an
important accompanying process of the contact metamorphism. It means the
substitution of any cation with another or the complete replacement of one
mineral by another without losing the original texture. Contact
metasomatism can lead to the formation of industrially perspective
metalliferous ore deposits such as copper ores of Utah, Arizona, and New
Mexico.
If the magma intrudes into sedimentary rocks rich in pore water, the
thermal effect can lead to the appearance of migrating hot water solutions.
In this way hydrothermal phenomena can also accompany the thermal
metamorphism leading to the formation of subdominant hydrothermal
mineral associations even causing the alteration of the intruding magma.
If the mass of the intruding magma is not large enough to keep a
long-lasting thermal effect and both the magma and the wall rock are too dry
to induce metasomatic or hydrothermal processes, the phenomena of the
thermal metamorphism are subdominant. Nevertheless even in these cases
there are small, partially re-melted and frequently vitreous rock fragments
along the boundary of the intruding magma mass indicating temporary
thermal effects.
5.3 Dynamothermal metamorphism
5.3.1 Classification and processes of the dinamothermal metamorphism –
mineral facies
In the course of regional metamorphism the alteration (recrystallization) of the pre-metamorphic (parent) rock is basically determined
by the accompanied changes of temperature and pressure. According to the
gradual and progressive increase of pressure and temperature the process of
the progressive dynamothermal metamorphism can be divided into four
separate stages, namely the anchi-, epi-, meso- and katazones.
These zones can be well defined by the mineral associations and the
texture of the forming metamorphic rock. Because of the increasing
temperature, decarbonisation and dehydration of the minerals are
important processes accompanied by structural reorganization of crystal
lattices under the increasing pressure.
Although dynamothermal metamorphism is determined by pressure
and temperature, the geochemical composition (e.g. silicon content) of the
pre-metamorphic material also has a determinative role in the formation of
metamorphic mineral associations. So the forming metamorphic mineral
association, depending mainly on the degree of temperature, pressure and
79
on geochemical components refers directly to the conditions of the
metamorphic stage.
This is well expressed in the determination of the so-called
metamorphic mineral facies, which in the sense of Eskola (1921)
“comprises all the rocks that have originated under temperature and pressure
conditions so similar that a definite chemical composition has resulted in the
same set of minerals quite regardless of their mode of crystallization,
whether from magma or aqueous solution or gas, and whether by direct
crystallization from solution … or by gradual change of earlier minerals”.
Hereinafter we show the stages of the metamorphism partly to show
the process and partly to describe the metamorphic rocks (mineral
associations) forming in the different zones (figure 36).
5.3.2 Stages of the dynamothermal metamorphism
The anchizone comprises metamorphic processes hardly separable
from those of diagenesis. It is rather common in the lower part of basin
filling sedimentary series with great thickness (e.g. Pannonian series in
south Hungary) where the increasing lithostatic pressure together with the
increasing geothermal effects can lead to a gradual change from the
advanced diagenesis to the initial metamorphism.
The minerals are usually the same as in the pre-metamorphic
sedimentary rocks, but a progressive change can be observed e.g. in the
organization of clay minerals (e.g. increasing illite crystallinity, occurrence
of chlorite-montmorillonite mixed layer structures). The forming
metamorphic minerals are zeolites (zeolite facies) namely laumontite and
heulandite.
In the epizone the re-crystallisation takes place under low
temperature and often under relatively high shearing stress. So the forming
rocks exhibit a foliated texture. In the epimetamorphism the so-called
Greenschist Facies appears referring to temperature ranging from 100 to
250 °C, pressure is about 5 kbar and the so-called Blueschist Facies
referring to lower temperature but higher (10-12 kbar) pressure. The latter
frequently takes place at the forearc region of the converging plate margins
and because of its high pressure and low temperature conditions it can be
regarded as a type of dinamometamorphism.
80
Figure 36 Subdvision of the regional metamorphism according to the degree
of the metamorphism and the original (parent) rock
Epizone
Mesozone
Katazone
Metamorphic
zone
temperature: 100-250 0C
temperature: 250-500 0C
temperature: 500-650 0C
pressure: 5 kbar
pressure: 5-6 kbar
pressure: 3-7 kbar
Metamorphic
facies
Greenschist facies
Blueschist facies
Rock forming
minerals
talc, dolomite, magnezite,
quartz, albite
(after Ramberg)
Amphibolite facies
biotite, tremolite, amphibole,
diopside, plagiochlase,
garnet, staurolite
Granulite facies
Eclogite facies
garnet, pyroxene, disthene,
sillimanite, rutile, ilmenite,
omphacite, olivine
Characteristic
texture
Orthometamorphism (parent rocks have magmatic origin)
Forming metamorphic rock
Parent rock
Peridotite
Basalt
Andesite
DacyteRhyolite
Talc schist
Pyroxene hornfels
Eclogite
Green schist
Chlorite schist
Porphyroid
Amphibolite
Mica gneiss
Feldspar gneiss
(orthogneiss)
Granulite
Parametamorphism (parent rocks have sedimentary origin)
Forming metamorphic rock
Parent rock
Sandstone
Clayey
sandstone
Clay
Marl
Quartzite
Sericite quartzite
Phyllite
Carbonate phyllite
Mica quartzite
Mica schist
Carbonate mica schist
Bauxite
Smirgel
Limestone
Marble
Granulite
Gneiss
(paragneiss)
81
The dominant minerals are quartz, albite, muscovite and in the
greenschist facies kaolinite, chlorite, serpentinite, talc, epidote, actinolite
from the silicate group and calcite dolomite and magnesite from carbonates.
Accessory minerals can be tourmaline and rutile. The important index
mineral of the blueschist facies is the glaucophane (a species of
amphiboles).
In the course of orthometamorphism from the basic and neutral
igneous rocks as pre-metamorphic (parent) rocks, green metamorphic schists
are formed such as chlorite schist, talc schist, serpentinites, epidote-albite
schist and actinolite schist.
In the course of parametamorphism, from pelitic sediments the
general metamorphic rock is phyllite, the main rock forming minerals of
which are mica structures (sericite), chlorite, quartz, graphite (from the
enclosed organic material). Because of the oriented distribution of the mica
structures, in the texture of the phyllite the cleavage structure is common
frequently with sharp angle to the original bedding.
In the mesozone the temperature is somewhat higher but not above
500 °C, the pressure is about 5-6 kbar as it is indicated by the formation of
the Amphibolite Facies. The foliated texture is also common similarly to
the epimetamorphism referring to the important role of the oriented stress.
The dominant minerals are amphiboles (hornblende) and mica
structures (muscovite and biotite) depending on the silicon content of the
geochemical system. Garnets (grossularite, almandine) also has an
important indicative role, while in rocks deficient in silica olivine, spinel
and corundum may appear.
In the way of orthometamorphism from the basic and neutral rocks
as pre-metamorphic rocks, amphibolite can be formed as one of the most
important products of mesometamorphism. It has anisotropic texture due to
the oriented distribution of the amphibole crystals. Plagioclase has also
dominant role in the texture of the rock.
In the parametamorphism from sedimentary rocks with high silicon
content as parent rock, mica schist can be formed. The foliated texture
contains oriented mica structures (muscovite and biotite) and the garnet is
also an important phase of the association. Quartz and graphite can also
appear especially in mica schist formed under lower pressure (2-3 kbar).
In the way of metamorphism the stages of meso- and
katametamorphism show gradual transition. Gneiss represents the
intermediate mineral association between the two stages of the
metamorphism with less foliated, occasionally simply oriented texture with
82
higher quartz and feldspar content contrary e.g. to the mica schists.
According to their parent rocks ortho- and paragneiss can be distinguished.
The katazone represents the stage of metamorphism with the highest
temperature between 500-650 °C and pressure (3-7 kbar) indicated by the
appearance of the Granulite and Eclogite Facies. The pressure is frequently
hydrostatic so the foliated texture is usually substituted by simply oriented
or totally isotropic ones.
The dominant forming minerals are quartz, feldspar, disthene
pyroxene, garnet and mica structures.
Some kinds of granulites (pyroxenic granulites) represent the
highest degree of metamorphism in the way of the orthometamorphism.
Their texture is isotropic containing dominantly pyroxenes and plagioclase
(andesine). In the orthometamorphism from the pre-metamorphic rocks with
low silicon content, eclogite may be formed. Its texture contain dominantly
pyroxenes (omphacite) enclosing Mg-garnets (pirope).
In the parametamorphism granulite (quartz-feldspar granulite) also
has a dominant role. The main mineral rock forming minerals are quartz,
orthoclase, plagiochlase, distene and garnet. The texture is anisotropic,
occasionally foliated.
5.4 The upper boundary of the metamorphism - ultrametamorphism
The final phase of metamorphism takes place at extreme PT
conditions and is called ultrametamorphism. It is usually active in the
deepest parts of orogene regions (see plate tectonics).
When extremely high pressure and temperature is exerted upon a
metamorphic rock it will be partly melted so that fluid and solid state will be
present together. This partial melting or re-melting of the metamorphic
rock is called anatexis. In the course of anatexis the material having a lower
melting point, due to the continuous presence of volatiles and alkaline
elements will be melted first and those having higher melting point will
melt later. Therefore this process results in a so-called metamorphic
differentiation.
The result of the partial re-melting is the migma (meaning mixed
magma), that is the mixture of melted and solid rock fragments. The resolidified part of the migma is the so called neosome (meaning re-born
rock) and the original residual migma material is called paleosome. Due to
the metamorphic differentiation the different parts of the material of the
migma will be segregated. This segregation results in the alternation of
lighter and darker bands in the neosome material. The lighter one is called
leucosome and the darker one is called melanosome. Leucosome is
83
composed of mainly quartz and feldspars while melansome consists of
biotite and amphibole.
Rocks that are produced by the solidification of the migma are
migmatites, the parent material of which is a mixture of re-melted
orthometamorphic and parametamorphic rocks. The end of the
metamorphism is the complete re-melting of a given rock. This process is
called palingenesis. Palingenesis produces palingenous rocks that are the
granites. These granites represent “life” after metamorphism for a rock that
might have gone through all of the steps in the rock cycle and when elevated
up to the surface can start the whole rock cycle all over again.
Selected titles for further reading
AUGUSTITHIS, S. S.: Atlas of the Textural Patterns of Metamorphosed
(Transformed and Deformed) Rocks and their Genetic Significance.
Theophrastus PublicationsS. A., Athens 1985.
BARTH T. F. W.: Theoretical Petrology. John Wiley and Sons Inc. New York
1962.
READ, H. H.: The Granite Controversy. Thomas Murby & Co, London,
1957.
84