Igneous Rocks and Magma

1) Introduction

Igneous rocks are defined as those rocks which have crystallised
from a silicate melt (magma) either within the Earth or at the surface. If the magma cools at depth, slow cooling will occur and a coarse-grained igneous rock will result. If cooling is more rapid, a medium-grained rock is formed ( for example at shallow depths – the hypabyssal environment – sills, centres of dykes), and if the magma erupts at the surface, cooling is very quick so a fine-grained volcanic, rock is formed. With very rapid chilling, a volcanic glass can be formed (obsidian).

Magma contains dissolved gases that remain in solution under pressure. When the pressure is released this can lead to an eruption. This pressure release is called exsolution. Such a process can lead to a fragmental or frothy rock being erupted called a pyroclastic eruption. Magma erupting underwater (sea, lake) produces characteristic shapes known as pillow lavas. These have glassy chilled margins with interiors full of holes, where gases were exsolved but trapped within the rock. The holes are termed vesicles (the sites of former gas bubbles) and are also common in lava flows.

Magmas intruded into small fissures near the surface of the Earth form dykes if vertical and discordant to the bedding, and sills if near horizontal and concordant with the bedding. These dykes and sills may have chilled margins and coarser interiors, and often show cooling joints perpendicular to the cooling surfaces.

2) Classification of igneous rocks

Igneous rocks may be classified using grain-size.
This classification roughly corresponds to plutonic, hypabyssal and volcanic environments. The crystals of coarse-grained rocks can be seen easily with the naked eye; those of medium-grained rocks need a hand-lens; those of fine-grained rocks require a microscope to be resolved.

Some igneous rocks show evidence of having undergone two stages of cooling. The magma may pause en route to the final place of intrusion or extrusion, and cool slightly. Crystals will form and grow. Subsequently eruption or final intrusion takes place and the remainder of the magma will crystallise with a generally finer- grained texture. So we see large crystals (termed phenocrysts) in a finer-grained matrix (called the groundmass). This is known as porphyritic texture.

Igneous rocks can also be classified using mineralogy and chemistry. Textural terms (above) are also useful because the same magma (for example a magma of basaltic composition) can crystallise under different conditions to give very different looking rocks. The chemical composition of the rocks will be the same, and the mineralogy may or may not be similar depending on the physical history, but the rocks may look very different.

3) Chemical definitions

Chemical analysis of igneous rocks gives a classification according to their chemical compositions. A common classification is based on the amount of silica in the rock. Note that this is the chemical amount of silicon dioxide (SiO2), not the quantity of the mineral quartz.

4) Mineralogical definitions

Igneous rocks are composed of varying percentages of mafic (ferromagnesian) minerals such as olivine, pyroxene, amphibole, biotite mica, and felsic minerals such as plagioclase, alkali feldspar and quartz. Generally, mafic minerals tend to be dark in colour and felsic ones light in colour, although this is not always the case.

The percentage of mafic minerals in a rock is called the Colour Index (C.I.) A high colour index is associated with ultrabasic and basic rocks containing >50% mafic minerals. These rocks are sometimes referred to as melanocratic. Colour indices of 30 – 50% are referred to as mesocratic, and are associated with intermediate rocks, while a low C.I. of <30%, referred to as leucocratic, is associated with acid rocks. Do not be led astray by the black obsidian glass which is not a mafic mineral.


We define igneous rocks using a combination of the percentages of quartz, alkali feldspars, plagioclase feldspars and ferromagnesian (mafic) minerals. Grain size also plays a part in naming igneous rocks.

Ultramafic rocks; These contain nearly 100% ferromagnesian minerals.

Dunite is a rock which is rich in the mineral olivine. Pyroxenite is a rock rich in the mineral pyroxene. Peridotite is a rock consisting mainly of olivine and pyroxene. All of these are intrusive rocks and therefore have a coarse or medium grain size. Extrusive rocks of this composition and mineralogy are called komatiites. These are rare today, but more abundant in the early history of the Earth when the mantle was hotter.

Basic rocks:

Basic rocks contain approximately 50% plagioclase (composition typically of labradorite >An50) and 50% mafic minerals (pyroxene). Olivine may also be present, in which case we can preface the name of the rock with the mineral name, for example, olivine basalt. Coarse-grained rocks of this composition are called gabbros; medium-grained ones are dolerite, and fine-grained ones are basalts.

Intermediate rocks:

The diorites are characterised by felsic minerals such as plagioclase (usually andesine), and ferromagnesian minerals (mafics) which may include hornblende and biotite (rarely pyroxene). The ratios are generally such that plagioclase is more abundant than mafics. Quartz and alkali feldspars also may be present. In coarse-grained form they are termed diorite or quartz diorite, and when fine-grained, andesite or dacite.

Acidic rocks:

Quartz and alkali feldspar are abundant (>50% of the total). Plagioclase is andesine or oligoclase. Alkali feldspar includes albite, microcline, orthoclase. Mafic minerals are less abundant and are commonly biotite or hornblende. Muscovite (white mica) also occurs in some granites. Coarse-grained rocks of this composition are called granodiorite or granites and fine-grained varieties are rhyodacite or rhyolite. Other types occur depending on the cooling history.

The relationship between SiO2 and the relative proportions of different minerals is illustrated by a diagrammatic model. If a sample has 50% silica, then using the model it could be ascertained that it should contain 5% olivine, 70% pyroxene and 25% plagioclase feldspar. If it is coarse-grained it would be a gabbro and if fine, a basalt.

5) Origin of basaltic magmas

Magmas are formed by the melting of pre-existing rocks. This occurs in the upper mantle, perhaps where water has lowered the melting temperature. The upper mantle is composed of peridotite.


Peridotite is a complex solid, commonly consisting of 4 minerals (olivine, clinopyroxene, orthopyroxene and spinel or garnet). Each of these minerals melts at a different temperature and so peridotite does not completely melt at any single temperature.

As the temperature rises, some minerals melt and others remain solid. The melting process is aided by the presence of water or by the release of pressure. The parts that melt first (consisting of the minerals with the lowest melting points) rise to higher crustal levels to become rocks of different overall compositions.

Mantle peridotites have been subjected to melting experiments at appropriate pressures in the laboratory, and basaltic magmas are produced in such experiments. Therefore, low P and S wave velocities in the upper mantle are attributed to local basaltic melts. The process of melting a solid to form a melt of different composition is called partial melting.

Evolution of magmas:


Once basaltic magmas have been formed by partial melting of the mantle, they rise upward towards the surface of the Earth. This is due to their buoyancy arising from the lower density of the partial melt than that of its parent. During the process of rising magmas cool and crystals begin to form. If these crystals are removed from the melt (for example they fall to the floor of the magma chamber) or are otherwise prevented from continuously reacting with the magma (formation of zoned crystals), then the magma will change in composition. This is the process of fractional crystallisation and it leads to the formation of families of related igneous rocks (sometimes called a suite or a rock series). Rocks such as basalts, andesites, dacites and rhyolites may all be erupted from a volcano beneath which a magma chamber is undergoing fractional crystallisation.This process can be repeated until acidic rocks (rhyolites) are formed. However, most granitic magmas are probably formed by another process: partial melting of crustal rocks.

Sources and References:https://giphy.com/gifs/12442m0WZFHvuE


Earth Vs Venus – November 2017

A Lecture I attended a last year! Found the video by the Geological Society of London!

Copied Video’s Description:
Why Earth developed into the crucible of life, and Venus into a hostile wasteland

The present-day differences in the expression and intensity of volcanism on the planets of the inner solar system serves a testament to the dynamic nature of planetary formation and evolution. For example, Earth and Venus are colloquially referred to as sister planets because of their similar size and composition. However, their contrasting volcanology, atmospheric mass and chemistry, climate, and geomorphology are striking.
In short, the Venusian atmosphere and surface contains five orders of magnitude less water than Earth and the average surface temperature on Venus is 460 °C. In addition, Venus is a relatively flat planet, where only 2% of the surface is shows any appreciable topography. Earth, by contrast, has a wet and cold surface with a bimodal topography (e.g. orogenic belts and ocean basins). Suffice to say, these are not identical siblings.
Here I will show how we can combine data from rock-deformation experiments with the chemistry of the Venusian and Terrestrial atmospheres to explain the flatness and relative volcanic quiescence of Venus. In short, I will outline why Earth developed into the crucible of life, and Venus into a hostile wasteland.

Sami Mikhail (University of St Andrew’s)
Dr. Mikhail is a lecturer in Earth Sciences at the University of St Andrews (since May 2015), after spending two years as a Carnegie Postdoctoral Fellow at the Geophysical Laboratory (Washington DC, USA) and a couple of postdoctoral positions at the Universities of Bristol and Edinburgh (UK).
Prior to this Dr. Mikhail gained an BSc in Geology from Kingston University (2006), an MSc in Isotope Geochemistry from Royal Holloway and Bedford New College (2007) and a PhD on the origin of diamond-forming carbon at University College London (2011).
The motivation behind Dr. Mikhail’s research is to understand how the interior of a planet affects and controls the composition of its surface and to long-term habitability. Dr. Mikhail’s approach combines investigations of natural samples with high-pressure and -temperature experiments and theoretical models.
Dr. Mikhail has worked on diverse projects such as the source of Icelandic volcanism, diamond-formation in the deep Earth, and more recently, on linking mantle processes to atmospheric chemistry on Earth, Mars, and Venus.

Geological Society Youtube Channel:

Lecture Details:

My Personal Blog:

Retrograde Orbits

(note: This animation has no audio track.) – The Open University

Although many moons in the Solar System follow prograde orbits, there are some notable exceptions. The gas giant planets Jupiter, Saturn, Uranus and Neptune have several small outer moons that follow retrograde orbits; this means that they orbit their planet in the opposite direction to the planet’s rotation. In a retrograde orbit, a moon revolves in its orbit in the opposite direction from that in which the planet rotates about its axis.

Video by The Open University.

More information at https://www.futurelearn.com

Physical Properties of Minerals

1) What is a mineral?

A mineral is a naturally occurring inorganic crystalline substance, whose compositions are either fixed or vary between certain fixed limits. This excludes, for example, artificial diamonds, coal, volcanic glass. What makes each mineral unique is a combination of its chemical composition and the internal arrangement of its constituent atoms.

The chemical composition may be fixed, as in quartz (SiO2) in which, for every single atom of silicon there are 2 atoms of oxygen. Or it may be variable for example the mineral olivine, which varies between two formulae; Mg2SiO4 and Fe2SiO4. Some minerals have a single chemical composition but a different arrangement of atoms in their three-dimensional structure. This is called “polymorphism” and is well seen in the polymorphs of carbon: graphite and diamond. One is grey/black and very soft, the other is usually colourless and is the hardest natural substance.

There are many different classes of minerals, but only the most common will be examined. These are: silicates (containing the SiO44^- anion), carbonates (with the CO3^2- anion), halides (which contain F^-, Cl), sulphides (with S^2-), sulphates (SO4^2-), oxides and native metals. The following are examples of each of these classes, which will become familiar in this lecture.

2) Crystallinity

The extent to which atomic structure controls the outward shape of a substance is called its crystallinity. For example, quartz can develop in 4 ways. Firstly, as a crystal showing crystal faces, which is termed CRYSTALLISED.

Secondly, as irregular grains of quartz with fully developed internal structure, but not displaying crystal faces, which is termed CRYSTALLINE.

Thirdly, as a finely distributed crystalline aggregate called chalcedony whose grains are only visible under powerful magnification, this is termed CRYPTOCRYSTALLINE.

Fourthly, as a precipitate such as OPAL where there is no regular arrangement of atoms which is termed AMORPHOUS.

3) Physical Properties

What physical properties can be used to identify minerals?

In hand-specimen, minerals can be identified using a combination of physical properties, and it is these that will be studied in this lecture and practical.

a) Crystal Form/Habit

The characteristic shape of an individual crystal of a mineral is called its “habit”. Sometimes it is possible to identify a mineral from its habit alone, for example, quartz often forms a six-sided column with a set of pyramid-like faces at the top. We would describe this as prismatic.There are many common terms to describe the habit of crystals for example:

  • acicular (needle-like);
  • bladed (elongate crystals flattened in one direction);
  • botryoidal (rounded masses looking like bunches of grapes);
  • reniform (kidney like);
  • fibrous (groups of parallel thread-like crystals);
  • massive (no regular form);
  • platy (very flat crystals);
  • prismatic (elongated crystals with well developed prism faces);
  • tabular (crystals slightly flattened in one direction).

A fully developed form is referred to as euhedral; the opposite, an
irregular form, is called anhedral.

b) Hardness

A diamond is considered to be very hard because no other mineral can scratch it. Quartz can scratch a large number of other minerals. A scale has been devised to describe the hardness of minerals; Mohs’ scale. It varies from 1 (softest) to 10 (hardest). To test the hardness of an unknown mineral a specimen set of minerals called a hardness set is used. Starting with a hard specimen such as corundum the user tries with each in turn to scratch the unknown mineral until one is found that will not scratch the unknown mineral. The hardness of the unknown mineral is between the number of the mineral that will scratch it and the one that will not.

It is useful in the field to know the hardnesses of some everyday objects. For example, a fingernail has a hardness slightly over 2, and can scratch minerals of hardness 2 or less. Teeth have a hardness of around 4 and a copper coin 4.5 to 5; a steel knife blade is 6 and a hard file is around 7. Silicate minerals vary in hardness – the softest is talc (1) the hardest is topaz (8). Oxides, sulphides and many native minerals are very soft (less than 3), but exceptions include corundum Al2O3 which has a hardness of 9.

c) Specific Gravity / Density

Some minerals have an unusually high or low density (mass per unit volume), although most are between 2.5-3.0 g/cm^3. Specific gravity, or relative density, is the ratio of the mass of a mineral to the mass of an equal volume of water. However, it is difficult to estimate in the field except in some cases. Some common ore minerals are particularly heavy, for example, galena PbS (7.5), cassiterite SnO2 (6.9). This is because of the high atomic weight of the elements Pb and Sn. One commonly occurring mineral, barite (barytes) BaSO4, is also unusually heavy (4.5) compared with other superficially similar minerals such as calcite.

d) Colour

Colour is rarely a reliable indicator of the identity of a mineral. This is because in many minerals, impurities elements in trace amounts) can change the colour drastically. An example is a quartz, which is colourless when pure (“rock crystal”) but occurs in the coloured forms of amethyst (purple), rose quartz (pink), smokey quartz (grey) and others. Fluorite (CaF2) also occurs in purple, blue, green and yellow varieties. Some minerals are less variable, and in some cases, the colour is characteristic, for example, malachite (green), pyrite (golden), Galena (silver-grey). Nevertheless, it is better to identify these minerals on other characteristics in addition to the colour, for example, using cleavage or habit.

e) Streak

A slightly more reliable method of determining a mineral’s true colour is to use its streak. This is the name given to the powdered material left behind when a mineral is rubbed on an unglazed porcelain plate. Most pale-coloured minerals have a white streak which is of little use in identification, but the method is very useful with dark coloured opaque minerals such as hematite (Fe2O3) which has a characteristic red/brown streak, or pyrite (FeS2) which has a characteristic black streak.

f) Degree of transparency

Terms such as transparent, translucent and opaque are used to describe the degree to which a mineral can transmit light. However, this often depends on the thickness of a specimen and other factors such as internal inclusions. It is, thus, not a particularly useful guide for mineral identification. However, most ore minerals (pyrite, galena) are opaque, while many silicates are translucent to transparent.

g) Cleavage

Cleavage in minerals is the splitting or breaking of a crystal along planar surfaces which are determined by the crystal structure. There are often only a small number of possible cleavage planes in a mineral, whereas there can be many possible crystal faces. Cleavage of mica is along sheets, so that parallel smooth flat surfaces can be seen. This is due to weak bonds between the sheets in the structure of mica. As mica can only cleave in this one plane, it is said to have only one cleavage direction. Other minerals such as calcite have three excellent cleavage directions, giving a rhomboid shape. Distinctive patterns of cleavage are good identifying marks for many minerals. Galena and halite both have three good cleavages at 90°, yielding almost perfect cubes.

This was a rather boring post, but informative 🙂

Sources and References:




The Structure of the Earth

Physical Properties of the Earth

The Earth is an oblate spheroid, being slightly flattened at the

Equatorial radius = 6378 km Polar radius = 6357 km

These measurements are calculated on the assumption that the Earth’s surface is smooth, but this is only an approximation since it disregards mountains and ocean depths. However, the difference between the height of Mount Everest and the depth of the Marianas trench is only about 20 km. Most land is concentrated in seven continents each fringed by shallow seas (flooded continent). Separating these are a number of major oceans including the Pacific, Atlantic and the Indian oceans.

It was Cavendish in 1798 who first calculated the mass of the Earth as 5.977 x 1024kg, and since its volume is known (from 4/3 ∏ r^3 where r is the radius of the Earth), then it can be calculated that the average density is 5.516 g/cm3. However, most rocks exposed at the surface have densities of less than 3g/cc, for example:

sandstone: 1.9 - 2.4 g/cm3
limestone: 1.9 - 2.7 g/cm3
granite: 2.6 - 2.7 g/cm3
basalt: 2.8 - 3.0 g/cm3 

Therefore, a material of greater density must exist at deeper levels within the Earth. The Earth has a series of layers or “shells”, but only the outer few km of the Earth can be directly observed; the upper crust, and the deepest boreholes which reach to only about 12.5 kms. Earthquakes provide the key to the structure at depth.


Stresses which develop in the Earth may become great enough to break the rocks, and cause slip along the resulting in fractures (faults). Although the slip distance in a given earthquake may be small (cm to metres), the rock masses involved are large and so the energy released is great. The resulting shock waves, or earthquakes, may cause great damage; greatest near the centre or focus, and less further away. The epicentre is the point on the surface of the Earth vertically above the focus.

Detection of seismic waves.
Earthquake energy is transmitted by several types of waves. Two types will be described:

P waves (primary or compressional) are transmitted by vibrations oscillating in the direction of propagation (push/pull).

S waves (secondary or shear), which vibrate at right angles to the direction of propagation. S waves cannot be transmitted through liquids because liquids have no elastic strength.

Recording Earthquakes

The arrival of earthquake waves is recorded by a seismograph. A mass is loosely coupled to the Earth by a spring. A chart is firmly coupled to the Earth. A pen linking them traces the difference in motion between the mass and the Earth’s surface. The arrival of waves from a distant earthquake is recorded as a seismogram on the rotating drum.

Consider what happens to P and S waves as they travel through the Earth.

The most important property of seismic waves is their speed of propagation. The velocity is governed by the physical properties (density, compressibility, rigidity) of the medium through which the wave is travelling.

Earlier in this lecture, it was deduced that the density of the Earth increases with depth. The wave propagation velocity must, therefore, change with depth, and this causes the wave to refract.


If a wave travelling through a medium with a fixed density encounters a new medium with a different density, the wave will change its direction. This “bending” of the wave is called refraction.

Data from seismometers located around the world can record waves from any given earthquake. The differences between recordings at different seismometers reveal properties of the sub-surface and hence the internal structure of the Earth.

For example, it has been discovered that the mantle is solid rock, but the outer core is a liquid. This was discovered, because for any given earthquake:-

  1. Both P and S waves are recorded by seismometers at distances of up to 103o from the epicentre.
  2. At distances greater than 103o, no S waves are recorded. This means that S waves that would have reappeared at > 103o have not propagated. The material at depths travelled by such waves must be liquid and be unable to transmit S waves.

Also, it has been discovered that the outer core must have a lower P wave velocity than the mantle. This is because at distances of 103o to 142o, no strong P waves are recorded. The liquid outer core has a lower P wave velocity, causing the P waves to be refracted to a steeper angle, so they cannot re-emerge between 103o to 142o. They actually re-emerge at angles > 186o. There is one small caveat to this observation. The inner core appears to be solid because some weak P wave arrivals occur between 103o to 142o. This is thought to be due to a slight increase in P wave velocity as waves enter the inner core, causing them to be refracted to a shallower angle, to re-emerge between 103o to 142o. If the inner core is solid, S waves could propa- gate there. The graph shows some calculations of what expected S wave velocities would be, but the inner core structure is still a source of controversy.


In the early 20th century a Yugoslavian seismologist by the name of Mohorovicic was studying seismograms from shallow focus earthquakes (< 40 km) that were nearby <800km. He noticed that there were 2 distinct sets of P waves and S waves involved. He interpreted these waves as a direct set and a refracted set. In the refracted set, waves travel down and are refracted at a boundary by a medium of higher velocity.

This boundary separates the crust with VP of 6-7km/sec from the upper mantle where VP starts at 8km/sec. It is called the Mohorovicic discontinuity but is commonly known as the MOHO.

Today, seismologists use artificial explosions to determine the structure beneath the surface and it is from these data that the depth of the MOHO can be calculated and thus the thickness of the crust. The MOHO is at 5-15 km under ocean crust and 35 km beneath normal thickness continental crust. The MOHO can be as much as 70 km deep beneath mountain belts where converging plates have caused an orogeny or mountain building event.

The Structure of the Earth

Recent advances in seismology now allow tomographic images of the interior of the Earth to be produced from P and S wave velocity data. Just as tomographic images of the interior of human bodies are produced by density contrasts in human tissue and bone subject to wave propagation, density contrasts in the Earth can be mapped by combining wave velocity data from large numbers of earthquakes.

The basic idea is that where the solid mantle is relatively hot, the P and S wave velocities should be anomalously low because the heat will result in a density decrease. One should be able to image hot, ascending plumes of mantle asthenosphere by looking for areas of anomalously low seismic velocity. Conversely, where the solid mantle is relatively cool, the P and S wave velocities should be anomalously fast because the lack of heat will result in a relatively high density.

One should be able to image cool, descending slabs of mantle lithosphere by looking for areas of anomalously high seismic velocity. Such images allow us to study subduction zones and constrain how deep the slabs penetrate. It appears that some slabs do not penetrate beneath 670 km whereas others continue down to the core-mantle boundary. This is an area of controversy in geology.