Carbonate Petrography

Carbonate petrography is the study of limestones, dolomites and associated deposits under optical or electron microscopes greatly enhances field studies or core observations and can provide a frame of reference for geochemical studies.

25 strangest Geologic Formations on Earth

The strangest formations on Earth.

What causes Earthquake?

Of these various reasons, faulting related to plate movements is by far the most significant. In other words, most earthquakes are due to slip on faults.

The Geologic Column

As stated earlier, no one locality on Earth provides a complete record of our planet’s history, because stratigraphic columns can contain unconformities. But by correlating rocks from locality to locality at millions of places around the world, geologists have pieced together a composite stratigraphic column, called the geologic column, that represents the entirety of Earth history.

Folds and Foliations

Geometry of Folds Imagine a carpet lying flat on the floor. Push on one end of the carpet, and it will wrinkle or contort into a series of wavelike curves. Stresses developed during mountain building can similarly warp or bend bedding and foliation (or other planar features) in rock. The result a curve in the shape of a rock layer is called a fold.

Showing posts with label igneous. Show all posts
Showing posts with label igneous. Show all posts

Geologic Contacts

A geologic contact is where one rock type touches another. There are three types of geologic contact:1. Depositional contacts are those where a sedimentary rock (or a lava flow) was deposited on an older rock
2. Intrusive contacts are those where one rock has intruded another
3. Fault contacts are those where rocks come into contact across fault zones.
Learn in detail about fault here

Following are the some pictures showing each type of geologic contact

Depositional Contacts

1. Angular Unconformity, Siccar Point, Scotland

This place is known as Siccar Point which is the most important unconformity described by James Hutton (1726-1797) in support of his world-changing ideas on the origin and age of the Earth.
Here 
gently sloping strata of 370-million-year-old Famennian Late Devonian Old Red Sandstone and a basal layer of conglomerate overlie near vertical layers of 435-million-year-old lower Silurian Llandovery Epoch greywacke, with an interval of around 65 million years.


2. Cretaceous Sandstone overlying Conglomerate    Kootenai Formation, SW Montana

Photo Courtesy: marlimillerphoto.com

3. Dun Briste Sea Stack, IrelandDun Briste is a truly incredible site to see but must be visited to appreciate its splendour. It was once joined to the mainland. The sea stack stands 45 metres (150 feet) tall.

Dun Briste and the surrounding cliffs were formed around 350 million years ago (during the 'Lower Carboniferous Period'), when sea temperatures were much higher and the coastline at a greater distance away.  There are many legends describing how the Sea Stack was formed but it is widely accepted that an arch leading to the rock collapsed during very rough sea conditions in 1393. This is remarkably recent in geological terms

Photo Courtesy: dunbriste.com 


Fault Contacts

1. Normal Faulting in the Cutler Formation near Arches National Park

Photo Courtesy: travelinggeologist.com

2. Normal Fault in Titus Canyon, Death Valley, California 

Photo Courtesy: travelinggeologist.com


3. 
Horst and Graben Structure in Zanjan, Iran

Photo Courtesy: Amazhda

Intrusive Contacts 

1. 
Pegmatite and aplite dikes and veins in granitic rocks on Kehoe Beach, Point Reyes National Seashore, California.


2. Spectacular mafic dyke from Isla de Socorro from Pep Cabré. The Isla de Socorro is a volcanic island off the west coast of Mexico and it is the only felsic volcano in the Pacific Ocean

Photo Courtesy: travelinggeologist.com

3. The margins of this Granite dyke cooled relatively quickly in contact with this much older Gabbro.
Photo near Ai-Ais Namibia

Photo Courtesy: travelinggeologist

10 of the Best Learning Geology Photos of 2016

A picture is worth a thousand words, but not all pictures are created equal. The pictures we usually feature on Learning Geology are field pictures showing Geological structures and features and many of them are high quality gem and mineral pictures. The purpose is to encourage students and professionals' activities by promoting "learning and scope" of Geology through our blogs.
In the end of 2016, we are sharing with you the 10 best photos of 2016 which we have posted on our page.

P.S: we always try our best to credit each and every photographer or website, but sometimes it’s impossible to track some of them. Please leave a comment if you know about the missing ones.

1. Folds from Basque France

 Image Credits: Yaqub ShahYaqub Shah

2. Horst and Graben Structure in Zanjan, Iran


Image Credits: https://www.instagram.com/amazhda



3. A unique Normal Fault

4. The Rock Cycle
The
 rock cycle illustrates the formation, alteration, destruction, and reformation of earth materials, and typically over long periods of geologic time. The rock cycle portrays the collective system of processes, and the resulting products that form, at or below the earth surface.The illustration below illustrates the rock cycle with the common names of rocks, minerals, and sediments associated with each group of earth materials: sediments, sedimentary rocks, metamorphic rocks, and igneous rocks.


Image Credits: Phil Stoffer


5. An amazing Botryoidal specimen for Goethite lovers! 


Image Credits: Moha Mezane 
   

6. Basalt outcrop of the Semail Ophiolite, Wadi Jizzi, Oman

Image Credits: Christopher Spencer
Christopher Spencer is founder of an amazing science outreach program named as Traveling Geologist. Visit his website to learn from him


7. Val Gardena Dolomites, Northern Italy





8. Beautiful fern fossil found in Potsville Formation from Pennsylvania.
The ferns most commonly found are Alethopteris, Neuropteris, Pecopteris, and Sphenophyllum.


Image Credits: Kurt Jaccoud


9. Snowball garnet in schist

Syn-kinematic crystals in which “Snowball garnet” with highly rotated spiral Si. 

Porphyroblast is ~ 5 mm in diameter.
From Yardley et al. (1990) Atlas of Metamorphic Rocks and their Textures.



10. Trilobite Specimen from Wheeler Formation, Utah
The Wheeler Shale is of Cambrian age and is a world famous locality for prolific trilobite remains. 


Image Credits: Paleo Fossils

How Do You Describe an Igneous Rock?

How Do You Describe an Igneous Rock? 

Different parameters are used to describe an igneous rock which are described in detail.

Characterizing Color and Texture 

If you wander around a city admiring building facades, you’ll find that many facades consist of igneous rock, for such rocks tend to be very durable. If you had to describe one of these rocks to a friend, what words might you use? You would  probably start by noting the rock’s colour. Overall, is the rock dark or light? More specifically, is it gray, pink, white, or black? Describing colour may not be easy, because some igneous rocks contain many visible mineral grains, each with a different colour; but even so, you’ll probably be able to characterize the overall hue of the rock. Generally, the colour reflects the rock’s composition, but it isn't always so simple, because colour may also be influenced by grain size and by the presence of trace amounts of impurities. (For example, the presence of a small amount of iron oxide gives rock a reddish tint.) Next, you would probably characterize the rock’s texture. A description of igneous texture indicates whether the rock consists of glass, crystals, or fragments. If the rock consists of crystals or fragments, a description of texture also specifies the grain size. Here are the common terms for defining texture:

How Do Extrusive and Intrusive Environments Differ?

How Do Extrusive and Intrusive Environments Differ? 

With a background on how melts form and freeze, we can now introduce key features of the two settings intrusive and extrusive in which igneous rocks form.

Extrusive Igneous Settings 

Different volcanoes extrude molten rock in different ways. Some volcanoes erupt streams of low-viscosity lava that flood down the flanks of the volcano and then cover broad swaths of the countryside. When this lava freezes, it forms a relatively thin lava flow. Such flows may cool in days to months. In contrast, some volcanoes erupt viscous masses of lava that pile into rubbly domes. And still others erupt explosively, sending clouds of volcanic ash and debris skyward, and/or avalanches of ash tumbling down the sides of the volcano.

Movement and Solidification of Molten Rock

Movement and Solidification of Molten Rock 

If magma stayed put once it formed, new igneous rocks would not develop in or on the crust. But it doesn't stay put; magma tends to move upward, away from where it formed. In some cases, it reaches the Earth’s surface and erupts at a volcano. This movement is a key component of the Earth System, because it transfers material from deeper parts of the Earth upward and provides the raw material from which new rocks and the atmosphere and ocean form. Eventually, magma freezes and transforms into a new solid rock.

Why Does Magma Rise? 

Magma rises for two reasons. First, buoyancy drives magma upward just as it drives a wooden block up through water, because magma is less dense than the surrounding rock. Second, magma rises because the weight of overlying rock creates pressure at depth that literally squeezes magma upward. The same process happens when you step into a puddle barefoot and mud squeezes up between your toes.

What Controls the Speed of Flow? 

Viscosity affects lava behaviour .
Viscosity, or resistance to flow, affects the speed with which magmas or lavas move. Magmas with low viscosity flow more easily than those with high viscosity, just as water flows more easily than molasses. Viscosity depends on temperature, volatile content, and silica content. Hotter magma is less viscous than cooler magma, just as hot tar is less viscous than cool tar, because thermal energy breaks bonds and allows atoms to move more easily. Similarly, magmas or lavas containing more volatiles are less viscous than dry (volatile-free) magmas, because the volatiles also tend to break apart silicate molecules and may also form gas bubbles. Mafic magmas are less viscous than felsic magmas, because silicon-oxygen tetrahedra tend to link together in magma to create long molecular chains that can’t move past each other easily, and there are more of these chains in a felsic magma than in a mafic magma. Thus, hotter mafic lavas have relatively low viscosity and flow in thin sheets over wide regions, but cooler felsic lavas are highly viscous and may clump into a dome-like mound at the volcanic vent (figure above a, b). 

Transforming Melt into Rock 

If a melt stayed at its point of origin, and nothing in its surroundings changed, it would stay molten. But melts don’t last forever. Rather, they eventually solidify or “freeze.” This process happens, in some cases, because volatiles escape from the melt, so that the freezing temperature rises if the melt’s temperature stays the same but its freezing temperature rises, it will solidify. Most often, however, freezing takes place simply when melt cools below its freezing temperature. Temperature decreases upward, toward the Earth’s surface, so magma enters a cooler environment automatically as it rises. If it is trapped underground as an intrusion, it slowly loses heat to surrounding wall rock, drops below its freezing temperature, and solidifies. If melt extrudes as lava at the ground surface, it cools in contact with air or water. 
The time it takes for a magma to cool depends on how fast it is able to transfer heat into its surroundings. To see why, think about the process of cooling coffee. If you pour hot coffee into a thermos bottle and seal it, the coffee stays hot for hours; because it’s insulated, the coffee in the thermos loses heat to the air outside only very slowly. Like the thermos bottle, surrounding wall rock acts as an insulator in that it transports heat away from a magma only very slowly, so magma underground (in an intrusive environment) cools slowly. In contrast, if you spill coffee on a table, it cools quickly because it loses heat to the cold air. Similarly, lava that erupts at the ground surface cools quickly because the air or water surrounding it can conduct heat away quickly. 
Three factors control the cooling time of magma that freezes below the surface in the intrusive realm. 

Factors that affect the freezing of molten rocks.
  • The depth of intrusion: Magma intruded deep in the crust, where it is surrounded by warm wall rock, cools more slowly than does magma intruded into cold wall rock near the ground surface. 
  • The shape and size of a magma body: Heat escapes from magma at an intrusion’s surface, so the greater the surface area for a given volume of intrusion, the faster it cools. Thus, a body of magma roughly with the shape of a pancake cools faster than one with the shape of a melon. And since the ratio of surface area to volume increases as size decreases, a body of magma the size of a car cools faster than one the size of a ship (figure above a, b). 
  • The presence of circulating groundwater: Water passing through magma absorbs and carries away heat, much like the coolant that flows around an automobile engine.

Changes in Magma during Cooling:  Fractional Crystallization 

Most people are familiar with the process of forming ice out of liquid water cool the water to a temperature of 0nC and crystals of ice start to form. Keep the temperature cold enough for long enough and all the water becomes solid, composed entirely of one type of mineral water ice. The process of freezing magma or lava is much more complex, because molten rock contains many different compounds, not just water, so during freezing of molten rock, many different minerals form. Further, not all of these minerals form at the same time (Bowen’s Reaction Series). To get a sense of this complexity, let’s look at an example.
When a mafic magma starts to freeze, mafic (iron- and magnesium-rich) minerals such as olivine and pyroxene start to crystallize first. These solid crystals are denser than the remaining magma, so they start to sink (figure above c). Some react chemically with the remaining magma as they sink, but some reach the floor of the magma chamber and become isolated from the magma. This process of sequential crystal formation and settling is called fractional crystallization it progressively extracts iron and magnesium from the magma, so the remaining magma becomes more felsic. If a magma freezes completely before much fractional crystallization has occurred, the magma becomes mafic igneous rock. But freezing of a magma that has been left over after lots of fractional crystallization has occurred produces felsic igneous rock.

Bowen’s Reaction Series

In the 1920s, the Canadian geologist Norman L. Bowen began a series of laboratory experiments designed to determine the sequence in which silicate minerals crystallize from a melt. First, Bowen melted powdered mafic igneous rock by raising its temperature to about 1280C. Then he cooled the melt just enough to cause part of it to solidify. Finally, he “quenched” the remaining melt by submerging it quickly in cold mercury. Quenching, which means sudden cooling to form a solid, transformed any remaining liquid into glass. The glass trapped the earlier-formed crystals within it. Bowen identified mineral crystals formed before quenching with a microscope, and he analyzed the chemical composition of the remaining glass. 

Bowen's reaction series indicates the succession of crystallization in cooling magma.
After experiments at different temperatures, Bowen found that, as new crystals form, they extract certain chemicals preferentially from the melt (figure above a). Thus, the chemical composition of the remaining melt progressively changes as the melt cools. Bowen described the specific sequence of mineral-producing reactions that take place in a cooling, initially mafic, magma. This sequence is now called Bowen’s reaction series in his honour.
Let’s examine the sequence more closely. In a cooling melt, olivine and calcium-rich plagioclase form first. This plagioclase reacts with the melt to form more, but different plagioclase; the plagioclase formed at a later stage contains more sodium (Na). Meanwhile, some olivine crystals react with the remaining melt to produce pyroxene, which may encase early olivine crystals or even replace them. However, some of the early olivine and Caplagioclase crystals settle out of the melt, taking iron, magnesium, and calcium atoms with them. By this process, the remaining melt becomes progressively enriched in silica. As the melt continues to cool, plagioclase continues to form, with later-formed plagioclase having progressively more sodium than earlier-formed plagioclase. Pyroxene crystals react with melt to form amphibole, and then amphibole reacts with the remaining melt to form biotite. All the while, crystals continue to settle out, so the remaining melt continues to become more felsic. At temperatures of between 650°C and 850°C, only about 10% melt remains, and this melt has a high silica content. At this stage, the final melt freezes, yielding quartz, K-feldspar (orthoclase), and muscovite. 
On the basis of his observations, Bowen realized that there are two tracks  to the reaction series. The “discontinuous” reaction series refers to the sequence  olivine, pyroxene, amphibole, biotite, K-feldspar-muscovite-quartz in that each step yields a different class of silicate mineral. The “continuous” reaction series refers to the progressive change from  calcium-rich to Na-rich plagioclase: the steps yield different versions of the  same mineral (figure above b). It’s important to note that not all minerals listed in the series appear in all igneous rock.  For example, a mafic magma may completely crystallize before felsic minerals such as quartz or K-feldspar have a chance to form.
Credits: Stephen Marshak (Essentials of Geology)

Nature of magma

Magma nature


Magma is a term first introduced into geologic literature in 1825 by Scope, who referred to it as a “compound liquid” consisting of solid particles suspended in a liquid, like mud. Measurements on extruded magma (lava), together with evaluations of mineral geothermometers in magmatic rocks and experimental determinations of their melting relations, indicate that temperatures of magmas near the surface of the Earth generally range from about 1200°C to 700°C; the higher values pertain to mafic compositions, the lower to silicic. Very rare alkali carbonatitic lavas that contain almost no silica have eruptive temperatures of about 600°C. Extruded magmas are rarely free of crystals, indicating that they rarely are superheated above temperatures of crystallization. Densities of magmas range from about 2.2 to 3.0 g/cm3 and are generally about 90% of that of the equivalent crystalline rock.
Magma in general consists of a mobile mixture of solid, liquid, and gaseous phases. The number and nature of the phases constituting a magma depend, under stable equilibrium conditions, on the three intensive variables P, T, and X (concentrations of chemical components in the magma). At sufficiently high T, any rock melts completely to form a homogeneous liquid solution, or melt. Except for carbonatite magmas, melts consist mostly of ions of O and Si hence the alternate appellation silicate liquid but always contain in addition significant amounts of Al, Ca, H, Na, and so on.
Examples of different types of magmas are shown schematically in below figure. Only in some unusually hot systems will a magma consist wholly of melt and no other phases. In most instances, a melt is only part of the whole magma, but is always present and gives it mobility. Hence, melt and magma are generally not the same. To a significant extent, the properties of the melt largely govern the overall dynamic behaviour of the whole magma. Rare magma systems consist at equilibrium of two physically distinct melts one essentially of carbonate and the other of silicate, or both are silicate but one is silicic and the other very rich in Fe. Each of these immiscible melts has distinctive properties, such as density. Oil and water are familiar immiscible liquids.
Schematic possible magmas

Atomic Structure of Melts

The configuration of ions in a melt its atomic structure largely dictates many of its significant properties. In pictorial representations of crystalline, liquid, and gaseous states, individual atoms have to be drawn as fixed in position relative to one another, but these are only their average, or instantaneous, positions. Even in crystals above absolute zero (0K), individual ions have motion. In glasses that are supercooled very viscous melts, ions experience vibrational motion: small periodic displacements about an average position. But at temperatures above a glass-melt transition, approximately two-thirds to three-quarters the melting T in degrees Kelvin, ions in the melt have more mobility and can break their bonds with neighbouring ions and wander about, forming new configurations. In a flowing melt, bonds are broken and bond angles and distances are distorted, but after deformation ceases, the ionic array may have sufficient time to reform into a “relaxed” equilibrium structure.
Many studies of melts in the laboratory using nuclear magnetic resonance, vibrational spectroscopy, and X-ray analyses reveal a lack of long-range (on the scale of more than a few atomic bond lengths) structural order and symmetry that characterize crystals. However, melts possess a short-range structural order in which tetrahedrally coordinated Si and Al cations are surrounded by four O anions and octahedrally bonded cations such as Ca and Fe2+ surrounded by six O anions roughly resemble those in crystals. Because silica is the most abundant constituent in most natural melts, the fundamental structural unit is the (SiO4)4 - tetrahedron, as it is in silicate minerals. Conceptual models of the atomic structure of silicate liquids can be constructed on the basis of these observations. Figure below depicts these models for liquid silica (SiO2) and CaMgSi2O6; the latter in crystalline form is diopside pyroxene.
Conceptual models of atomic structures of silicate melts compared with the symmetric lattice of a crystalline solid. (a) Crystalline silica (high tridymite). Layers of hexagonal rings of Si-O tetrahedra with alternating apices pointing up and down are stacked on top of one another, creating a three-dimensional structure in which each oxygen is shared by two silicons. Tetrahedra with apices pointing up have the upper apical oxygen left out of the drawing so as to reveal underlying silicon. Dashed line indicates outline of one unit cell in the lattice. (b) Model of liquid silica. Si-O tetrahedra are slightly distorted relative to the crystalline lattice. Long-range order is absent. Structure is highly polymerized because all tetrahedra are interconnected by bridging oxygen anions. (c) Model of liquid CaMgSi2O6 showing less polymerization than that of liquid silica. Note presence of network-modifying cations (Ca and Mg) and non-bridging oxygen, neither of which occurs in the silica melt. 
Because the entropy of melting of crystalline silica (i.e., the change in entropy from the crystalline to the liquid state) is relatively small, there can be little change in the degree of order in the atomic structure of the melt relative to the crystalline state. Thus, a model for liquid silica is a three-dimensional network of somewhat distorted Si-O tetrahedra, not unlike the corresponding structure of crystalline silica. Short range order is roughly similar to that in the crystalline state, but long-range order, as would be evident in a symmetrical crystal lattice, is absent. The silica melt can be viewed as a three-dimensional network of interlinked chains, or polymers, of Si-O tetrahedra.
On the other hand, in the model of the CaMgSi2O6 melt, these string like polymers are shorter, less intricately linked, and interspersed among octahedrally coordinated cations of Ca and Mg. This melt is not as polymerized as liquid silica.
Four different types of ions can be recognized in these models (Figure above) on the basis of their relation to the polymers: 
  1. Network-forming cations of Si4 within the interconnected tetrahedra of the polymers are strongly linked by 
  2. bridging oxygens. 
  3. Network-modifying cations of Ca and Mg are more weakly bonded to 
  4. non-bridging oxygens in non-tetrahedral bonding arrangements. 
The ratio of non-bridging oxygens to network-forming, tetrahedrally coordinated cations chiefly Si and Al is a measure of the degree of polymerization in a melt; small ratios correspond to high degrees of polymerization. In completely polymerized liquid silica, the ratio = 0. In partially polymerized liquid CaMgSi2O6 it is = 2/1 = 2.
The atomic structure of naturally occurring melts is more complex than these simple models. Despite considerable research, many details are not understood. Other ions of different size, charge, and electro-negativity, such as Al3+ , Ti4 +, Fe3+ , P5+ , H- , or F- make natural melts more complex. In this milieu, mobile cations compete for available anions, principally oxygen, in order to satisfy bonding requirements and to minimize the free energy of the melt. This is not quite the same situation as in crystals, where cations have more or less fixed sites of a particular coordination in the ordered lattice. In addition to the widespread (SiO4)4 - tetrahedra in melts, there are less abundant neighbouring tetrahedra of more negatively charged (Al3+ O4)5 - and (Fe3+ O4)5 -. The ionic charge and size of network-modifying cations, which generally form weaker bonds with non-bridging oxygens, can play an important role in melt structure. Network modifiers most commonly include monovalent K and Na; divalent Ca, Mg, Fe, and Mn, and more highly charged, but less abundant high-field-strength cations including P5+ , Ti4 +, and the still less abundant trace elements.
The most important dynamic property of a melt its viscosity depends strongly on its atomic structure. Because viscosity is a measure of the ease of flow of a melt and the mobility of ions, it should be intuitively obvious that more highly polymerized melts are more viscous. Alternatively, it can be said that, because non-bridging oxygen anions are less strongly bonded to neighbouring cations than bridging oxygens to Si and Al, viscosity correlates with the ratio of non-bridging to bridging oxygens. Increased concentrations of some components can depolymerize melts and reduce viscosity. Even small weight proportions of dissolved water or fluorine can depolymerize silicate melts, making them much less viscous. Also, high-field-strength, network-modifying cations whose charge is generally > 3+ have a strong affinity for oxygen anions and may successfully compete against network-forming Si4+ , Al3+ , and Fe3+ , thus depolymerizing the melt. The role of Fe in melt structures is especially significant because it occurs in two oxidation states. Fe2+ appears to be exclusively a network modifier, whereas Fe3+ can be either a network modifier or a network former. Changes in the oxidation state can therefore affect the degree of polymerization of a melt.
Increasing pressure appears to reduce the degree of polymerization somewhat. Because octahedral coordination of Si and Al is favoured in crystalline structures at high P over tetrahedral coordination, similar coordination changes might occur in melts at high P. Some experiments suggest that Al more readily shifts toward octahedral coordination with increasing P than does Si.
In conclusion, water-free (“dry”) rhyolite melts have virtually no non-bridging oxygens and are nearly completely polymerized and highly viscous. In andesite melts the ratio of non-bridging oxygens to network forming, tetrahedrally coordinated cations is about 0.2, and in basalt melts it is 0.4–1.2. Consequently, mafic, silica-poor melts are significantly less polymerized and less viscous than dry silicic melts.

Elementary concepts of thermodynamics

Elementary concepts of thermodynamics

Thermodynamic States, Processes, and State Variables

Geologic systems parts of the universe set aside in our mind for investigation are commonly more complex than, for example, the systems of a laboratory chemist. Geologic systems can be as large as the entire Earth and may endure for millions of years; they tend to be poorly definable and ever changing. Ideal end member systems in figure below are as follows:
  1. Isolated system: No matter or energy can be transferred across the boundary of the system and no work can be done on or by the system. Considering the span of time over which most identifiable geologic systems operate, no part of the Earth, nor even all of it, can be considered perfectly isolated because transfers of energy and movement of matter typify the dynamic Earth.
  2. Open system: The opposite of an isolated system. Matter and energy can flow across the boundary and work can be done on and by the system. Most geologic systems are open, at least in the context of their long lifetimes.
  3. Closed system: Energy, such as heat, can flow across the boundary of the system, but matter cannot. The composition of the system remains exactly constant. Because movement of matter is slow across boundaries of, for example, a rapidly cooling thin dike, it can be considered to be virtually closed. On the other hand, a large intrusion of magma that remains molten for tens of thousands of years while various fluids move across the wall rock contact is not a closed system.
  4. Adiabatic system: A special category of an isolated system is one in which no heat can be exchanged between the system and the surroundings. It is thermally insulated, but energy can be transferred across the system boundary through work done on or by it. Thus, an ascending, decompressing mantle plume or magma body cools as it expands against and does PV work on the surroundings, into which little heat is conducted because the rate of conduction is so slow.
End-member thermodynamic systems. Classification is according to flows of matter and energy across their boundary with the surroundings.
Any thermodynamic state of a system can be characterized by its state properties, or state variables.
Some state properties are extensive in that they depend on the amount of material, such as the mass and volume of some chemicals constituting a system. Other state properties are intensive and are independent of the amount of material present; they have a definite value at each point within the system, such as T, P, and concentration of a particular chemical species. Density (mass/volume) is also an intensive property. For example, at T 25°C and P 1 atm, the density of stable graphite anywhere within a crystal is 2.23 g/cm3 whereas that of diamond is 3.51 g/cm3.
P and T are extremely important intensive variables whose change in rock-forming systems is responsible for changes in their state; for example, increasing T causes solid rock to melt. Values of the geobaric and geothermal gradients in the Earth can be used to approximate P and T at a particular depth, assuming these gradients to be uniform. Otherwise uniform gradients can locally be perturbed, such as the geothermal gradient around a shallow crustal magma intrusion.
Rock-forming geologic processes can be thought of as thermodynamic processes: those that affect a thermodynamic system as it changes from one state to another.. In a thermodynamic process there could be all kinds of work done on or by the system, flows of heat in or out, chemical reactions and movement of matter, tortuous paths taken, and so forth. However, two ideal, end-member thermodynamic processes can be recognized: irreversible and reversible. In an irreversible thermodynamic process the initial state is metastable, and the spontaneous change in the system leads to a more stable, lower-energy final state. The conversion of metastable volcanic glass into more stable crystals a process called devitrification under near-atmospheric conditions is one example of an irreversible process. The energy trough in which metastable glass is stuck has to be surmounted, overcoming the activation energy, before it can move to a lower energy state. The atomic bonds in the glass must be broken or reformed into more stable crystalline structures. Devitrification of glass occurs spontaneously in the direction of diminishing energy, never in reverse (Figure below).
Spontaneous, irreversible thermodynamic process. A system in an initial higher-energy metastable state moves spontaneously to a lower-energy, stable state. The reverse never happens spontaneously.
In a reversible process, both initial and final states are stable equilibrium states and the path between them is a continuous sequence of equilibrium states. A reversible process is never actually realized in nature; it is a hypothetical concept that is used to make the mathematical models of thermodynamics work, and that’s all that can be said without an extended discussion.

First Law of Thermodynamics

Suppose that a change in the internal energy, dEi of a rock system, such as a mineral grain, is produced by adding some amount of heat to it, dq. As a result of absorbing heat, the mineral grain expands by an increment in volume, dV, doing an increment of PV work, dw PdV, on the surrounding mineral grains. According to the first law of thermodynamics, or law of conservation of energy, the increase in internal energy due to heat absorbed is diminished by the amount of work done on the surroundings, or
(1) dEi = dq - dw = dq - PdV
In the “scientific” convention, heat added to (or work done on) a system, as in this formulation, is positive and work done by (or heat withdrawn from) the system on its surroundings is negative.

Enthalpy

A new state property, the enthalpy, H, is defined here as
(2) H = Ei + PV
Upon differentiation, this equation becomes dH = dEi + PdV + VdP. Combining with equation (1) gives
(3) dH = dq + VdP
For isobaric (constant-pressure) changes, dP = 0, equation (3) becomes
(4) dHp = dqp
Although H may appear to be simply another and unnecessary label for q, the enthalpy is significant in providing a measure at constant P of how much heat is gained or lost from a system. Three possible changes in a system are accompanied by a gain or loss of heat, as follows:
  1. 1. Chemical reactions, such as the reaction of quartz plus calcite creating wollastonite plus CO2.
  2. 2. A change in state, such as crystals melting to liquid, that occurs at a fixed T once that T is reached in heating a system.
  3. 3. A change in T of the system where no change in state occurs, such as simply heating crystals below their melting T. 
The last two changes can be illustrated by a plot of enthalpy versus T for the diopside CaMgSi2O6 system as it is heated at constant P (Figure below). As diopside crystals absorb heat up to their melting point, T increases proportionally with the heat capacity, Cp. The slope of this line is (dH/dT)p Cp. At the melting T absorbed heat does not increase T, but is consumed in breaking the atomic bonds of the crystalline structure to produce the more random liquid array. This relatively large amount of absorbed heat at the constant T of melting, is the latent heat of melting, or enthalpy of melting, Hm. Once melting is complete, additional input of heat into the system raises the T of the melt proportional to its heat capacity. The slopes of the T-H lines above and below the
melting point are the heat capacities of the melt and the
crystals, respectively.
Note in Figure below that the latent heat involved in changing the state of the system from liquid to crystal, or vice versa, is similar to the heat absorbed in changing the T of the crystals or liquid by hundreds of degrees. Thus, melting of rock in the Earth to generate magma absorbs a vast amount of thermal energy, which moderates changes in T in the system.
Enthalpy-T relations for CaMgSi2O6 at 1 atm. 
They are either exothermic or endothermic, respectively. If catalyzed by a spark, the reaction between hydrogen and oxygen releases a burst of heat a rapid exothermic reaction, or explosion. Solidification of magma exothermically releases the latent heat of crystallization. In contrast, dissolution of potassium nitrate in water endothermically absorbs heat so that the container becomes cold. Thus, whether heat is released or absorbed in a chemical reaction provides no consistent clue as to the direction a reaction moves spontaneously.

Entropy and the Second and Third Laws of Thermodynamics

Another way to look at spontaneity is in changes in the distribution or concentration of energy. Spontaneous thermal processes lead to a more even concentration of heat. A bowl of hot soup on the table eventually reaches the same T as the room. Heat flows spontaneously from a hot body to a cold, eliminating the difference in T. Without an uneven concentration of thermal energy, or the opportunity for heat flow, no work can be done. As heat flows from an intrusion of hotter magma into the cooler wall rocks, PV work of volumetric expansion is performed on them. The heat also drives endothermic chemical reactions of wall rock metamorphism. Water in an enclosed lake on a high plateau has gravitational potential energy relative to sea level, but so long as it is isolated from sea level and cannot flow, the concentration of energy is uniform and no work can be done. However, in the natural course of events, a river drains the lake into the sea, forming a process path along the potential energy gradient between the high- and low potential-energy levels. Work can then be done, driving turbines to generate electricity, eroding the river channel, transporting sediment, and so on.
One statement of the second law of thermodynamics is that spontaneous natural processes tend to even out the concentration of some form of energy, smoothing the energy gradient. A hot lava flow extruded from a lofty volcano cools to atmospheric T as it descends down slope, thereby reducing differences in thermal and gravitational potential energy between initial and final states in accordance with the second law.
Eventually, billions of years from now, all of the thermal energy in the Earth will be consumed in tectonism, volcanism, and other processes and dispersed into outer space. No mountains or volcanoes will be erected and erosion in the solar-powered hydrologic system will wear everything down to some common level (assuming the Sun does not run out of nuclear energy!). Without differences in the concentration of thermal and gravitational potential energy no geologic work can be accomplished and the planet will be geologically dead!. 
The measure of the uniformity in concentration of energy in a system is called the entropy, S. The more uniform the concentration of some form of energy, the greater the entropy. The geologically dead planet will have maximal entropy.
Another, more useful, way to define entropy is to relate it to the internal disorder in the system. This provides an alternate statement of the second law: In any spontaneous process in an isolated system there is an increase in entropy, that is, an increase in disorder. The law in this form is illustrated by Figure 3.4, where white and black balls in the boxes represent molecules of two different gases. The spontaneous mixing of the two gas molecules results in an increase in “mixedupness,” disorder, randomness, or entropy. Note that there is no accompanying change in energy in this mixing process. Thus, another driving “force” for a spontaneous process is an increase in entropy, even though there may be no change in the energy.
At decreasing T, crystals become increasingly ordered, less atomic substitution is possible, and their entropy decreases. The third law of thermodynamics states that at absolute zero, where the Kelvin temperature is zero (0K = - 273.15°C), crystals are perfectly ordered and all atoms are fixed in space so that the entropy is zero.
A convenient way to think of relative entropies is that a gas made of high-speed molecules in random trajectories has a greater entropy than the compositionally equivalent liquid array, which, though still somewhat disordered, has linked atoms. The compositionally equivalent crystalline solid has still lower entropy, because its atoms form an ordered array. As an example, for water,
Ssteam > Sliquid water > Sice

Gibbs Free Energy

The boulder-on-the-hill example has two major flaws as an analogy for the way natural systems, in general, change spontaneously from a higher to a lower energy state. First, it isn’t always gravitational potential energy that is minimized. Second, the analogy does not take into account the fact that in an isolated system a process can proceed spontaneously without any change in energy, but it does proceed with increasing entropy figure below. To overcome these two flaws, a new extensive property of a system is defined in such a way as to serve as a universal directionality pointer for spontaneous reactions. 
Entropy increases in a spontaneous, irreversible process in an isolated system. Bottom left, a hypothetical isolated system a box filled with atoms of two gases (black and white balls) separated by an impermeable wall. Bottom right, the wall has been removed in the box, and the atoms of the two gases have mixed spontaneously and irreversibly as a result of their motion (kinetic energy). An increase in disorder or randomness of the atoms in the system and an increase in entropy, S > 0. No change in energy has occurred.

This new property is called the Gibbs free energy, G, and is defined by the expression
G = H + TS. Combining with equation (2)
(6) G = Ei + PV - TS
In differential form this becomes
(7) dG = dEi + PdV + VdP - TdS - SdT
Remembering the work-pressure-volume relation (dW = PdV), we can write a parallel expression for the heat temperature-entropy relation
(8) dq = TdS
Equation 97) can be simplified by substituting this equality and also by making a substitution from equation (2) to obtain
(9) dG = VdP - SdT
Equation (9) is a useful thermodynamic expression that allows us to make powerful statements regarding the direction of changes in geologic systems as the independent intensive variables of state, T and P, change. The extensive entropy and volume properties of the system are also relevant factors.
If P and T remain the same through any spontaneous change in state, that is, dP = dT = 0, then, from
equation (9)
(10) dGP,T = 0
This is simply the condition for a minimum (or maximum) in G in P-T space where the slope of the tangent to the energy function is zero, or horizontal. Figure below shows a system that has moved to a state of minimal energy and stable equilibrium from a higher-energy, metastable state at constant P and T in a closed (constant composition) system. Note that the energy change,  GP,T, between the initial metastable state and the final stable state is negative: GP,T < 0.
In some spontaneously changing systems, increasing entropy is the dominant factor, whereas in  thers, decreasing energy is the dominant factor. The Gibbs free energy is formulated in such a way that it always decreases in a spontaneous change in the state of a system.
In other words, the Gibbs free energy of the final stable state is lower than that of the initial metastable state in figure below. The Gibbs free energy is a thermodynamic potential energy that, like gravitational potential energy for the hypothetical boulders, is the lowest possible in a state of stable equilibrium for a changing system.

Gibbs free energy decreases in a spontaneous change in a closed system where the initial and final states are at the same P and T. In this example, the energy of diamond is greater than that of graphite at the same P and T, or Gdiamond > Ggraphite, so the change in energy, GP,T, in the spontaneous process is negative, or Gdiamond - Ggraphite = GP,T < 0. Note the activation energy barrier, Ea , that must be surmounted in order for the change to occur.


Igneous rocks

Characterizing Color and Texture 

If you wander around a city admiring building façades, you'll find that many façades consist of igneous rock, for such rocks tend to be very durable. If you had to describe one of these rocks to a friend, what words might you use? You would  probably start by noting the rock’s colour. Overall, is the rock dark or light? More specifically, is it Gray, pink, white, or black? Describing colour may not be easy, because some igneous rocks contain many visible mineral grains, each with a different colour; but even so, you’ll probably be able to characterize the overall hue of the rock. Generally, the colour reflects the rock’s composition, but it isn't always so simple, because colour may also be influenced by grain size and by the presence of trace amounts of impurities. (For example, the presence of a small amount of iron oxide gives rock a reddish tint.) Next, you would probably characterize the rock’s texture. A description of igneous texture indicates whether the rock consists of glass, crystals, or fragments. If the rock consists of crystals or fragments, a description of texture also specifies the grain size. Here are the common terms for defining texture:
Textures and types of igneous rocks.
  1. Crystalline texture: Rocks that consist of minerals that grow when a melt solidifies interlock like pieces of a jigsaw puzzle (a in figure above). Rocks with such a texture are called crystalline igneous rocks. The interlocking of crystals in these rocks occurs because once some grains have developed, they interfere with the growth of  later-formed grains. The last grains to form end up filling irregular spaces between already existing grains. Geologists distinguish subcategories of crystalline igneous rocks according to the size of the crystals. Coarse-grained (phaneritic) rocks have crystals large enough to be identified with the naked eye. Fine-grained (aphanitic) rocks have crystals too small to be identified with the naked eye. Porphyritic rocks have larger crystals surrounded by a mass of fine crystals. In a porphyritic rock, the larger crystals are called phenocrysts, while the mass of finer crystals is called ground mass. 
  2. Fragmental texture: Rocks consisting of igneous chunks and/ or shards that are packed together, welded together, or cemented together after having solidified are fragmental igneous rocks (a in figure above). 
  3. Glassy texture: Rocks made of a solid mass of glass, or of tiny crystals surrounded by glass, are glassy igneous rocks. Glassy rocks fracture conchoidally (b in figure above). 
What factors control the texture of igneous rocks? In the case of non-fragmental rocks, texture largely reflects cooling rate. The presence of glass indicates that cooling happened so quickly that the atoms within a lava didn't have time to arrange into crystal lattices. Crystalline rocks form when a melt cools more slowly. In crystalline rocks, grain size depends on cooling time. A melt that cools rapidly, but not rapidly enough to make glass, forms fine-grained rock, because many crystals form but none has time to grow large (c figure above). A melt that cools very slowly forms a coarse-grained rock, because a few crystals have time to grow large.
Because of the relationship between cooling rate and texture, lava flows, dikes, and sills tend to be composed of fine grained igneous rock. In contrast, plutons tend to be composed of coarse-grained rock. Plutons that intrude into hot wall rock at great depth cool very slowly and thus tend to have larger crystals than plutons that intrude into cool country rock at shallow depth, where they cool relatively rapidly. Porphyritic rocks form when a melt cools in two stages. First, the melt cools slowly at depth, so that phenocrysts form. Then, the melt erupts and the remainder cools quickly, so that groundmass crystallizes around the phenocrysts.
There is, however, an exception to the standard cooling rate and grain size relationship. A very coarse-grained igneous rock called pegmatite doesn't necessarily cool slowly. Pegmatite contains crystals up to tens of centimetres across and occurs in dikes. Because pegmatite occurs in dikes, which generally cool quickly, the coarseness of the rock may seem surprising. Researchers have shown that pegmatites are coarse because they form from water-rich melts in which atoms can move around so rapidly that large crystals can grow very quickly.

Classifying Igneous Rocks 

Because melts can have a variety of compositions and can freeze to form igneous rocks in many different environments above and below the surface of the Earth, we observe a wide spectrum of igneous rock types. We classify these according to their texture and composition. Studying a rock’s texture tells us about the rate at which it cooled, as we've seen, and therefore the environment in which it formed. Studying its composition tells us about the original source of the magma and the way in which the magma evolved before finally solidifying. Below, we introduce some of the more important igneous rock types. 

Crystalline igneous rocks

Igneous rocks are classified based on composition and texture.
The scheme for classifying the principal types of crystalline igneous rocks is quite simple. The different compositional classes are distinguished on the basis of silica content ultramafic, mafic, intermediate, or felsic whereas the different textural classes are distinguished according to whether the grains are coarse or fine.  The chart in figure above gives the texture and composition of the most commonly used crystalline igneous rock names. As a rough guide, the colour of an igneous rock reflects its composition: mafic rocks tend to be black or dark Gray, intermediate rocks tend to be lighter Gray or greenish Gray, and felsic rocks tend to be light tan to pink or maroon. Note that rhyolite and granite have the same chemical composition but differ in grain size. Which of these two rocks develops from a melt of felsic composition depends on the cooling rate. A felsic lava that solidifies quickly at the Earth’s surface or in a thin dike or sill turns into fine-grained rhyolite; but the same magma, if solidifying slowly at depth in a pluton, turns into coarse-grained granite. A similar situation holds for mafic lavas a mafic lava that cools quickly in a lava flow forms basalt, but a mafic magma that cools slowly forms gabbro. 

Glassy igneous rocks

Glassy texture develops more commonly in felsic igneous rocks because the high concentration of silica inhibits the easy growth of crystals. But basaltic and intermediate lavas can form glass if they cool rapidly enough. In some cases, a rapidly cooling lava freezes while it still contains a high concentration of gas bubbles these bubbles remain as open holes known as vesicles. Geologists distinguish among several different kinds of glassy rocks.
Pumice, a vesicle-filled volcanic rock, is so light that paper can hold it up. The vesicles it contains tend to be small.
  • Obsidian is a mass of solid, felsic glass. It tends to be black or brown (b in first figure). Because it breaks conchoidally, sharp-edged pieces split off its surface when you hit a sample with a hammer. Pre- industrial people worldwide used such pieces for arrowheads, scrapers, and knife blades. 
  • Pumice is a felsic volcanic rock that contains  abundant vesicles, giving it the appearance of a sponge. Pumice forms by the quick cooling of frothy lava that  resembles the head of foam in a glass of beer. In some cases, pumice contains so many air-filled pores that it can actually float on water, like styrofoam (figure above). 
  • Scoria is a mafic volcanic rock that contains abundant vesicles (more than about 30%). Generally, the bubbles in scoria are bigger than those in pumice, and the rock, overall, looks darker.

Pyroclastic igneous rocks  

When volcanoes erupt explosively, they spew out fragments of lava. Geologists refer to all such fragments as pyroclasts. Accumulations of fragmental volcanic debris are called pyroclastic deposits, and when the material in these deposits consolidates into a solid mass, due either to welding together of still-hot clasts or to cementation by minerals precipitating from water passing through, it becomes a pyroclastic rock. Geologists distinguish among several types of pyroclastic rocks based on grain size. Let’s consider two examples. 
  • Tuff is a fine-grained pyroclastic igneous rock composed of volcanic ash. It may contain fragments of pumice. 
  • Volcanic breccia consists of larger fragments of volcanic debris that either fall through the air and accumulate, or form when a lava flow breaks into pieces.
Credits: Stephen Marshak (Essentials of Geology)

    Metals and Ores

    Metal and Ores

    Metal and Its Discovery 

    Metals are opaque, shiny, smooth solids that can conduct electricity and can be bent, drawn into wire, or hammered into thin sheets. In this regard, they look and behave quite differently from wood, plastic, meat, or rock. This is because, unlike in other substances, the atoms that make up metals are held together by metallic bonds, so electrons can flow from atom to atom fairly easily and atoms can, in effect, slide past each other without breaking apart. The first metals that people used copper, silver, and gold can occur in rock as native metals. Native metals consist only of metal atoms, and thus look and behave like metal.
    Gold occurs as native metal within quartz veins. The quartz breaks up to form sand, leaving nuggets of gold.
    Gold nuggets, for example, are chunks of native metal that have eroded free of bedrock (figure above). Over the ages, people have collected nuggets of native metal from stream beds and pounded them together with stone hammers to make arrowheads, scrapers, and later, coins and jewellery. But if we had to rely solely on native metals as our source of metal, we would have access to only a tiny fraction of our current metal supply. Most of the metal atoms we use today originated as ions bonded to non-metallic elements in a great variety of minerals that themselves look nothing like metal. Only because of the chance discovery by some prehistoric genius that certain rocks, when heated to high temperatures in fire (a process called smelting), decompose to yield metal plus a non-metallic residue called slag, do we now have the ability to produce sufficient metal for the needs of industrialized society. 

    What Is an Ore? 

    Examples of ore minerals.
    The minerals from which metals can be extracted are called ore minerals, or economic minerals. These minerals contain metal in high concentrations and in a form that can be easily extracted. Galena (PbS), for example, is about 50% lead, so we consider it to be an ore mineral of lead (a in figure above). We obtain most of our iron from haematite and magnetite. Copper comes from a variety of minerals, none of which look like copper (b in figure above). Geologists have identified a great variety of ore minerals. Many ore minerals are sulphides, in which the metal occurs in combination with sulphur (S), or oxides, in which the metal occurs in combination with oxygen (O).
    To obtain the metals needed for industrialized society, we mine ore, rock containing native metals or a concentrated accumulation of ore minerals. To be an ore, rock must not only contain ore minerals, it must also contain a sufficient amount to make the rock worth mining. Iron constitutes only about 6.2% of the continental crust's weight but makes up about 30% to 60% of iron ore. The concentration of a useful metal in an ore determines the grade of the ore the higher the concentration, the higher the grade. Whether or not an ore of a given grade is worth mining depends on the price of metal in the market. 

    How Do Ore Deposits Form? 

    Ore minerals do not occur uniformly through rocks of the crust. If they did, we would not be able to extract them economically. Fortunately for humanity, geologic processes concentrate these minerals into accumulations called ore deposits. Simply put, an ore deposit is an economically significant occurrence of ore. The various kinds of ore deposits differ from each other in terms of which ore minerals they contain and which geologic conditions led to their formation. Below, we introduce a few examples.
    Various processes that form ore deposits.

    Magmatic deposits 

    When a magma cools, sulphide ore minerals crystallize early, then, because sulphides tend to be dense, hey sink to the bottom of the magma chamber, where they accumulate; this accumulation is a magmatic deposit. When the magma freezes solid, the resulting igneous body may contain a concentration of sulphide minerals at its base. Because of their composition, such concentrations are known as “massive- sulphide deposits” (a in figure above).

    Hydrothermal deposits 

    Hydrothermal activity involves the circulation of hot-water solutions through a magma or through the rocks surrounding an igneous intrusion. These fluids dissolve metal ions. When a solution enters a region of lower pressure, lower temperature, different acidity, and/ or different availability of oxygen, the metals come out of solution and form ore minerals that precipitate in fractures and pores, creating a hydrothermal deposit (figure above b). Such deposits may form within an igneous intrusion or in surrounding country rock. If the resulting ore minerals disperse through the intrusion, we can also call the deposit a disseminated deposit, but if they precipitate to fill cracks in pre-existing rock, we can call the deposit a vein deposit; veins are mineral-filled cracks.
    In recent decades, geologists have discovered that hydrothermal activity at the submarine volcanoes along mid-ocean ridges leads to the eruption of hot water, containing high concentrations of dissolved metal and sulphur, from a vent. When this hot water comes in contact with cold seawater, the dissolved components instantly precipitate as tiny crystals of metal-sulphide minerals (c in figure above). The erupting water, therefore, looks like a black cloud, so the vents are called “black smokers”. The minerals in the cloud eventually sink and form a pile of ore minerals around the vent. Since the ore minerals typically are sulphides, the resulting hydrothermal deposits constitute another type of massive-sulphide deposit.

    Secondary-enrichment deposits

    Sometimes groundwater passes through ore-bearing rock long after the rock first formed. This groundwater dissolves some of the ore minerals and carries the dissolved ions away. When the water eventually flows into a different chemical environment (for instance, one with a different amount of oxygen or acid), it precipitates new ore minerals, commonly in concentrations exceeding that of the original deposit. A new ore deposit formed from metals that were dissolved and carried away from a pre-existing ore deposit is called a secondary-enrichment deposit. Some of these deposits contain spectacularly beautiful copper-bearing carbonate minerals, such as azurite and malachite. 

    MVT ores 


    Rain falling along one margin of a large sedimentary basin may sink into the subsurface and then flow as groundwater along a curving path that takes it first down to the bottom of the basin, and then eventually back up to the opposite margin of the basin, hundreds of kilometres away. At the bottom of the basin, temperatures become high enough that the water dissolves metals. As the water returns to the surface and enters cooler rock, these metals precipitate in ore minerals. Ore deposits formed in this way, containing lead- and zinc-bearing minerals, appear in dolomite beds of the Mississippi Valley region of the United States, and thus have come to be known as Mississippi Valley–type (MVT) ores. 

    Sedimentary deposits of metals 

    A Precambrian banded iron formation from northern Michigan.
    Some ore minerals accumulate in sedimentary environments under special circumstances. For example, between 2.5 and 1.8 billion years ago, the atmosphere, which previously had contained very little oxygen, gained oxygen because of the evolution of abundant photo synthetic organisms. This change affected the chemistry of seawater so that large quantities of dissolved iron precipitated as iron oxide minerals that settled as sediment on the sea floor. The resulting iron-rich sedimentary deposits are known as a banded iron formation (BIF) (figure above), because after lithification they consist of alternating beds of Gray iron oxide (magnetite or haematite) and red beds of jasper (iron-rich chert).
    The chemistry of seawater in some parts of the ocean today leads to the deposition of manganese-oxide minerals on the sea floor. These minerals grow into lumpy accumulations known as manganese nodules. Mining companies have begun to explore technologies for vacuuming up these nodules; geoscientists estimate that the worldwide supply of nodules contains 720 years’ worth of copper and 60,000 years’ worth of manganese, at current rates of consumption. 

    Residual mineral deposits 

    Recall from Interlude B that as rainwater sinks into the Earth, it leaches (dissolves) certain elements and leaves behind others, as part of the process of forming soil. In rainy, tropical environments, the residue left behind in soils after leaching includes concentrations of iron or aluminium. Locally, these metals become so concentrated that the soil itself becomes an ore deposit. We refer to such deposits as residual mineral deposits. Most of the aluminium ore mined today comes from bauxite, a residual mineral deposit created by the extreme leaching of rocks (such as granite) containing aluminium-bearing minerals.  The figure below shows the residual mineral deposits as well as placer deposits.

    Placer deposits 

    Placer deposits form where erosion produces clasts of native metals. Sorting by the stream concentrates the metals.
    Ore deposits may develop when rocks containing native metals erode, producing a mixture of sand grains and metal flakes or nuggets (pebble-sized fragments). For example, gold accumulates in sand or gravel bars along the course of rivers, for the moving water carries away lighter mineral grains (quartz and feldspar) but can’t move the heavy metal grains (gold) so easily. Concentrations of metal grains in stream sediments are a type of placer deposit (figure above). Panning further concentrates gold flakes or nuggets swirling water in a pan causes the lighter sand grains to wash away, leaving the gold behind.

    Where Are Ore Deposits Found? 

    The Inca Empire of fifteenth-century Peru boasted elaborate cities and temples, decorated with fantastic masks, jewellery, and sculptures made of gold. Then, around 1532, Spanish ships arrived, led by conquistadors who quipped, “We Spaniards suffer from a disease that only gold can cure.” The Incas, already weakened by civil war, were no match for the armorclad Spaniards with their guns and horses. Within six years, the Inca Empire had vanished, and Spanish ships were transporting Inca treasure back to Spain. Why did the Incas possess so much gold? Or to ask the broader question, what geologic factors control the distribution of ore? Once again, we can find the answer by considering the consequences of plate tectonics. Several of the ore-deposit types mentioned above occur in association with igneous rocks. Igneous activity does not happen randomly around the Earth, but rather concentrates along convergent plate boundaries (specifically, in the overriding plate of a subduction zone), along divergent plate boundaries (along mid-ocean ridges), continental rifts, or hot spots. Thus, magmatic and hydrothermal deposits (and secondary-enrichment deposits derived from these) occur in these geologic settings. Placer deposits are typically found in the sediments eroded from such magmatic or hydrothermal deposits. The Inca gold formed in the Andes along a convergent plate boundary.

    Credits: Stephen Marshak (Essentials of Geology)