Petrology

A volcanic sand grain seen under the microscope, with plane-polarized light in the upper picture,

and cross polarized light in the lower picture. Scale box is 0.25 mm.

Petrology (from Greek: πέτρα, petra, rock; and λόγος, logos, knowledge) is the branch of

geology that studies rocks, and the conditions in which rocks form.

Lithology was once approximately synonymous with petrography, but in current usage, lithology

focusses on macroscopic hand-sample or outcrop-scale description of rocks, while petrography is

the speciality that deals with microscopic details.

In the oil industry, lithology, or more specifically mud logging, is the graphic representation of

geological formations being drilled through, and drawn on a log called a mud log. As the cuttings

are circulated out of the borehole they are sampled, examined (typically under a 10x microscope)

and tested chemically when needed.

Methodology

Petrology utilizes the classical fields of mineralogy, petrography, optical mineralogy, and

chemical analyses to describe the composition and texture of rocks. Modern petrologists also

include the principles of geochemistry and geophysics through the studies of geochemical trends

and cycles and the use of thermodynamic data and experiments to better understand the origins

of rocks.

Branches

There are three branches of petrology, corresponding to the three types of rocks: igneous,

metamorphic, and sedimentary, and another dealing with experimental techniques:

Igneous petrology focuses on the composition and texture of igneous rocks (rocks such as

granite or basalt which have crystallized from molten rock or magma). Igneous rocks

include volcanic and plutonic rocks.

Sedimentary petrology focuses on the composition and texture of sedimentary rocks

(rocks such as sandstone, shale, or limestone which consist of pieces or particles derived

from other rocks or biological or chemical deposits, and are usually bound together in a

matrix of finer material).

Metamorphic petrology focuses on the composition and texture of metamorphic rocks

(rocks such as slate, marble, gneiss, or schist which started out as sedimentary or igneous

rocks but which have undergone chemical, mineralogical or textural changes due to

extremes of pressure, temperature or both)

Experimental petrology employs high-pressure, high-temperature apparatus to investigate

the geochemistry and phase relations of natural or synthetic materials at elevated

pressures and temperatures. Experiments are particularly useful for investigating rocks of

the lower crust and upper mantle that rarely survive the journey to the surface in pristine

condition. The work of experimental petrologists has laid a foundation on which modern

understanding of igneous and metamorphic processes has been built.

Igneous rock

Geologic provinces of the world (USGS) Shield Platform Orogen Basin Large igneous

province Extended crust Oceanic crust: 0–20 Ma 20–65

Ma >65 Ma

Volcanic rock in North America.

Plutonic rock in North America.

Igneous rock (derived from the Latin word igneus meaning of fire, from ignis meaning fire) is

one of the three main rock types, the others being sedimentary and metamorphic rock. Igneous

rock is formed through the cooling and solidification of magma or lava. Igneous rock may form

with or without crystallization, either below the surface as intrusive (plutonic) rocks or on the

surface as extrusive (volcanic) rocks. This magma can be derived from partial melts of pre-

existing rocks in either a planet's mantle or crust. Typically, the melting is caused by one or more

of three processes: an increase in temperature, a decrease in pressure, or a change in

composition. Over 700 types of igneous rocks have been described, most of them having formed

beneath the surface of Earth's crust. These have diverse properties, depending on their

composition and how they were formed.

Geological significance

The upper 16 kilometres (10 mi) of Earth's crust is composed of approximately 95% igneous

rocks with only a thin, widespread covering of sedimentary and metamorphic rocks.[1]

Igneous rocks are geologically important because:

which some igneous rocks are extracted, and the temperature and pressure conditions that

allowed this extraction, and/or of other pre-existing rock that melted;

their absolute ages can be obtained from various forms of radiometric dating and thus can

be compared to adjacent geological strata, allowing a time sequence of events;

their features are usually characteristic of a specific tectonic environment, allowing

tectonic reconstitutions (see plate tectonics);

in some special circumstances they host important mineral deposits (ores): for example,

tungsten, tin, and uranium are commonly associated with granites and diorites, whereas

ores of chromium and platinum are commonly associated with gabbros.

Morphology and setting

In terms of modes of occurrence, igneous rocks can be either intrusive (plutonic), extrusive

(volcanic) or hypabyssal.

Intrusive igneous rocks

Close-up of granite (an intrusive igneous rock) exposed in Chennai, India.

Intrusive igneous rocks are formed from magma that cools and solidifies within the crust of a

planet. Surrounded by pre-existing rock (called country rock), the magma cools slowly, and as a

result these rocks are coarse grained. The mineral grains in such rocks can generally be identified

with the naked eye. Intrusive rocks can also be classified according to the shape and size of the

intrusive body and its relation to the other formations into which it intrudes. Typical intrusive

formations are batholiths, stocks, laccoliths, sills and dikes.

The central cores of major mountain ranges consist of intrusive igneous rocks, usually granite.

When exposed by erosion, these cores (called batholiths) may occupy huge areas of the Earth's

surface.

Coarse grained intrusive igneous rocks which form at depth within the crust are termed as

abyssal; intrusive igneous rocks which form near the surface are termed hypabyssal.

Extrusive igneous rocks

Basalt (an extrusive igneous rock in this case); light coloured tracks show the direction of lava

flow.

Extrusive igneous rocks are formed at the crust's surface as a result of the partial melting of rocks

within the mantle and crust. Extrusive Igneous rocks cool and solidify quicker than intrusive

igneous rocks. Since the rocks cool very quickly they are fine grained.

The melted rock, with or without suspended crystals and gas bubbles, is called magma. Magma

rises because it is less dense than the rock from which it was created. When it reaches the

surface, magma extruded onto the surface either beneath water or air, is called lava. Eruptions of

volcanoes into air are termed subaerial whereas those occurring underneath the ocean are termed

submarine. Black smokers and mid-ocean ridge basalt are examples of submarine volcanic

activity.

The volume of extrusive rock erupted annually by volcanoes varies with plate tectonic setting.

Extrusive rock is produced in the following proportions:[2]

divergent boundary: 73%

convergent boundary (subduction zone): 15%

hotspot: 12%.

Magma which erupts from a volcano behaves according to its viscosity, determined by

temperature, composition, and crystal content. High-temperature magma, most of which is

basaltic in composition, behaves in a manner similar to thick oil and, as it cools, treacle. Long,

thin basalt flows with pahoehoe surfaces are common. Intermediate composition magma such as

andesite tends to form cinder cones of intermingled ash, tuff and lava, and may have viscosity

similar to thick, cold molasses or even rubber when erupted. Felsic magma such as rhyolite is

usually erupted at low temperature and is up to 10,000 times as viscous as basalt. Volcanoes with

rhyolitic magma commonly erupt explosively, and rhyolitic lava flows typically are of limited

extent and have steep margins, because the magma is so viscous.

Felsic and intermediate magmas that erupt often do so violently, with explosions driven by

release of dissolved gases — typically water but also carbon dioxide. Explosively erupted

pyroclastic material is called tephra and includes tuff, agglomerate and ignimbrite. Fine volcanic

ash is also erupted and forms ash tuff deposits which can often cover vast areas.

Because lava cools and crystallizes rapidly, it is fine grained. If the cooling has been so rapid as

to prevent the formation of even small crystals after extrusion, the resulting rock may be mostly

glass (such as the rock obsidian). If the cooling of the lava happened slowly, the rocks would be

coarse-grained.

Because the minerals are mostly fine-grained, it is much more difficult to distinguish between

the different types of extrusive igneous rocks than between different types of intrusive igneous

rocks. Generally, the mineral constituents of fine-grained extrusive igneous rocks can only be

determined by examination of thin sections of the rock under a microscope, so only an

approximate classification can usually be made in the field.

Hypabyssal igneous rocks

Hypabyssal igneous rocks are formed at a depth in between the plutonic and volcanic rocks.

Hypabyssal rocks are less common than plutonic or volcanic rocks and do often form dikes, sills

or laccoliths.

Classification

Igneous rocks are classified according to mode of occurrence, texture, mineralogy, chemical

composition, and the geometry of the igneous body.

The classification of the many types of different igneous rocks can provide us with important

information about the conditions under which they formed. Two important variables used for the

classification of igneous rocks are particle size, which largely depends upon the cooling history,

and the mineral composition of the rock. Feldspars, quartz or feldspathoids, olivines, pyroxenes,

amphiboles, and micas are all important minerals in the formation of almost all igneous rocks,

and they are basic to the classification of these rocks. All other minerals present are regarded as

nonessential in almost all igneous rocks and are called accessory minerals. Types of igneous

rocks with other essential minerals are very rare, and these rare rocks include those with essential

carbonates.

In a simplified classification, igneous rock types are separated on the basis of the type of feldspar

present, the presence or absence of quartz, and in rocks with no feldspar or quartz, the type of

iron or magnesium minerals present. Rocks containing quartz (silica in composition) are silica-

oversaturated. Rocks with feldspathoids are silica-undersaturated, because feldspathoids cannot

coexist in a stable association with quartz.

Igneous rocks which have crystals large enough to be seen by the naked eye are called

phaneritic; those with crystals too small to be seen are called aphanitic. Generally speaking,

phaneritic implies an intrusive origin; aphanitic an extrusive one.

An igneous rock with larger, clearly discernible crystals embedded in a finer-grained matrix is

termed porphyry. Porphyritic texture develops when some of the crystals grow to considerable

size before the main mass of the magma crystallizes as finer-grained, uniform material.

Texture

Gabbro specimen showing phaneritic texture; Rock Creek Canyon, eastern Sierra Nevada,

California; scale bar is 2.0 cm.

Main article: Rock microstructure

Texture is an important criterion for the naming of volcanic rocks. The texture of volcanic rocks,

including the size, shape, orientation, and distribution of mineral grains and the intergrain

relationships, will determine whether the rock is termed a tuff, a pyroclastic lava or a simple

lava.

However, the texture is only a subordinate part of classifying volcanic rocks, as most often there

needs to be chemical information gleaned from rocks with extremely fine-grained groundmass or

from airfall tuffs, which may be formed from volcanic ash.

Textural criteria are less critical in classifying intrusive rocks where the majority of minerals will

be visible to the naked eye or at least using a hand lens, magnifying glass or microscope.

Plutonic rocks tend also to be less texturally varied and less prone to gaining structural fabrics.

Textural terms can be used to differentiate different intrusive phases of large plutons, for

instance porphyritic margins to large intrusive bodies, porphyry stocks and subvolcanic dikes

(apophyses). Mineralogical classification is used most often to classify plutonic rocks. Chemical

classifications are preferred to classify volcanic rocks, with phenocryst species used as a prefix,

e.g. "olivine-bearing picrite" or "orthoclase-phyric rhyolite".

Basic classification scheme for igneous rocks on their mineralogy. If the approximate volume

fractions of minerals in the rock are known the rock name and silica content can be read off the

diagram. This is not an exact method because the classification of igneous rocks also depends on

other components than silica, yet in most cases it is a good first guess.

Chemical classification

Igneous rocks can be classified according to chemical or mineralogical parameters:

Chemical: total alkali-silica content (TAS diagram) for volcanic rock classification used when

modal or mineralogic data is unavailable:

acid igneous rocks containing a high silica content, greater than 63% SiO2 (examples

granite and rhyolite)

intermediate igneous rocks containing between 52 - 63% SiO2 (example andesite and

dacite)

basic igneous rocks have low silica 45 - 52% and typically high iron - magnesium

content (example gabbro and basalt)

ultrabasic igneous rocks with less than 45% silica. (examples picrite and komatiite)

alkalic igneous rocks with 5 - 15% alkali (K2O + Na2O) content or with a molar ratio of

alkali to silica greater than 1:6. (examples phonolite and trachyte)

Note: the acid-basic terminology is used more broadly in older (generally British)

geological literature. In current literature felsic-mafic roughly substitutes for acid-basic.

Chemical classification also extends to differentiating rocks which are chemically similar

according to the TAS diagram, for instance;

Ultrapotassic; rocks containing molar K2O/Na2O >3

Peralkaline; rocks containing molar (K2O + Na2O)/ Al2O3 >1

Peraluminous; rocks containing molar (K2O + Na2O)/ Al2O3 <1

An idealized mineralogy (the normative mineralogy) can be calculated from the chemical

composition, and the calculation is useful for rocks too fine-grained or too altered for

identification of minerals that crystallized from the melt. For instance, normative quartz

classifies a rock as silica-oversaturated; an example is rhyolite. A normative feldspathoid

classifies a rock as silica-undersaturated; an example is nephelinite.

History of classification

In 1902 a group of American petrographers proposed that all existing classifications of igneous

rocks should be discarded and replaced by a "quantitative" classification based on chemical

analysis. They showed how vague and often unscientific was much of the existing terminology

and argued that as the chemical composition of an igneous rock was its most fundamental

characteristic it should be elevated to prime position.

Geological occurrence, structure, mineralogical constitution—the hitherto accepted criteria for

the discrimination of rock species—were relegated to the background. The completed rock

analysis is first to be interpreted in terms of the rock-forming minerals which might be expected

to be formed when the magma crystallizes, e.g., quartz feldspars, olivine, akermannite,

feldspathoids, magnetite, corundum and so on, and the rocks are divided into groups strictly

according to the relative proportion of these minerals to one another.[3][4]

Mineralogical classification

For volcanic rocks, mineralogy is important in classifying and naming lavas. The most important

criterion is the phenocryst species, followed by the groundmass mineralogy. Often, where the

groundmass is aphanitic, chemical classification must be used to properly identify a volcanic

rock.

Mineralogic contents - felsic versus mafic

felsic rock, highest content of silicon, with predominance of quartz, alkali feldspar and/or

feldspathoids: the felsic minerals; these rocks (e.g., granite, rhyolite) are usually light

coloured, and have low density.

mafic rock, lesser content of silicon relative to felsic rocks, with predominance of mafic

minerals pyroxenes, olivines and calcic plagioclase; these rocks (example, basalt, gabbro)

are usually dark coloured, and have a higher density than felsic rocks.

ultramafic rock, lowest content of silicon, with more than 90% of mafic minerals (e.g.,

dunite).

For intrusive, plutonic and usually phaneritic igneous rocks where all minerals are visible at least

via microscope, the mineralogy is used to classify the rock. This usually occurs on ternary

diagrams, where the relative proportions of three minerals are used to classify the rock.

The following table is a simple subdivision of igneous rocks according both to their composition

and mode of occurrence.

Composition

Mode of occurrence Felsic Intermediate Mafic Ultramafic

Intrusive Granite Diorite Gabbro Peridotite

Extrusive Rhyolite Andesite Basalt Komatiite

Essential rock forming silicates

Felsic Intermediate Mafic Ultramafic

Coarse Grained Granite Diorite Gabbro Peridotite

Medium Grained

Diabase

Fine Grained Rhyolite Andesite Basalt Komatiite

Example of classification

Granite is an igneous intrusive rock (crystallized at depth), with felsic composition (rich in silica

and predominately quartz plus potassium-rich feldspar plus sodium-rich plagioclase) and

phaneritic, subeuhedral texture (minerals are visible to the unaided eye and commonly some of

them retain original crystallographic shapes).

Magma origination

The Earth's crust averages about 35 kilometers thick under the continents, but averages only

some 7-10 kilometers beneath the oceans. The continental crust is composed primarily of

sedimentary rocks resting on crystalline basement formed of a great variety of metamorphic and

igneous rocks including granulite and granite. Oceanic crust is composed primarily of basalt and

gabbro. Both continental and oceanic crust rest on peridotite of the mantle.

Rocks may melt in response to a decrease in pressure, to a change in composition such as an

addition of water, to an increase in temperature, or to a combination of these processes.

Other mechanisms, such as melting from impact of a meteorite, are less important today, but

impacts during accretion of the Earth led to extensive melting, and the outer several hundred

kilometers of our early Earth probably was an ocean of magma. Impacts of large meteorites in

last few hundred million years have been proposed as one mechanism responsible for the

extensive basalt magmatism of several large igneous provinces.

Decompression

Decompression melting occurs because of a decrease in pressure.[5]

The solidus temperatures of

most rocks (the temperatures below which they are completely solid) increase with increasing

pressure in the absence of water. Peridotite at depth in the Earth's mantle may be hotter than its

solidus temperature at some shallower level. If such rock rises during the convection of solid

mantle, it will cool slightly as it expands in an adiabatic process, but the cooling is only about

0.3°C per kilometer. Experimental studies of appropriate peridotite samples document that the

solidus temperatures increase by 3°C to 4°C per kilometer. If the rock rises far enough, it will

begin to melt. Melt droplets can coalesce into larger volumes and be intruded upwards. This

process of melting from upward movement of solid mantle is critical in the evolution of Earth.

Decompression melting creates the ocean crust at mid-ocean ridges. Decompression melting

caused by the rise of mantle plumes is responsible for creating ocean islands like the Hawaiian

islands. Plume-related decompression melting also is the most common explanation for flood

basalts and oceanic plateaus (two types of large igneous provinces), although other causes such

as melting related to meteorite impact have been proposed for some of these huge volumes of

igneous rock.

Effects of water and carbon dioxide

The change of rock composition most responsible for creation of magma is the addition of water.

Water lowers the solidus temperature of rocks at a given pressure. For example, at a depth of

about 100 kilometers, peridotite begins to melt near 800°C in the presence of excess water, but

near or above about 1500°C in the absence of water.[6]

Water is driven out of the oceanic

lithosphere in subduction zones, and it causes melting in the overlying mantle. Hydrous magmas

of basalt and andesite composition are produced directly and indirectly as results of dehydration

during the subduction process. Such magmas and those derived from them build up island arcs

such as those in the Pacific ring of fire. These magmas form rocks of the calc-alkaline series, an

important part of continental crust.

The addition of carbon dioxide is relatively a much less important cause of magma formation

than addition of water, but genesis of some silica-undersaturated magmas has been attributed to

the dominance of carbon dioxide over water in their mantle source regions. In the presence of

carbon dioxide, experiments document that the peridotite solidus temperature decreases by about

200°C in a narrow pressure interval at pressures corresponding to a depth of about 70 km. At

greater depths, carbon dioxide can have more effect: at depths to about 200 km, the temperatures

of initial melting of a carbonated peridotite composition were determined to be 450°C to 600°C

lower than for the same composition with no carbon dioxide.[7]

Magmas of rock types such as

nephelinite, carbonatite, and kimberlite are among those that may be generated following an

influx of carbon dioxide into mantle at depths greater than about 70 km.

Temperature increase

Increase of temperature is the most typical mechanism for formation of magma within

continental crust. Such temperature increases can occur because of the upward intrusion of

magma from the mantle. Temperatures can also exceed the solidus of a crustal rock in

continental crust thickened by compression at a plate boundary. The plate boundary between the

Indian and Asian continental masses provides a well-studied example, as the Tibetan Plateau just

north of the boundary has crust about 80 kilometers thick, roughly twice the thickness of normal

continental crust. Studies of electrical resistivity deduced from magnetotelluric data have

detected a layer that appears to contain silicate melt and that stretches for at least 1000

kilometers within the middle crust along the southern margin of the Tibetan Plateau.[8]

Granite

and rhyolite are types of igneous rock commonly interpreted as products of melting of

continental crust because of increases of temperature. Temperature increases also may contribute

to the melting of lithosphere dragged down in a subduction zone.

Magma evolution

Schematic diagrams showing the principles behind fractional crystallisation in a magma. While

cooling, the magma evolves in composition because different minerals crystallize from the melt.

1: olivine crystallizes; 2: olivine and pyroxene crystallize; 3: pyroxene and plagioclase

crystallize; 4: plagioclase crystallizes. At the bottom of the magma reservoir, a cumulate rock

forms.

Most magmas only entirely melt for small parts of their histories. More typically, they are mixes

of melt and crystals, and sometimes also of gas bubbles. Melt, crystals, and bubbles usually have

different densities, and so they can separate as magmas evolve.

As magma cools, minerals typically crystallize from the melt at different temperatures (fractional

crystallization). As minerals crystallize, the composition of the residual melt typically changes. If

crystals separate from melt, then the residual melt will differ in composition from the parent

magma. For instance, a magma of gabbroic composition can produce a residual melt of granitic

composition if early formed crystals are separated from the magma. Gabbro may have a liquidus

temperature near 1200°C, and derivative granite-composition melt may have a liquidus

temperature as low as about 700°C. Incompatible elements are concentrated in the last residues

of magma during fractional crystallization and in the first melts produced during partial melting:

either process can form the magma that crystallizes to pegmatite, a rock type commonly enriched

in incompatible elements. Bowen's reaction series is important for understanding the idealised

sequence of fractional crystallisation of a magma.

Magma composition can be determined by processes other than partial melting and fractional

crystallization. For instance, magmas commonly interact with rocks they intrude, both by

melting those rocks and by reacting with them. Magmas of different compositions can mix with

one another. In rare cases, melts can separate into two immiscible melts of contrasting

compositions.

There are relatively few minerals that are important in the formation of common igneous rocks,

because the magma from which the minerals crystallize is rich in only certain elements: silicon,

oxygen, aluminium, sodium, potassium, calcium, iron, and magnesium. These are the elements

which combine to form the silicate minerals, which account for over ninety percent of all igneous

rocks. The chemistry of igneous rocks is expressed differently for major and minor elements and

for trace elements. Contents of major and minor elements are conventionally expressed as weight

percent oxides (e.g., 51% SiO2, and 1.50% TiO2). Abundances of trace elements are

conventionally expressed as parts per million by weight (e.g., 420 ppm Ni, and 5.1 ppm Sm).

The term "trace element" typically is used for elements present in most rocks at abundances less

than 100 ppm or so, but some trace elements may be present in some rocks at abundances

exceeding 1000 ppm. The diversity of rock compositions has been defined by a huge mass of

analytical data—over 230,000 rock analyses can be accessed on the web through a site sponsored

by the U. S. National Science Foundation (see the External Link to EarthChem).

Etymology

The word "igneous" is derived from the Latin ignis, meaning "of fire". Volcanic rocks are named

after Vulcan, the Roman name for the god of fire.

Intrusive rocks are also called plutonic rocks, named after Pluto, the Roman god of the

underworld.

Bowen's reaction series

Discontinuous

Series Continuous

Series High

Olivine

Plagioclase

(Calcium rich)

Pyroxene

Biotite

(Black Mica) Plagioclase

(Sodium rich)

Relative

Crystallization

Temperature

Orthoclase

Muscovite

(White Mica)

Quartz

Low

Within the field of geology, Bowen's reaction series is the work of the petrologist, Norman L.

Bowen who was able to explain why certain types of minerals tend to be found together while

others are almost never associated with one another. He experimented in the early 1900s with

powdered rock material that was heated until it melted and then allowed to cool to a target

temperature whereupon he observed the types of minerals that formed in the rocks produced. He

repeated this process with progressively cooler temperatures and the results he obtained led him

to formulate his reaction series which is still accepted today as the idealized progression of

minerals produced by cooling magma. Based upon Bowen's work, one can infer from the

minerals present in a rock the relative conditions under which the material had formed.

Description

Olivine weathering to iddingsite within a mantle xenolith, demonstrating the principles of the

Goldich dissolution series

The series is broken into two branches, the continuous and the discontinuous. The branch on the

right is the continuous. The minerals at the top of the illustration (given aside) are first to

crystallize and so the temperature gradient can be read to be from high to low with the high

temperature minerals being on the top and the low temperature ones on the bottom. Since the

surface of the Earth is a low temperature environment compared to the zones of rock formation,

the chart also easily shows the stability of minerals with the ones at bottom being most stable and

the ones at top being quickest to weather, known as the Goldich dissolution series. This is

because minerals are most stable in the conditions closest to those under which they had formed.

Put simply, the high temperature minerals, the first ones to crystallize in a mass of magma, are

most unstable at the Earth's surface and quickest to weather because the surface is most different

from the conditions under which they were created while the low temperature minerals are much

more stable because the conditions at the surface are much more similar to the conditions under

which they formed.

Pluton

Plutonic redirects here, for the Australian gold mine see Plutonic Gold Mine

A Jurassic pluton of pink monzonite intruded below and beneath a section of gray

sedimentary rocks and then was subsequently uplifted and exposed, near Notch Peak,

House Range, Utah.

A pluton in geology is an intrusive igneous rock (called a plutonic rock) body that crystallized

from magma slowly cooling below the surface of the Earth. Plutons include batholiths, dikes,

sills, laccoliths, lopoliths, and other igneous bodies. In practice, "pluton" usually refers to a

distinctive mass of igneous rock, typically kilometers in dimension, without a tabular shape like

those of dikes and sills. Batholiths commonly are aggregations of plutons. Examples of plutons

include Cardinal Peak and Mount Kinabalu.

The most common rock types in plutons are granite, granodiorite, tonalite, monzonite, and quartz

diorite. The term granitoid is used for a general, light colored, coarse-grained igneous rock in

which a proper, or more specific name, is not known. Use of granitoid should be restricted to the

field wherever possible.

The term originated from Pluto, the ancient Roman god of the underworld. The use of the name

and concept goes back to the beginnings of the science of geology in the late 18th century and

the then hotly debated theories of plutonism (or vulcanism), and neptunism regarding the origin

of basalt.

Batholith

Half Dome, a granite monolith in Yosemite National Park and part of the Sierra Nevada

batholith.

A batholith (from Greek bathos, depth + lithos, rock) is a large emplacement of igneous

intrusive (also called plutonic) rock that forms from cooled magma deep in the earth's crust.

Batholiths are almost always made mostly of felsic or intermediate rock-types, such as granite,

quartz monzonite, or diorite (see also granite dome).

Formation

Although they may appear uniform, batholiths are in fact structures with complex histories and

compositions. They are composed of multiple masses, or plutons, bodies of igneous rock of

irregular dimensions (typically at least several kilometers) that can be distinguished from

adjacent igneous rock by some combination of criteria including age, composition, texture, or

mappable structures. Individual plutons are crystallized from magma that traveled toward the

surface from a zone of partial melting near the base of the Earth's crust.

Traditionally, these plutons have been considered to form by ascent of relatively buoyant magma

in large masses called plutonic diapirs. Because, the diapirs are liquefied and very hot, they tend

to rise through the surrounding country rock, pushing it aside and partially melting it. Most

diapirs do not reach the surface to form volcanoes, but instead slow down, cool and usually

solidify 5 to 30 kilometers underground as plutons (hence the use of the word pluton; in

reference to the Roman god of the underworld Pluto). It has also been proposed[who?]

that plutons

commonly are formed not by diapiric ascent of large magma diapirs, but rather by aggregation of

smaller volumes of magma that ascended as dikes.[citation needed]

A batholith is formed when many plutons converge to form a huge expanse of granitic rock.

Some batholiths are mammoth, paralleling past and present subduction zones and other heat

sources for hundreds of kilometers in continental crust. One such batholith is the Sierra Nevada

Batholith, which is a continuous granitic formation that makes up much of the Sierra Nevada in

California. An even larger batholith, the Coast Plutonic Complex is found predominantly in the

Coast Mountains of western Canada, and extends for 1,800 kilometers and reaches into

southeastern Alaska.

Surface expression and erosion

A batholith is an exposed area of mostly continuous plutonic rock that covers an area larger than

100 square kilometers. Areas smaller than 100 square kilometers are called stocks. However, the

majority of batholiths visible at the surface (via outcroppings) have areas far greater than 100

square kilometers. These areas are exposed to the surface through the process of erosion

accelerated by continental uplift acting over many tens of millions to hundreds of millions of

years. This process has removed several tens of square kilometers of overlying rock in many

areas, exposing the once deeply buried batholiths.

Batholiths exposed at the surface are subjected to huge pressure differences between their former

homes deep in the earth and their new homes at or near the surface. As a result, their crystal

structure expands slightly and over time. This manifests itself by a form of mass wasting called

exfoliation. This form of erosion causes convex and relatively thin sheets of rock to slough off

the exposed surfaces of batholiths (a process accelerated by frost wedging). The result: fairly

clean and rounded rock faces. A well-known result of this process is Half Dome, located in

Yosemite Valley.

Dike (geology)

Banded gneiss with dike of granite orthogneiss.

An intrusion (Notch Peak monzonite) inter-fingers (partly as a dike) with highly-metamorphosed

host rock (Cambrian carbonate rocks). From near Notch Peak, House Range, Utah.

A dike or dyke in geology is a type of sheet intrusion referring to any geologic body that cuts

discordantly across

planar wall rock structures, such as bedding or foliation

massive rock formations, like igneous/magmatic intrusions and salt diapirs.

Dikes can therefore be either intrusive or sedimentary in origin.

Magmatic dikes

A diabase dike crosscutting horizontal limestone beds in Arizona.

A small dike on the Baranof Cross-Island Trail, Alaska.

An intrusive dike is

an igneous body with

a very high aspect

ratio, which means

that its thickness is

usually much smaller

than the other two

dimensions.

Thickness can vary

from sub-centimeter

scale to many

meters, and the

lateral dimensions can extend over many kilometers. A dike is an intrusion into an opening

cross-cutting fissure, shouldering aside other pre-existing layers or bodies of rock; this implies

that a dike is always younger than the rocks that contain it. Dikes are usually high angle to near

vertical in orientation, but subsequent tectonic deformation may rotate the sequence of strata

through which the dike propagates so that the latter becomes horizontal. Near horizontal, or

conformable intrusions, along bedding planes between strata are called intrusive sills.

Sometimes dikes appear as swarms, consisting of several to hundreds of dikes emplaced more or

less contemporaneously during a single intrusive event. The world's largest dike swarm is the

Mackenzie dike swarm in the Northwest Territories, Canada.[1]

Shiprock, New Mexico, the volcanic neck in the distance, with radiating dike on its south side.

Photo credit: USGS Digital Data Series

Dikes often form as either radial or concentric swarms around plutonic intrusives, volcanic necks

or feeder vents in volcanic cones. The latter are known as ring dikes.

Dikes can vary in texture and their composition can range from diabase or basaltic to granitic or

rhyolitic, but on a global perspective the basaltic composition prevails, manifesting ascent of vast

volumes of mantle-derived magmas through fractured lithosphere throughout Earth history.

Pegmatite dikes are extremely coarse crystalline granitic rocks often associated with late-stage

granite intrusions or metamorphic segregations. Aplite dikes are fine grained or sugary textured

intrusives of granitic composition.

Dikes in the Black Canyon of the Gunnison National Park, Colorado, USA

.

Sedimentary dikes

Clastic dike (left of notebook) in the Chinle Formation in the Island In the Sky District of

Canyonlands National Park, Utah.

Sedimentary dikes or clastic dikes are vertical bodies of sedimentary rock that cut off other rock

layers. They can form in two ways:

When a shallow unconsolidated sediment is composed of alternating coarse grained and

impermeable clay layers the fluid pressure inside the coarser layers may reach a critical

value due to lithostatic overburden. Driven by the fluid pressure the sediment breaks

through overlying layers and forms a dike.

When a soil is under permafrost conditions the pore water is totally frozen. When cracks

are formed in such rocks, they may fill up with sediments that fall in from above. The

result is a vertical body of sediment that cuts through horizontal layers: a dike.

Magmatic dikes radiating from West Spanish Peak

Sill (geology)

Illustration showing the difference between a dike and a sill.

Salisbury Crags in Edinburgh, Scotland, a sill partially exposed during the ice ages

Mid-Carboniferous dolerite sill cutting Lower Carboniferous shales and sandstones, Horton

Bluff, Minas Basin South Shore, Nova Scotia

In geology, a sill is a tabular sheet intrusion that has intruded between older layers of

sedimentary rock, beds of volcanic lava or tuff, or even along the direction of foliation in

metamorphic rock. The term sill is synonymous with concordant intrusive sheet. This means that

the sill does not cut across preexisting rocks, in contrast to dikes which do cut across older rocks.

Sills are always parallel to beds (layers) of the surrounding country rock. Usually they are in a

horizontal orientation, although tectonic processes can cause rotation of sills into near vertical

orientations. They can be confused with solidified lava flows; however, there are several

differences between them. Intruded sills will show partial melting and incorporation of the

surrounding country rock. On both the "upper" and "lower" contact surfaces of the country rock

into which the sill has intruded, evidence of heating will be observed (contact metamorphism).

Lava flows will show this evidence only on the lower side of the flow. In addition, lava flows

will typically show evidence of vesicles (bubbles) where gases escaped into the atmosphere.

Because sills generally form at depth (up to many kilometers), the pressure of overlying rock

prevents this from happening much, if at all. Lava flows will also typically show evidence of

weathering on their upper surface, whereas sills, if still covered by country rock, typically do not.

Associated ore deposits

Certain layered intrusions are a variety of sill that often contain important ore deposits.

Precambrian examples include the Bushveld, Insizwa and the Great Dyke complexes of southern

Africa, the Duluth intrusive complex of the Superior District, and the Stillwater igneous complex

of the United States. Phanerozoic examples are usually smaller and include the Rùm peridotite

complex of Scotland and the Skaergaard igneous complex of east Greenland. These intrusions

often contain concentrations of gold, platinum, chromium and other rare elements.

Transgressive sills

Despite their concordant nature, many large sills change stratigraphic level within the intruded

sequence, with each concordant part of the intrusion linked by relatively short dike-like

segments. Such sills are known as transgressive, examples include the Whin Sill and sills within

the Karoo basin.[1][2]

The geometry of large sill complexes in sedimentary basins has become

clearer with the availability of 3D seismic reflection data.[3]

Such data has shown that many sills

have an overall saucer shape and that many others are at least in part transgressive.[4]

Laccolith

A laccolith is a sheet intrusion (or concordant pluton) that has been injected between two layers

of sedimentary rock. The pressure of the magma is high enough that the overlying strata are

forced upward, giving the laccolith a dome or mushroom-like form with a generally planar base.

A laccolith intruding into and deforming strata

Laccolith exposed by erosion of overlying strata in Montana

Pink monzonite intrudes within the grey Cambrian and Ordovician strata near Notch Peak, Utah.

Laccoliths tend to form at relatively shallow depths and are typically formed by relatively

viscous magmas, such as those that crystallize to diorite, granodiorite, and granite. Cooling

underground takes place slowly, giving time for larger crystals to form in the cooling magma.

The surface rock above laccoliths often erodes away completely, leaving the core mound of

igneous rock. The term was first applied as laccolite by Grove Karl Gilbert after his study of

intrusions of diorite in the Henry Mountains of Utah in about 1875.

It is often difficult to reconstruct shapes of intrusions. For instance, Devils Tower in Wyoming

was thought to be a volcanic neck, but study has revealed it to be an eroded laccolith[1]

. The rock

would have had to cool very slowly so as to form the slender pencil-shaped columns of phonolite

porphyry seen today. However, erosion has stripped away the overlying and surrounding rock,

and so it is impossible to reconstruct the original shape of the igneous intrusion; that rock may

not be the remnant of a laccolith. At other localities, such as in the Henry Mountains and other

isolated mountain ranges of the Colorado Plateau, some intrusions demonstrably have shapes of

laccoliths. The small Barber Hill syenite-stock laccolith in Charlotte, Vermont USA, has several

volcanic trachyte dikes associated with it. Molybdenite is also visible in outcrops on this exposed

laccolith.

There are many examples of possible laccoliths on the surface of the Moon.[2]

These igneous

features may be confused with impact cratering.

Lopolith

Diagram showing the shape of a lopolith (7)

A lopolith is a large igneous intrusion which is lenticular in shape with a depressed central

region. Lopoliths are generally concordant with the intruded strata with dike or funnel-shaped

feeder bodies below the body. The term was first defined and used by Frank Fitch Grout during

the early 1900s in describing the Duluth gabbro complex in northern Minnesota and adjacent

Ontario.

Lopoliths typically consist of large layered intrusions that range in age from Archean to Eocene.

Examples include the Duluth gabbro, the Sudbury Igneous Complex of Ontario, the Bushveld

igneous complex of South Africa, the Skaergaard complex of Greenland and the Humboldt

lopolith of Nevada. The Sudbury and Bushveld occurrences have been attributed to impact

events and associated crustal melting.

Subvolcanic rock

A subvolcanic rock, also known as a hypabyssal rock, is an igneous rock that originates at

medium to shallow depths within the crust and contain intermediate grain size and often

porphyritic texture. They have textures between volcanic and plutonic rocks. Subvolcanic rocks

include diabase and porphyry.

Porphyry (geology)

.

A piece of porphyry

Rhyolite porphyry. Scale bar in lower left is 1 cm.

Porphyry is a variety of igneous rock consisting of large-grained crystals, such as feldspar or

quartz, dispersed in a fine-grained feldspathic matrix or groundmass. The larger crystals are

called phenocrysts. In its non-geologic, traditional use, the term "porphyry" refers to the purple-

red form of this stone, valued for its appearance.

The term "porphyry" is from Greek and means "purple". Purple was the color of royalty, and the

"Imperial Porphyry" was a deep purple igneous rock with large crystals of plagioclase. This rock

was prized for various monuments and building projects in Imperial Rome and later.

Subsequently the name was given to igneous rocks with large crystals. Porphyritic now refers to

a texture of igneous rocks. Its chief characteristic is a large difference between the size of the

tiny matrix crystals and other much larger phenocrysts. Porphyries may be aphanites or

phanerites, that is, the groundmass may have invisibly small crystals, like basalt, or the

individual crystals of the groundmass may be easily distinguished with the eye, as in granite.

Most types of igneous rocks may display some degree of porphyritic texture.

Granite

Granite

— Igneous Rock —

Granite containing potassium feldspar, plagioclase feldspar,

quartz, and biotite and/or amphibole

Composition

Potassium feldspar, plagioclase feldspar, and quartz;

differing amounts of muscovite, biotite, and hornblende-type

amphiboles.

Granite (pronounced /ˈɡrænɨt/) is a common and widely occurring type of intrusive, felsic,

igneous rock. Granites usually have a medium- to coarse-grained texture. Occasionally some

individual crystals (phenocrysts) are larger than the groundmass, in which case the texture is

known as porphyritic. A granitic rock with a porphyritic texture is sometimes known as a

porphyry. Granites can be pink to gray in color, depending on their chemistry and mineralogy.

By definition, granite has a color index (the percentage of the rock made up of dark minerals) of

less than 25%. Outcrops of granite tend to form tors and rounded massifs. Granites sometimes

occur in circular depressions surrounded by a range of hills, formed by the metamorphic aureole

or hornfels. Granite is usually found in the continental plates of the Earth's crust.

Granite is nearly always massive (lacking internal structures), hard and tough, and therefore it

has gained widespread use as a construction stone. The average density of granite is between

2.65[1]

and 2.75 g/cm3, its compressive strength usually lies above 200 MPa, and its viscosity at

standard temperature and pressure is 3-6 • 1019

Pa·s.[2]

The word granite comes from the Latin granum, a grain, in reference to the coarse-grained

structure of such a crystalline rock.

Granitoid is a general, descriptive field term for light-colored, coarse-grained igneous rocks.

Petrographic examination is required for identification of specific types of granitoids.[3]

Mineralogy

Orbicular granite near the town of Caldera, northern Chile

Granite is classified according to the QAPF diagram for coarse grained plutonic rocks and is

named according to the percentage of quartz, alkali feldspar (orthoclase, sanidine, or microcline)

and plagioclase feldspar on the A-Q-P half of the diagram. True granite according to modern

petrologic convention contains both plagioclase and alkali feldspars. When a granitoid is devoid

or nearly devoid of plagioclase, the rock is referred to as alkali granite. When a granitoid

contains less than 10% orthoclase, it is called tonalite; pyroxene and amphibole are common in

tonalite. A granite containing both muscovite and biotite micas is called a binary or two-mica

granite. Two-mica granites are typically high in potassium and low in plagioclase, and are

usually S-type granites or A-type granites. The volcanic equivalent of plutonic granite is rhyolite.

Granite has poor primary permeability but strong secondary permeability.

Chemical composition

A worldwide average of the chemical composition of granite, by weight percent:[4]

The Stawamus Chief is a granite monolith in British Columbia

SiO2 — 72.04%

Al2O3 — 14.42%

K2O — 4.12%

Na2O — 3.69%

CaO — 1.82%

FeO — 1.68%

Fe2O3 — 1.22%

MgO — 0.71%

TiO2 — 0.30%

P2O5 — 0.12%

MnO — 0.05%

Based on 2485 analyses

Occurrence

Granite is currently known only on Earth, where it forms a major part of continental crust.

Granite often occurs as relatively small, less than 100 km² stock masses (stocks) and in

batholiths that are often associated with orogenic mountain ranges. Small dikes of granitic

composition called aplites are often associated with the margins of granitic intrusions. In some

locations, very coarse-grained pegmatite masses occur with granite.

Granite has been intruded into the crust of the Earth during all geologic periods, although much

of it is of Precambrian age. Granitic rock is widely distributed throughout the continental crust

and is the most abundant basement rock that underlies the relatively thin sedimentary veneer of

the continents.

Origin

Close-up of granite exposed in Chennai, India.

Granite is an igneous rock and is formed from magma. Granitic magma has many potential

origins but it must intrude other rocks. Most granite intrusions are emplaced at depth within the

crust, usually greater than 1.5 kilometres and up to 50 km depth within thick continental crust.

The origin of granite is contentious and has led to varied schemes of classification. Classification

schemes are regional and include French, British, and American systems.

Geochemical origins

Various granites (cut and polished surfaces)

Granitoids are a ubiquitous component of the crust. They have crystallized from magmas that

have compositions at or near a eutectic point (or a temperature minimum on a cotectic curve).

Magmas will evolve to the eutectic because of igneous differentiation, or because they represent

low degrees of partial melting. Fractional crystallisation serves to reduce a melt in iron,

magnesium, titanium, calcium and sodium, and enrich the melt in potassium and silicon - alkali

feldspar (rich in potassium) and quartz (SiO2), are two of the defining constituents of granite.

Close-up of granite from Yosemite National Park, valley of the Merced River

This process operates regardless of the origin of the parental magma to the granite, and

regardless of its chemistry. However, the composition and origin of the magma which

differentiates into granite, leaves certain geochemical and mineral evidence as to what the

granite's parental rock was. The final mineralogy, texture and chemical composition of a granite

is often distinctive as to its origin. For instance, a granite which is formed from melted sediments

may have more alkali feldspar, whereas a granite derived from melted basalt may be richer in

plagioclase feldspar. It is on this basis that the modern "alphabet" classification schemes are

based.

Chappell & White classification system

The letter-based Chappell & White classification system was proposed initially to divide granites

into I-type granite (or igneous protolith) granite and S-type or sedimentary protolith granite.[5]

Both of these types of granite are formed by melting of high grade metamorphic rocks, either

other granite or intrusive mafic rocks, or buried sediment, respectively.

M-type or mantle derived granite was proposed later, to cover those granites which were clearly

sourced from crystallized mafic magmas, generally sourced from the mantle. These are rare,

because it is difficult to turn basalt into granite via fractional crystallisation.

A-type or anorogenic granites are formed above volcanic "hot spot" activity and have peculiar

mineralogy and geochemistry. These granites are formed by melting of the lower crust under

conditions that are usually extremely dry. The rhyolites of the Yellowstone caldera are examples

of volcanic equivalents of A-type granite.[6][7]

Granitization

An old, and largely discounted theory, granitization states that granite is formed in place by

extreme metasomatism by fluids bringing in elements e.g. potassium and removing others e.g.

calcium to transform the metamorphic rock into a granite. This was supposed to occur across a

migrating front. The production of granite by metamorphic heat is difficult, but is observed to

occur in certain amphibolite and granulite terrains. In-situ granitisation or melting by

metamorphism is difficult to recognise except where leucosome and melanosome textures are

present in gneisses. Once a metamorphic rock is melted it is no longer a metamorphic rock and is

a magma, so these rocks are seen as a transitional between the two, but are not technically

granite as they do not actually intrude into other rocks. In all cases, melting of solid rock requires

high temperature, and also water or other volatiles which act as a catalyst by lowering the solidus

temperature of the rock.

Ascent and emplacement

Roche Rock, Cornwall

The Cheesewring, a granite tor on the southern edge of Bodmin Moor, Cornwall

The ascent and emplacement of large volumes of granite within the upper continental crust is a

source of much debate amongst geologists. There is a lack of field evidence for any proposed

mechanisms, so hypotheses are predominantly based upon experimental data. There are two

major hypotheses for the ascent of magma through the crust:

Stokes Diapir

Fracture Propagation

Of these two mechanisms, Stokes diapir was favoured for many years in the absence of a

reasonable alternative. The basic idea is that magma will rise through the crust as a single mass

through buoyancy. As it rises it heats the wall rocks, causing them to behave as a power-law

fluid and thus flow around the pluton allowing it to pass rapidly and without major heat loss.[8]

This is entirely feasible in the warm, ductile lower crust where rocks are easily deformed, but

runs into problems in the upper crust which is far colder and more brittle. Rocks there do not

deform so easily: for magma to rise as a pluton it would expend far too much energy in heating

wall rocks, thus cooling and solidifying before reaching higher levels within the crust.

Nowadays fracture propagation is the mechanism preferred by many geologists as it largely

eliminates the major problems of moving a huge mass of magma through cold brittle crust.

Magma rises instead in small channels along self-propagating dykes which form along new or

pre-existing fault systems and networks of active shear zones (Clemens, 1998).[9]

As these

narrow conduits open, the first magma to enter solidifies and provides a form of insulation for

later magma.

Granitic magma must make room for itself or be intruded into other rocks in order to form an

intrusion, and several mechanisms have been proposed to explain how large batholiths have been

emplaced:

Stoping, where the granite cracks the wall rocks and pushes upwards as it removes blocks

of the overlying crust

Assimilation, where the granite melts its way up into the crust and removes overlying

material in this way

Inflation, where the granite body inflates under pressure and is injected into position

Most geologists today accept that a combination of these phenomena can be used to explain

granite intrusions, and that not all granites can be explained entirely by one or another

mechanism.

Granodiorite

A sample of granodiorite from Massif Central, France

Photomicrograph of thin section of granodiorite from Slovakia (in crossed polarised light)

Granodiorite (pronounced /ˌɡrænɵˈdaɪ.ɵraɪt/ or /ˌɡreɪnɵˈdaɪ.ɵraɪt/) is an intrusive igneous rock

similar to granite, but containing more plagioclase than potassium feldspar. Officially, it is

defined as a phaneritic igneous rock with greater than 20% quartz by volume where at least 65%

of the feldspar is plagioclase. It usually contains abundant biotite mica and hornblende, giving it

a darker appearance than true granite. Mica may be present in well-formed hexagonal crystals,

and hornblende may appear as needle-like crystals.

Geology

On average the upper continental crust has the same composition as granodiorite.

Granodiorite is a plutonic igneous rock, formed by an intrusion of silica-rich magma, which

cools in batholiths or stocks below the Earth's surface. It is usually only exposed at the surface

after uplift and erosion have occurred. The volcanic equivalent of granodiorite is dacite.

Syenite

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Syenite

leucocratic variety of nepheline syenite from Sweden (särnaite).

Syenite is a coarse-grained intrusive igneous rock of the same general composition as granite but

with the quartz either absent or present in relatively small amounts (<5%).

The feldspar component of syenite is predominantly alkaline in character (usually orthoclase) .

Plagioclase feldspars may be present in small quantities, less than 10%.

When present, ferromagnesian minerals are usually hornblende amphibole, rarely pyroxene or

biotite. Biotite is rare, because in a syenite magma most aluminium is used in producing feldspar.

Syenites are usually peralkaline and peraluminous, with high proportions of alkali elements and

aluminium.

Syenites are formed from alkaline igneous activity, generally formed in thick continental crustal

areas, or in Cordilleran subduction zones. To produce a syenite, it is necessary to melt a granitic

or igneous protolith to a fairly low degree of partial melting. This is required because potassium

is an incompatible element and tends to enter a melt first, whereas higher degrees of partial

melting will liberate more calcium and sodium, which produce plagioclase, and hence a granite,

adamellite or tonalite.

At very low degrees of partial melting a silica undersaturated melt is produced, forming a

nepheline syenite, where orthoclase is replaced by a feldspathoid such as leucite, nepheline or

analcime.

Syenite is not a common rock, some of the more important occurrences being in New England,

Arkansas, Montana, New York (syenite gneisses), Switzerland, Germany, and Norway.

Etymology

The term syenite was originally applied to hornblende granite like that of Syene in Egypt, from

which the name is derived.

Episyenite

Episyenite (or epi-syenite) is a term used in petrology to describe to the result of alteration of a

SiO2 rich rock to a more SiO2 depleted rock.

The process which results in SiO2 depletion can be termed episyenitization. This process is only

referring to the macroscopic result of relative SiO2 depletion in a rock. The actual physical

process leading to this SiO2 depletion may vary in a given metamorphic environment. Diffusion

of chemical components in a stagnant fluid, related to differences in chemical potential or

pressure as well as advection of a SiO2- undersaturated fluid may lead to the dissolution of quartz

from the un-altered rock, thus depleting it of this component.

Nepheline syenite

Nepheline syenite from Sweden

Nephelene syenite is a holocrystalline plutonic rock that consists largely of nepheline and alkali

feldspar. The rocks are mostly pale colored, grey or pink, and in general appearance they are not

unlike granites, but dark green varieties are also known. Phonolite is the fine-grained extrusive

equivalent.

Petrology

Nepheline syenites are silica-undersaturated and some are peralkaline (terms discussed in

igneous rock). Nepheline is a feldspathoid, a solid-solution mineral, that does not coexist with

quartz; rather, nepheline would react with quartz to produce alkali feldspar.

They are distinguished from ordinary syenites not only by the presence of nepheline but also by

the occurrence of many other minerals rich in alkalis and in rare earths and other incompatible

elements. Alkali feldspar dominates, commonly represented by orthoclase and the exsolved

lamellar albite, form perthite. In some rocks the potash feldspar, in others the soda feldspar

predominates. Fresh clear microcline is very characteristic of some types of nepheline syenite.

Sodalite, colorless and transparent in thin section, but frequently pale blue in the hand

specimens, is the principal feldspathoid mineral in addition to nepheline. Reddish-brown to black

triclinic aenigmatite occurs also in these rocks. Extremely iron-rich olivine is rare, but is present

in some nepheline syenite. Other minerals common in minor amounts include sodium-rich

pyroxene, biotite, titanite, zircon, iron oxides, apatite, fluorite, melanite garnet, and zircon.

Cancrinite occurs in several nepheline-syenites. A great number of interesting and rare minerals

have been recorded from nepheline syenites and the pegmatite veins which intersect them.

Genesis

Silica-undersaturated igneous rocks typically are formed by low degrees of partial melting in the

Earth's mantle. Carbon dioxide may dominate over water in source regions. Magmas of such

rocks are formed in a variety of environments, including continental rifts, ocean islands, and

supra-subduction positions in subduction zones. Nepheline syenite and phonolite may be derived

by crystal fractionation from more mafic silica-undersaturated mantle-derived melts, or as partial

melts of such rocks. Igneous rocks with nepheline in their normative mineralogy commonly are

associated with other unusual igneous rocks such as carbonatite.

Distribution

Nepheline syenites and phonolites occur in Canada, Norway, Greenland, Sweden, the Ural

Mountains, the Pyrenees, Italy, Brazil, China, the Transvaal region, and Magnet Cove igneous

complex of Arkansas, as well as on oceanic islands.

Phonolite lavas formed in the East African rift in particularly large quantity, and the volume

there may exceed the volume of all other phonolite occurrences combined, as discussed by

Barker (1983).

Nepheline-normative rocks occur in close association with the Bushveld Igneous Complex,

possibly formed from partial melting of the wall rocks to that large ultramafic layered intrusion.

Nepheline syenites are rare; there is only one occurrence in Great Britain and one in France and

Portugal. They are known also in Bohemia and in several places in Norway, Sweden and

Finland. In the Americas these rocks have been found in Texas, Arkansas and Massachusetts,

also in Ontario, British Columbia and Brazil. South Africa, Madagascar, India, Tasmania, Timor

and Turkestan are other localities for the rocks of this series.

Rocks of this class also occur in Brazil (Serra de Tingua) containing sodalite and often much

augite, in the western Sahara and Cape Verde Islands; also at Zwarte Koppies in the Transvaal,

Madagascar, São Paulo in Brazil, Paisano Pass in West Texas and Montreal, Canada. The rock of

Salem, Massachusetts, United States, is a mica-foyaite rich in albite and aegirine: it accompanies

granite and essexite. Litchfieldite is another well-marked type of nepheline-syenite, in which

albite is the dominant feldspar. It is named after Litchfield, Maine, United States, where it occurs

in scattered blocks. Biotite, cancrinite and sodalite are characteristic of this rock. A similar

nepheline-syenite is known from Hastings County, Ontario, and contains hardly any orthoclase,

but only albite feldspar. Nepheline is very abundant and there is also cancrinite, sodalite,

scapolite, calcite, biotite and hornblende. The lujaurites are distinguished from the rocks above

described by their dark color, which is due to the abundance of minerals such as augite, aegirine,

arfvedsonite and other kinds of amphibole. Typical examples are known near Lujaur on the

White Sea, where they occur with umptekites and other very peculiar rocks. Other localities for

this group are at Julianehaab in Greenland with sodalite-syenite; at their margins they contain

pseudomorphs after leucite. The lujaurites frequently have a parallel-banding or gneissose

structure. Sodalite-syenites in which sodalite very largely or completely takes the place of

nepheline occur in Greenland, where they contain also microcline-perthite, aegirine, arfvedsonite

and eudialyte.

Cancrinite syenite, with a large percentage of cancrinite, has been described from Dalekarlia,

Sweden and from Finland. We may also mention urtite from Lujaur Urt on the White Sea, which

consists very largely of nepheline, with aegirine and apatite, but no feldspar. Jacupirangite (from

Jacupiranga in Brazil) is a blackish rock composed of titaniferous augite, magnetite, ilmenite,

perofskite and nepheline, with secondary biotite.

Nomenclature

There is a wide variety of silica-undersaturated and peralkaline igneous rocks, including many

informal place-name varieties named after the locations in which they were first discovered. In

many cases these are plain nepheline syenites containing one or more rare minerals or

mineraloids, which do not warrant a new formal classification. These include;

Foyaite: foyaites are named after Foya in the Serra de Monchique, in southern Portugal. These

are K-feldspar-nepheline syenites containing <10% ferromagnesian minerals, usually pyroxene-,

hornblende- and biotite.

Laurdalite: The laurdalites, from Laurdal in Norway, are grey or pinkish, and in many ways

closely resemble the laurvikites of southern Norway, with which they occur. They contain

anorthoclase feldspars, biotite or greenish augite, much apatite and in some cases, olivine.

Ditroite: Ditroite derives is name from Ditrau, Transylvania, Romania. It is essentially a

microcline, sodalite and cancrinite variety of nepheline syenite. It contains also orthoclase,

nepheline, biotite, aegirine, acmite.

Chemical composition

The chemical peculiarities of the nepheline-syenites are well marked. They are exceedingly rich

in alkalis and in alumina (hence the abundance of felspathoids and alkali feldspars) with silica

varying from 50 to 56%, while lime, magnesia[disambiguation needed]

and iron are never present in

great quantity, though somewhat more variable than the other components. A worldwide average

of the major elements in nepheline syenite tabulated by Barker (1983) is listed below, expressed

as weight percent oxides.

SiO2 — 54.99%

TiO2 — 0.60%

Al2O3 — 20.96%

Fe2O3 — 2.25%

FeO — 2.05%

MnO — 0.15%

MgO — 0.77%

CaO — 2.31%

Na2O — 8.23%

K2O — 5.58%

H2O — 1.47%

P2O5 — 0.13%

The normative mineralogy of this average composition contains about 22 percent nepheline and

66 percent feldspar.

Monzonite

Photomicrograph of thin section of monzonite (in cross polarised light)

The QAPF diagram, by which a monzonite is defined

Photomicrograph of thin section of monzonite (in plane polarised light)

An intrusion (Notch Peak monzonite) inter-fingers (partly as a dike) with highly-metamorphosed

host rock (Cambrian carbonate rocks). From near Notch Peak, House Range, Utah.

Monzonite is an intermediate igneous intrusive rock composed of approximately equal amounts

of sodic to intermediate plagioclase and orthoclase feldspars with minor amounts of hornblende,

biotite and other minerals. Quartz a minor constituent or is absent; with greater than 10% quartz

the rock is termed a quartz monzonite.

If the rock has more orthoclase or potassium feldspar it grades into a syenite. With an increase of

calcic plagioclase and mafic minerals the rock type becomes a diorite. The volcanic equivalent is

the latite.

Tonalite

A piece of tonalite on red granite gneiss from Tjörn, Sweden

Tonalite is an igneous, plutonic (intrusive) rock, of felsic composition, with phaneritic texture.

Feldspar is present as plagioclase (typically oligoclase or andesine) with 10% or less alkali

feldspar. Quartz is present as more than 20% of the rock. Amphiboles and pyroxenes are

common accessory minerals.

In older references tonalite is sometimes used as a synonym for quartz diorite. However the

current IUGS classification defines tonalite as having greater than 20% quartz and quartz diorite

with from 5 to 20% quartz.

The name is derived from the type locality of tonalites, adjacent to the Tonale Line, a major

structural lineament and mountain pass, Tonale Pass, in the Italian and Austrian Alps.

Trondhjemite is an orthoclase-deficient variety of tonalite with minor biotite as the only mafic

mineral, named after Norway's third largest city, Trondheim.

Igneous rocks by composition

Type Ultramafic

< 45% SiO2

Mafic

< 52% SiO2

Intermediate

52–63% SiO2

Intermediate-

Felsic

63–69% SiO2

Felsic

>69 % SiO2

Volcanic

rocks:

Subvolcanic

rocks:

Plutonic rocks:

Komatiite

Kimberlite,

Lamproite

Peridotite

Basalt

Diabase

(Dolerite)

Gabbro

Andesite

Diorite

Dacite

Granodiorite

Rhyolite

Aplite—

Pegmatite

Granite

Diorite

Diorite

Diorite classification on QAPF diagram

Diorite (pronounced /ˈdaɪəraɪt/) is a grey to dark grey intermediate intrusive igneous rock

composed principally of plagioclase feldspar (typically andesine), biotite, hornblende, and/or

pyroxene. It may contain small amounts of quartz, microcline and olivine. Zircon, apatite,

sphene, magnetite, ilmenite and sulfides occur as accessory minerals.[1]

It can also be black or

bluish-grey, and frequently has a greenish cast. Varieties deficient in hornblende and other dark

minerals are called leucodiorite. When olivine and more iron-rich augite are present, the rock

grades into ferrodiorite, which is transitional to gabbro. The presence of significant quartz makes

the rock type quartz-diorite (>5% quartz) or tonalite (>20% quartz), and if orthoclase (potassium

feldspar) is present at greater than ten percent the rock type grades into monzodiorite or

granodiorite. Diorite has a medium grain size texture, occasionally with porphyry.

Diorites may be associated with either granite or gabbro intrusions, into which they may subtly

merge. Diorite results from partial melting of a mafic rock above a subduction zone. It is

commonly produced in volcanic arcs, and in cordilleran mountain building such as in the Andes

Mountains as large batholiths. The extrusive volcanic equivalent rock type is andesite.

Occurrence

Diorite

Diorite is a relatively rare rock; source localities include Leicestershire; UK [2]

(one name for

microdiorite - Markfieldite - exists due to the rock being found in the village of Markfield),

Sondrio, Italy; Thuringia and Saxony in Germany; Finland; Romania; Northeastern Turkey;

central Sweden; Scotland; the Darrans range of New Zealand; the Andes Mountains; the Isle of

Guernsey; Basin and Range province and Minnesota in the USA; Idahet in Egypt

An orbicular variety found in Corsica is called corsite.

Gabbro

Gabbro specimen; Rock Creek Canyon, eastern Sierra Nevada, California.

Close-up of gabbro specimen; Rock Creek Canyon, eastern Sierra Nevada, California.

Photomicrograph of a thin section of gabbro.

Gabbro (pronounced /ˈɡæbroʊ/) refers to a large group of dark, coarse-grained, intrusive mafic

igneous rocks chemically equivalent to basalt. The rocks are plutonic, formed when molten

magma is trapped beneath the Earth's surface and cools into a crystalline mass.

The vast majority of the Earth's surface is underlain by gabbro within the oceanic crust, produced

by basalt magmatism at mid-ocean ridges.

Petrology

A gabbro landscape on the main ridge of the Cuillin, Isle of Skye, Scotland.

Gabbro as a xenolith in a granite, eastern Sierra Nevada, Rock Creek Canyon, California.

Gabbro is dense, greenish or dark-colored and contains pyroxene, plagioclase, amphibole, and

olivine (olivine gabbro when olivine is present in a large amount).

The pyroxene is mostly clinopyroxene; small amounts of orthopyroxene may be present. If the

amount of orthopyroxene is substantially greater than the amount of clinopyroxene, the rock is

then a norite. Quartz gabbros are also known to occur and are probably derived from magma that

was over-saturated with silica. Essexites represent gabbros whose parent magma was under-

saturated with silica, resulting in the formation of the feldspathoid mineral nepheline. (Silica

saturation of a rock can be evaluated by normative mineralogy). Gabbros contain minor amounts,

typically a few percent, of iron-titanium oxides such as magnetite, ilmenite, and ulvospinel.

Gabbro is generally coarse grained, with crystals in the size range of 1 mm or greater. Finer

grained equivalents of gabbro are called diabase, although the vernacular term microgabbro is

often used when extra descriptiveness is desired. Gabbro may be extremely coarse grained to

pegmatitic, and some pyroxene-plagioclase cumulates are essentially coarse grained gabbro,

although these may exhibit acicular crystal habits.

Gabbro is usually equigranular in texture, although it may be porphyritic at times, especially

when plagioclase oikocrysts have grown earlier than the groundmass minerals.

Distribution

Gabbro can be formed as a massive, uniform intrusion via in-situ crystallisation of pyroxene and

plagioclase, or as part of a layered intrusion as a cumulate formed by settling of pyroxene and

plagioclase. Cumulate gabbros are more properly termed pyroxene-plagioclase orthocumulate.

Gabbro is an essential part of the oceanic crust, and can be found in many ophiolite complexes as

parts of zones III and IV (sheeted dyke zone to massive gabbro zone). Long belts of gabbroic

intrusions are typically formed at proto-rift zones and around ancient rift zone margins, intruding

into the rift flanks. Mantle plume hypotheses may rely on identifying mafic and ultramafic

intrusions and coeval basalt volcanism.

Norite

Norite is a mafic intrusive igneous rock composed largely of the calcium-rich plagioclase

labradorite and hypersthene with olivine. Norite is essentially indistinguishable from gabbro

without thin section study under the petrographic microscope. It occurs with gabbro and other

mafic to ultramafic rocks in layered intrusions which are often associated with platinum

orebodies such as in the Bushveld Igneous Complex in South Africa, the Skaergaard igneous

complex of Greenland, and the Stillwater igneous complex in Montana, USA. Norite is also the

basal igneous rock of the Sudbury Basin complex in Ontario which is the site of a meteorite

impact and the world's second largest nickel mining region. Norite is a common rock type of the

Apollo samples. On a smaller scale, norite can be found in small localized intrusions such as the

Gombak Norite in Bukit Gombak, Singapore.

The name Norite is derived from the Norwegian name for Norway: Norge.

Anorthosite

Anorthosite from Poland

Lunar anorthosite from Apollo 15 landing site

Anorthosite (pronounced /ænˈɔrθəsaɪt/) is a phaneritic, intrusive igneous rock characterized by a

predominance of plagioclase feldspar (90–100%), and a minimal mafic component (0–10%).

Pyroxene, ilmenite, magnetite, and olivine are the mafic minerals most commonly present.

Anorthosite on Earth can be divided into two types: Proterozoic anorthosite (also known as

massif or massif-type anorthosite) and Archean anorthosite. These two types of anorthosite have

different modes of occurrence, appear to be restricted to different periods in Earth's history, and

are thought to have had different origins.

Lunar anorthosites constitute the light-coloured areas of the Moon's surface and have been the

subject of much research.[1]

Proterozoic anorthosite

Age

Although a few Proterozoic anorthosite bodies were emplaced either late in the Archean Eon, or

early in the Phanerozoic Eon, the vast majority of Proterozoic anorthosites were emplaced, as

their name suggests, during the Proterozoic Eon (ca. 2,500-542 Ma).

Mode of occurrence

Anorthosite from southern Finland

Anorthosite plutons occur in a wide range of sizes. Some smaller plutons, exemplified by many

anorthosite bodies in the U.S. and Harris in Scotland, cover only a few dozen square kilometres.

Larger plutons, like the Mt. Lister Anorthosite, in northern Labrador, Canada, cover several

thousands of square kilometres.

Many Proterozoic anorthosites occur in spatial association with other highly distinctive,

contemporaneous rock types (the so-called 'anorthosite suite' or 'anorthosite-mangerite-

charnockite complex'). These rock types include iron-rich diorite, gabbro, and norite; leucocratic

mafic rocks such as leucotroctolite and leuconorite; and iron-rich felsic rocks, including

monzonite and rapakivi granite. Importantly, large volumes of ultramafic rocks are not found in

association with Proterozoic anorthosites.

Occurrences of Proterozoic anorthosites are commonly referred to as 'massifs'. However, there is

some question as to what name would best describe any occurrence of anorthosite together with

the rock types mentioned above. Early works used the term 'complex' The term 'plutonic suite'

has been applied to some large occurrences in northern Labrador, Canada; however, it has been

suggested (in 2004-2005) that 'batholith' would be a better term. 'Batholith' is used to describe

such occurrences for the remainder of this article.

The areal extent of anorthosite batholiths ranges from relatively small (dozens or hundreds of

square kilometres) to nearly 20,000 km2 (7,700 sq mi), in the instance of the Nain Plutonic Suite

in northern Labrador, Canada.

Major occurrences of Proterozoic anorthosite are found in the southwest U.S., the Appalachian

Mountains, eastern Canada, across southern Scandinavia and eastern Europe. Mapped onto the

Pangaean continental configuration of that eon, these occurrences are all contained in a single

straight belt, and must all have been emplaced intracratonally. The conditions and constraints of

this pattern of origin and distribution are not clear. However, see the Origins section below.

Anorthosites are also common in layered intrusions. Anorthosite in these layered intrusions can

form as cumulate layers in the upper parts of the intrusive complex[2]

or as later-stage intrusions

into the layered intrusion complex.[3]

Physical characteristics

Since they are primarily composed of plagioclase feldspar, most of Proterozoic anorthosites

appear, in outcrop, to be grey or bluish. Individual plagioclase crystals may be black, white, blue,

or grey, and may exhibit an iridescence known as labradorescence on fresh surfaces. The

feldspar variety labradorite is commonly present in anorthosites. Mineralogically, labradorite is a

compositional term for any calcium-rich plagioclase feldspar containing between 50–70

molecular percent anorthite (An 50–70), regardless of whether it shows labradorescence. The

mafic mineral in Proterozoic anorthosite may be clinopyroxene, orthopyroxene, olivine, or, more

rarely, amphibole. Oxides, such as magnetite or ilmenite, are also common.

Most anorthosite plutons are very coarse grained; that is, the individual plagioclase crystals and

the accompanying mafic mineral are more than a few centimetres long. Less commonly,

plagioclase crystals are megacrystic, or larger than one metre long. However, most Proterozoic

anorthosites are deformed, and such large plagioclase crystals have recrystallized to form smaller

crystals, leaving only the outline of the larger crystals behind.

While many Proterozoic anorthosite plutons appear to have no large-scale relict igneous

structures (having instead post-emplacement deformational structures), some do have igneous

layering, which may be defined by crystal size, mafic content, or chemical characteristics. Such

layering clearly has origins with a rheologically liquid-state magma.

Chemical and isotopic characteristics

The composition of plagioclase feldspar in Proterozoic anorthosites is most commonly between

An40 and An60 (40-60% anorthite). This compositional range is intermediate, and is one of the

characteristics which distinguish Proterozoic anorthosites from Archean anorthosites. Mafic

minerals in Proterozoic anorthosites have a wide range of composition, but are not generally

highly magnesian.

The trace-element chemistry of Proterozoic anorthosites, and the associated rock types, has been

examined in some detail by researchers with the aim of arriving at a plausible genetic theory.

However, there is still little agreement on just what the results mean for anorthosite genesis; see

the 'Origins' section below. A very short list of results, including results for rocks thought to be

related to Proterozoic anorthosites.[4]

Some research has focused on neodymium (Nd) and strontium (Sr) isotopic determinations for

anorthosites, particularly for anorthosites of the Nain Plutonic Suite (NPS). Such isotopic

determinations are of use in gauging the viability of prospective sources for magmas that gave

rise to anorthosites. Some results are detailed below in the 'Origins' section.

Origins of Proterozoic anorthosites

The origins of Proterozoic anorthosites have been a subject of theoretical debate for many

decades. A brief synopsis of this problem is as follows. The problem begins with the generation

of magma, the necessary precursor of any igneous rock.

Magma generated by small amounts of partial melting of the mantle is generally of basaltic

composition. Under normal conditions, the composition of basaltic magma requires it to

crystallize between 50 and 70% plagioclase, with the bulk of the remainder of the magma

crystallizing as mafic minerals. However, anorthosites are defined by a high plagioclase content

(90–100% plagioclase), and are not found in association with contemporaneous ultramafic rocks.

This is now known as 'the anorthosite problem'. Proposed solutions to the anorthosite problem

have been diverse, with many of the proposals drawing on different geological subdisciplines.

It was suggested early in the history of anorthosite debate that a special type of magma,

anorthositic magma, had been generated at depth, and emplaced into the crust. However, the

solidus of an anorthositic magma is too high for it to exist as a liquid for very long at normal

ambient crustal temperatures, so this appears to be unlikely. The presence of water vapour has

been shown to lower the solidus temperature of anorthositic magma to more reasonable values,

but most anorthosites are relatively dry. It may be postulated, then, that water vapour be driven

off by subsequent metamorphism of the anorthosite, but some anorthosites are undeformed,

thereby invalidating the suggestion.

The discovery, in the late 1970s, of anorthositic dykes in the Nain Plutonic Suite, suggested that

the possibility of anorthositic magmas existing at crustal temperatures needed to be reexamined.

However, the dykes were later shown to be more complex than was originally thought. In

summary, though liquid-state processes clearly operate in some anorthosite plutons, the plutons

are probably not derived from anorthositic magmas.

Many researchers have argued that anorthosites are the products of basaltic magma, and that

mechanical removal of mafic minerals has occurred. Since the mafic minerals are not found with

the anorthosites, these minerals must have been left at either a deeper level or the base of the

crust. A typical theory is as follows: partial melting of the mantle generates a basaltic magma,

which does not immediately ascend into the crust. Instead, the basaltic magma forms a large

magma chamber at the base of the crust and fractionates large amounts of mafic minerals, which

sink to the bottom of the chamber. The cocrystallizing plagioclase crystals float, and eventually

are emplaced into the crust as anorthosite plutons. Most of the sinking mafic minerals form

ultramafic cumulates which stay at the base of the crust.

This theory has many appealing features, of which one is the capacity to explain the chemical

composition of high-alimuna orthopyroxene megacrysts (HAOM). This is detailed below in the

section devoted to the HAOM. However, on its own, this hypothesis cannot coherently explain

the origins of anorthosites, because it does not fit with, among other things, some important

isotopic measurements made on anorthositic rocks in the Nain Plutonic Suite. The Nd and Sr

isotopic data shows the magma which produced the anorthosites cannot have been derived only

from the mantle. Instead, the magma that gave rise to the Nain Plutonic Suite anorthosites must

have had a significant crustal component. This discovery led to a slightly more complicated

version of the previous hypothesis: Large amounts of basaltic magma form a magma chamber at

the base of the crust, and, while crystallizing, assimilating large amounts of crust.[5]

This small addendum explains both the isotopic characteristics and certain other chemical

niceties of Proterozoic anorthosite. However, at least one researcher has cogently argued, on the

basis of geochemical data, that the mantle's role in production of anorthosites must actually be

very limited: the mantle provides only the impetus (heat) for crustal melting, and a small amount

of partial melt in the form of basaltic magma. Thus anorthosites are, in this view, derived almost

entirely from lower crustal melts.[6]

High-alumina orthopyroxene megacrysts

The high-alumina orthopyroxene megacrysts (HAOM) have, like Proterozoic anorthosites, been

the subject of great debate, although a tentative consensus about their origin appears to have

emerged. The peculiar characteristic worthy of such debate is reflected in their name. Normal

orthopyroxene has chemical composition (Fe,Mg)2Si2O6, whereas the HAOM have anomalously

large amounts of aluminium (up to about 9%) in their atomic structure.

Because the solubility of aluminium in orthopyroxene increases with increasing pressure, many

researchers,[7]

have suggested that the HAOM crystallized at depth, near the base of the Earth's

crust. The maximum amounts of aluminium correspond to a 30–35 km (19–22 mi) depth.

Other researchers consider the chemical compositions of the HAOM to be the product of rapid

crystallization at moderate or low pressures.[8]

Archaean anorthosite

Smaller amounts of anorthosite were emplaced during the Archaean eon (ca 3,800-2,400 Ma),

although most have been dated between 3,200 and 2,800 Ma. They are distinct texturally and

mineralogically from Proterozoic anorthosite bodies. Their most characteristic feature is the

presence of equant megacrysts of plagioclase surrounded by a fine-grained mafic groundmass.

Diabase

Diabase

Diabase (pronounced /ˈdaɪ.əbeɪs/) or Dolerite is a mafic, holocrystalline, subvolcanic rock

equivalent to volcanic basalt or plutonic gabbro. In North American usage, the term diabase

refers to the fresh rock, whilst elsewhere the term dolerite is used for the fresh rock and diabase

refers to altered material.[1][2]

Diabase dikes and sills are typically shallow intrusive bodies and

often exhibit fine grained to aphanitic chilled margins which may contain tachylite (dark mafic

glass).

Petrology

Diabase normally has a fine, but visible texture of euhedral lath-shaped plagioclase crystals

(62%) set in a finer matrix of clinopyroxene, typically augite (20–29%), with minor olivine (3%

up to 12% in olivine diabase), magnetite (2%), and ilmenite (2%).[3]

Accessory and alteration

minerals include hornblende, biotite, apatite, pyrrhotite, chalcopyrite, serpentine, chlorite, and

calcite. The texture is termed diabasic and is typical of diabases. This diabasic texture is also

termed interstitial[4]

. The feldspar is high in anorthite (as opposed to albite), the calcium

endmember of the plagioclase anorthite-albite solid solution series, most commonly labradorite.

Diabase/dolerite

The Candlestick, Tasman Peninsula, Tasmania, is composed of Jurassic Dolerite. Tasmania has

the world's largest areas of dolerite.

In non-North American usage dolerite is preferred due to the various conflicting uses of diabase.

Dolerite (Greek: doleros, meaning "deceptive") was the name given by Haüy in his 1822 Traité

de minéralogie. In continental Europe diabase was reserved by Brongniart for pre-Tertiary (pre-

Cenozoic) material[5]

, with dolerite used for more recent rock. The use of diabase in this sense

was abandoned in Britain in favor of dolerite for rocks of all ages by Allport (1874)[6]

, though

some British geologists continued to use diabase to describe slightly altered dolerite, in which

pyroxene has been altered to amphibole.[7]

Locations

A diabase dike crosscutting horizontal limestone beds in Arizona

Diabase is usually found in smaller relatively shallow intrusive bodies such as dikes and sills.

Diabase dikes occur in regions of crustal extension and often occur in dike swarms of hundreds

of individual dikes or sills radiating from a single volcanic center.

The Palisades Sill which makes up the New Jersey Palisades on the Hudson River, near New

York City, is an example of a diabase sill. The dike complexes of the British Tertiary Volcanic

Province which includes Skye, Rum, Mull, and Arran of western Scotland, the Slieve Gullion

region of Ireland, and extends across northern England contains many examples of diabase dike

swarms. Parts of the Deccan Traps of India, formed at the end of the Cretaceous also includes

dolerite[8]

. It is also abundant in large parts of Curaçao, an island off the coast of Venezuela.

In Western Australia a 200 km long dolerite dike, the Norseman–Wiluna Belt[9]

is associated

with the non-alluvial gold mining area between Norseman and Kalgoolie, which includes the

largest gold mine in Australia[10]

, the Super Pit gold mine. West of the Norseman–Wiluna Belt is

the Yalgoo–Singleton Belt, where complex dolerite dike swarms obscure the volcaniclastic

sediments.[11]

The vast areas of mafic volcanism/plutonism associated with the Jurassic breakup of

Gondwanaland in the Southern Hemisphere include many large diabase/dolerite sills and dike

swarms. These include the Karoo dolerites of South Africa, the Ferrar Dolerites of Antarctica,

and the largest of these, indeed the most extensive of all dolerite formations worldwide, are

found in Tasmania. Here, the volume of magma which intruded into a thin veneer of Permian

and Triassic rocks from multiple feeder sites, over a period of perhaps a million years, may have

exceeded 40,000 cubic kilometres.[12]

In Tasmania alone dolerite dominates the landscape.

Ring dikes are large, near vertical dikes showing above ground as circular outcrops up to 30 km

in diameter, with a depth from hundreds of metres to several kilometres. Thicker dikes are made

up of plutonic rocks, rather than hypabyssal and are centred around deep intrusions. The central

part may be a block sunken into underlying magma, the ring dikes forming in the fracture zone

around the sunken block.

Peridotite

Peridotite

— Igneous Rock —

Peridotite xenolith from San Carlos, southwestern United

States. The rock is typical olivine-rich peridotite, cut by a

centimeter-thick layer of greenish-black pyroxenite.

Composition

olivine, pyroxene

A peridotite is a dense, coarse-grained igneous rock, consisting mostly of the minerals olivine

and pyroxene. Peridotite is ultramafic, as the rock contains less than 45% silica. It is high in

magnesium, reflecting the high proportions of magnesium-rich olivine, with appreciable iron.

Peridotite is derived from the Earth's mantle, either as solid blocks and fragments, or as crystals

accumulated from magmas that formed in the mantle. The compositions of peridotites from these

layered igneous complexes vary widely, reflecting the relative proportions of pyroxenes,

chromite, plagioclase, and amphibole.

Peridotite is the dominant rock of the upper part of the Earth's mantle. The compositions of

peridotite nodules found in certain basalts and diamond pipes (kimberlites) are of special interest,

because they provide samples of the Earth's Mantle roots of continents brought up from depths

from about 30 km or so to depths at least as great as about 200 km. Some of the nodules preserve

isotope ratios of osmium and other elements that record processes over three billion years ago,

and so they are of special interest to paleogeologists because they provide clues to the

composition of the Earth's early mantle and the complexities of the processes that were involved.

The word peridotite comes from the gemstone peridot, which consists of pale green olivine.[1]

Types of peridotite

Dunite: more than 90% olivine, typically with Mg/Fe ratio of about 9:1.

Wehrlite: mostly composed of olivine plus clinopyroxene.

Harzburgite: mostly composed of olivine plus orthopyroxene, and relatively low

proportions of basaltic ingredients (because garnet and clinopyroxene are minor).

Lherzolite: mostly composed of olivine, orthopyroxene (commonly enstatite), and

clinopyroxene (diopside), and have relatively high proportions of basaltic ingredients

(garnet and clinopyroxene). Partial fusion of lherzolite and extraction of the melt fraction

can leave a solid residue of harzburgite.

Classification diagram for peridotite and pyroxenite, based on proportions of olivine and

pyroxene. The pale green area encompasses the most common compositions of peridotite in the

upper part of the Earth's mantle (partly adapted from Bodinier and Godard (2004)).

Composition

Peridotites are rich in magnesium, reflecting the high proportions of magnesium-rich olivine.

The compositions of peridotites from layered igneous complexes vary widely, reflecting the

relative proportions of pyroxenes, chromite, plagioclase, and amphibole. Minor minerals and

mineral groups in peridotite include plagioclase, spinel (commonly the mineral chromite), garnet

(especially the mineral pyrope), amphibole, and phlogopite. In peridotite, plagioclase is stable at

relatively low pressures (crustal depths), aluminous spinel at higher pressures (to depths of 60

km or so), and garnet at yet higher pressures.

Pyroxenites are related ultramafic rocks, which are composed largely of orthopyroxene and/or

clinopyroxene; minerals that may be present in lesser abundance include olivine, garnet,

plagioclase, amphibole, and spinel.

Distribution and location

Olivine in a peridotite weathering to iddingsite within a mantle xenolith

Peridotite is the dominant rock of the Earth's mantle above a depth of about 400 km; below that

depth, olivine is converted to the higher-pressure mineral wadsleyite. Oceanic plates consist of

up to about 100 km of peridotite covered by a thin crust; the crust, commonly about 6 km thick,

consists of basalt, gabbro, and minor sediments. The peridotite below the ocean crust, "abyssal

peridotite," is found on the walls of rifts in the deep sea floor. Oceanic plates are usually

subducted back into the mantle in subduction zones. However, pieces can be emplaced into or

overthrust on continental crust by a process called obduction, rather than carried down into the

mantle; the emplacement may occur during orogenies, as during collisions of one continent with

another or with an island arc. The pieces of oceanic plates emplaced within continental crust are

referred to as ophiolites; typical ophiolites consist mostly of peridotite plus associated rocks such

as gabbro, pillow basalt, diabase sill-and-dike complexes, and red chert. Other masses of

peridotite have been emplaced into mountain belts as solid masses but do not appear to be related

to ophiolites, and they have been called "orogenic peridotite massifs" and "alpine peridotites."

Peridotites also occur as fragments (xenoliths) carried up by magmas from the mantle. Among

the rocks that commonly include peridotite xenoliths are basalt and kimberlite. Certain volcanic

rocks, sometimes called komatiites, are so rich in olivine and pyroxene that they also can be

termed peridotite. Small pieces of peridotite have even been found in lunar breccias.

The rocks of the peridotite family are uncommon at the surface and are highly unstable, because

olivine reacts quickly with water at typical temperatures of the upper crust and at the Earth's

surface. Many, if not most, surface outcrops have been at least partly altered to serpentinite, a

process in which the pyroxenes and olivines are converted to green serpentine. This hydration

reaction involves considerable increase in volume with concurrent deformation of the original

textures. Serpentinites are mechanically weak and so flow readily within the earth. Distinctive

plant communities grow in soils developed on serpentinite, because of the unusual composition

of the underlying rock. One mineral in the serpentine group, chrysotile, is a type of asbestos.

Morphology and texture

Some peridotites are layered or are themselves layers; others are massive. Many layered

peridotites occur near the base of bodies of stratified gabbroic complexes. Other layered

peridotites occur isolated, but possibly once composed part of major gabbroic complexes. Both

layered and massive peridotites can have any of three principal textures: (1) rather well formed

crystals of olivine separated by other minerals. This probably reflects the original deposition of

olivine sediment from magma. (2) Equigranular crystals with straight grain boundaries

intersecting at about 120°. This may result from slow cooling whereby recrystallization leads to a

minimization of surface energy. (3) Long crystals with ragged curvilinear boundaries. This

probably results from internal deformation.

Many peridotite occurrences have characteristic textures. For example, peridotites with well-

formed olivine crystals occur mainly as layers in gabbroic complexes. "Alpine" peridotites

generally have irregular crystals that occur as more or less serpentinized lenses bounded by faults

in belts of folded mountains such as the Alpines, the Pacific coast ranges, and in the Appalachian

piedmont. Peridotite nodules with irregular equigranular textures are often found in alkaline

basalts and in kimberlite pipes. Some peridotites rich in amphibole have a concentric layered

structure and form parts of plutons called Alaskan-type zoned ultramafic complexes.

Origin

Peridotites have two primary modes of origin, as mantle rocks formed during the accretion and

differentiation of the Earth, or as cumulate rocks formed by precipitation of olivine ± pyroxenes

from basaltic or ultramafic magmas; these magmas are ultimately derived from the upper mantle

by partial melting of mantle peridotites.

Mantle peridotites are sampled as alpine-type massifs in collisional mountain ranges or as

xenoliths in basalt or kimberlite. In all cases these rocks are pyrometamorphic (that is,

metamorphosed in the presence of molten rock) and represent either fertile mantle (lherzolite) or

partially depleted mantle (harzburgite, dunite). Alpine peridotites may be either of the ophiolite

association and representing the uppermost mantle below ocean basins, or masses of

subcontinental mantle emplaced along thrust faults in mountain belts.

Layered peridotites are igneous sediments and form by mechanical accumulation of dense

olivine crystals. Some peridotite forms by precipitation and collection of cumulate olivine and

pyroxene from mantle-derived magmas, such as those of basalt composition. Peridotites

associated with Alaskan-type ultramafic complexes are cumulates that probably formed in the

root zones of volcanoes. Cumulate peridotites are also formed in komatiite lava flows.

Mantle lherzolites may be the principal source rock for basaltic magmas, whereas mantle

harzburgites probably form both from the crystalline residue left after basaltic magma migrates

out of lherzolite and from a crystalline accumulation of early solidification products of some

basaltic magmas within the mantle.

Associated rocks

Komatiites are the rare volcanic equivalent of peridotite.

Eclogite, a rock similar to basalt in composition, is composed primarily of sodic clinopyroxene

and garnet. Eclogite is associated with peridotite in some xenolith occurrences; it also occurs

with peridotite in rocks metamorphosed at high pressures during processes related to subduction.

Pyroxenite

A sample of the orthopyroxenite meteorite ALH84001

Pyroxenite is an ultramafic igneous rock consisting essentially of minerals of the pyroxene

group, such as augite and diopside, hypersthene, bronzite or enstatite. They are classified (see

diagram below) into clinopyroxenites, orthopyroxenites, and the websterites which contain both

pyroxenes. Closely allied to this group are the hornblendites, consisting essentially of hornblende

and other amphiboles.

They are essentially of igneous origin, though some pyroxenites are included in the metamorphic

Lewisian complex of Scotland. The pyroxene-rich rocks which result from the contact

metamorphism of impure limestones are described as pyroxene hornfelses (calc-silicate

hornfelses).

Intrusive and mantle pyroxenites

The igneous pyroxenites are closely allied to the gabbros and norites, from which they differ by

the absence of feldspar, and to the peridotites, which are distinguished from them by containing

more than 40% olivine. This connection is indicated also by their mode of occurrence, for they

usually accompany masses of gabbro and peridotite and seldom are found by themselves.

They are often very coarse-grained, containing individual crystals which may be several inches

in length. The principal accessory minerals, in addition to olivine and feldspar, are chromite and

other spinels, garnet, magnetite, rutile, and scapolite.

Pyroxenites can be formed as cumulates in ultramafic intrusions by accumulation of pyroxene

crystals at the base of the lava chamber. Here they are generally associated with gabbro and

anorthite cumulate layers and are typically high up in the intrusion. They may be accompanied

by magnetite layers, ilmenite layers, but rarely chromite cumulates.

Pyroxenites are also found as layers within masses of peridotite. These layers most commonly

have been interpreted as products of reaction between ascending magmas and peridotite of the

upper mantle. The layers typically are a few centimeters to a meter or so in thickness.

Pyroxenites that occur as xenoliths in basalt and in kimberlite have been interpreted as fragments

of such layers. Although some mantle pyroxenites contain garnet, they are not eclogites, as

clinopyroxene in them is less sodic than omphacite and the pyroxenite compositions typically are

unlike that of basalt. It has been proposed that large volumes of pyroxenite form in the upper

mantle as a result of reaction between peridotite and magma derived from partial melting of

eclogite, and that such pyroxenite volumes are important sources of basalt magma (e.g., Sobolev

and others, 2007).

Pyroxenite lavas

Purely pyroxene-bearing volcanic rocks are rare, restricted to spinifex textured sills, lava tubes

and thick flows in the Archaean greenstone belts. Here, the pyroxenite lavas are created by in-

situ crystallisation and accumulation of pyroxene on the floor of a lava flow, creating the

distinctive spinifex texture, but also occasionally mesocumulate and orthocumulate segregations.

This is in essence similar to the formation of olivine spinifex textures in komatiite lava flows, the

chemistry of the magma differing only to favor crystallisation of pyroxene.

A type locality is the Gullewa Greenstone Belt, in the Murchison region of Western Australia,

and the Duketon Belt near Laverton, where pyroxene spinifex lavas are closely associated with

gold deposits.

Distribution

They frequently occur in the form of dikes or segregations in gabbro and peridotite: in Shetland,

Cortland on the Hudson river, North Carolina (websterite), Baltimore, New Zealand, and in

Saxony.

Classification diagram for peridotite and pyroxenite, based on proportions of olivine and

pyroxene. The pale green area encompasses the most common compositions of peridotite in the

upper part of the Earth's mantle

The pyroxenites are often subject serpentinization under low temperature retrograde

metanorphism and weathering. The rocks are often completely replaced by serpentines, which

sometimes preserve the original structures of the primary minerals, such as the lamination of

hypersthene and the rectangular cleavage of augite. Under pressure-metamorphism hornblende is

developed and various types of amphibolite and hornblende-schist are produced. Occasionally

rocks rich in pyroxene are found as basic facies of nepheline syenite; a good example is provided

by the melanite pyroxenites associated with the borolanite variety found in the Loch Borralan

igneous complex of Scotland.

Dunite

Small volcanic bomb of (black) basanite with (green) dunite

Dunite (pronounced /ˈdʌnaɪt/ or /ˈdjuːnaɪt/) is an igneous, plutonic rock, of ultramafic

composition, with coarse-grained or phaneritic texture. The mineral assemblage is greater than

90% olivine, with minor amounts of other minerals such as pyroxene, chromite and pyrope.

Dunite is the olivine-rich end-member of the peridotite group of mantle-derived rocks. Dunite

and other peridotite rocks are considered the major constituents of the Earth's mantle above a

depth of about 400 kilometers. Dunite is rarely found within continental rocks, but where it is

found, it typically occurs at the base of ophiolite sequences where slabs of mantle rock from a

subduction zone have been thrust onto continental crust by obduction during continental or

island arc collisions (orogeny). It is also found in alpine peridotite massifs that represent slivers

of sub-continental mantle exposed during collisional orogeny. Dunite typically undergoes

retrograde metamorphism in near-surface environments and is altered to serpentinite and

soapstone.

Dunite may represent the refractory residue left after the extraction of basaltic magmas in the

upper mantle. This is the type of dunite found in the lowermost parts of ophiolites, alpine

peridotite massifs, and xenoliths. Dunite may also form by the accumulation of olivine crystals

on the floor of large basaltic or picritic magma chambers. These "cumulate" dunites typically

occur in thick layers in layered intrusions, associated with cumulate layers of wehrlite, olivine

pyroxenite, harzburgite, and even chromitite (a cumulate rock consisting largely of chromite).

Small layered intrusions may be of any geologic age, for example, the Triassic Palisades Sill in

New York and the larger Eocene Skaergaard complex in Greenland. The largest layered mafic

intrusions are tens of kilometers in size and almost all are Proterozoic in age, e.g., the Stillwater

igneous complex (Montana), the Muskox intrusion (Canada), and the Great Dyke (Zimbabwe).

Cumulate dunite may also be found in ophiolite complexes, associated with layers of wehrlite,

pyroxenite, and gabbro.

Dunite was named by the Austrian geologist, Ferdinand von Hochstetter in 1859 after Dun

Mountain near Nelson, New Zealand. Dun Mountain was given its name because of the dun

colour of the underlying ultramafic rocks. This color results from surface weathering that

oxidizes the iron in olivine in temperate climates (weathering in tropical climates creates a deep

red soil). Dun Mountain is separated from its sister massif, Red Mountain, at the southern end of

South Island, New Zealand, by the Alpine Fault, an approximately 600 km long right lateral

strike slip fault similar to the San Andreas fault in California.

A massive exposure of dunite in the United States can be found as the Twin Sisters Peaks, near

Mt. Baker, in the northern Cascade Mountains of Washington State.

Dunite could be used to sequester CO2 and help mitigate global climate change via accelerated

rock weathering. This would involve spreading large quantities of finely ground dunite in

tropical regions known near sources of dunite. One significant environmental side effect would

be a significant increase the pH of nearby rivers.

Hornblendite

Hornblendite from Poland

Hornblendite is a plutonic rock consisting mainly of the amphibole hornblende. Hornblende

rich ultramafic rocks are rare and when hornblende is the dominant mineral phase they are

classified as hornblendites with qualifiers such as garnet hornblendite identifying a second

abundant contained mineral.

Metamorphic rocks composed dominantly of amphiboles are referred to as amphibolites

Kimberlite

Kimberlite from U.S.A.

QEMSCAN mineral map of kimberlite from South Africa

Kimberlite is a type of potassic volcanic rock best known for sometimes containing diamonds. It

is named after the town of Kimberley in South Africa, where the discovery of an 83.5-carat (16.7

g) diamond in 1871 spawned a diamond rush, eventually creating the Big Hole.

Kimberlite occurs in the Earth's crust in vertical structures known as kimberlite pipes. Kimberlite

pipes are the most important source of mined diamonds today. The consensus on kimberlites is

that they are formed deep within the mantle. Formation occurs at depths between 150 and 450

kilometres (93 and 280 mi), from anomalously enriched exotic mantle compositions, and are

erupted rapidly and violently, often with considerable carbon dioxide and other volatile

components. It is this depth of melting and generation which makes kimberlites prone to hosting

diamond xenocrysts.

Kimberlite has attracted more attention than its relative volume might suggest that it deserves.

This is largely because it serves as a carrier of diamonds and garnet peridotite mantle xenoliths to

the Earth's surface. Its probable derivation from depths greater than any other igneous rock type,

and the extreme magma composition that it reflects in terms of low silica content and high levels

of incompatible trace element enrichment, make an understanding of kimberlite petrogenesis

important. In this regard, the study of kimberlite has the potential to provide information on the

composition of the deep mantle, and melting processes occurring at or near the interface between

the cratonic continental lithosphere and the underlying convecting asthenospheric mantle.

Morphology and volcanology

Kimberlites occur as carrot-shaped, vertical intrusions termed 'pipes'. This classic carrot shape is

formed due to a complex intrusive process of kimberlitic magma which inherits a large

proportion of both CO2 and H2O in the system, which produces a deep explosive boiling stage

that causes a significant amount of vertical flaring (Bergman, 1987). Kimberlite classification is

based on the recognition of differing rock facies. These differing facies are associated with a

particular style of magmatic activity, namely crater, diatreme and hypabyssal rocks (Clement and

Skinner 1985, and Clement, 1982).

The morphology of kimberlite pipes, and the classical carrot shape, is the result of explosive

diatreme volcanism from very deep mantle-derived sources. These volcanic explosions produce

vertical columns of rock that rise from deep magma reservoirs. The morphology of kimberlite

pipes is varied but generally includes a sheeted dyke complex of tabular, vertically dipping

feeder dykes in the root of the pipe which extends down to the mantle. Within 1.5–2 km (0.93–

1.2 mi) of the surface, the highly pressured magma explodes upwards and expands to form a

conical to cylindrical diatreme, which erupts to the surface. The surface expression is rarely

preserved, but is usually similar to a maar volcano. The diameter of a kimberlite pipe at the

surface is typically a few hundred meters to a kilometer (up to 0.6 mile).

Two Jurassic kimberlite dikes exist in Pennsylvania. One, the Gates-Adah Dike, outcrops on the

Monongahela River on the border of Fayette and Greene Counties. The other, the Dixonville-

Tanoma Dike in central Indiana County, does not outcrop at the surface and was discovered by

miners.[1]

Petrology

Both the location and origin of kimberlitic magmas are areas of contention. Their extreme

enrichment and geochemistry has led to a large amount of speculation about their origin, with

models placing their source within the sub-continental lithospheric mantle (SCLM) or even as

deep as the transition zone. The mechanism of enrichment has also been the topic of interest with

models including partial melting, assimilation of subducted sediment or derivation from a

primary magma source.

Historically, kimberlites have been subdivided into two distinct varieties termed 'basaltic' and

'micaceous' based primarily on petrographic observations (Wagner, 1914). This was later revised

by Smith (1983) who re-named these divisions Group I and Group II based on the isotopic

affinities of these rocks using the Nd, Sr and Pb systems. Mitchell (1995) later proposed that

these group I and II kimberlites display such distinct differences, that they may not be as closely

related as once thought. He showed that Group II kimberlites actually show closer affinities to

lamproites than they do to Group I kimberlites. Hence, he reclassified Group II kimberlites as

orangeites to prevent confusion.

Group I kimberlites

Group-I kimberlites are of CO2-rich ultramafic potassic igneous rocks dominated by a primary

mineral assemblage of forsteritic olivine, magnesian ilmenite, chromium pyrope, almandine-

pyrope, chromium diopside (in some cases subcalcic), phlogopite, enstatite and of Ti-poor

chromite. Group I kimberlites exhibit a distinctive inequigranular texture caused by macrocrystic

(0.5–10 mm, 0.020–0.39 in) to megacrystic (10–200 mm, 0.39–7.9 in) phenocrysts of olivine,

pyrope, chromian diopside, magnesian ilmenite and phlogopite, in a fine to medium grained

groundmass.

The groundmass mineralogy, which more closely resembles a true composition of the igneous

rock, contains forsteritic olivine, pyrope garnet, Cr-diopside, magnesian ilmenite and spinel.

Group II kimberlites

Group-II kimberlites (or orangeites) are ultrapotassic, peralkaline rocks rich in volatiles

(dominantly H2O). The distinctive characteristic of orangeites is phlogopite macrocrysts and

microphenocrysts, together with groundmass micas that vary in composition from phlogopite to

"tetraferriphlogopite" (anomalously Fe-rich phlogopite). Resorbed olivine macrocrysts and

euhedral primary crystals of groundmass olivine are common but not essential constituents.

Characteristic primary phases in the groundmass include: zoned pyroxenes (cores of diopside

rimmed by Ti-aegirine); spinel-group minerals (magnesian chromite to titaniferous magnetite);

Sr- and REE-rich perovskite; Sr-rich apatite; REE-rich phosphates (monazite, daqingshanite);

potassian barian hollandite group minerals; Nb-bearing rutile and Mn-bearing ilmenite.

Kimberlitic indicator minerals

Kimberlites are peculiar igneous rocks because they contain a variety of mineral species with

peculiar chemical compositions. These minerals such as potassic richterite, chromian diopside (a

pyroxene), chromium spinels, magnesian ilmenite, and garnets rich in pyrope plus chromium, are

generally absent from most other igneous rocks, making them particularly useful as indicators for

kimberlites.

These indicator minerals are generally sought in stream sediments in modern alluvial material.

Their presence may indicate the presence of a kimberlite within the erosional watershed which

produced the alluvium.

Volcanic rock

.

Ignimbrite is a deposit of a pyroclastic flow.

Volcanic rock is an igneous rock of volcanic origin.

Texture

Photomicrograph of a volcanic lithic fragment (sand grain); upper picture is plane-polarized

light, bottom picture is cross-polarized light, scale box at left-center is 0.25 millimeter.

Volcanic rocks are usually fine-grained or aphanitic to glass in texture. They often contain clasts

of other rocks and phenocrysts. Phenocrysts are crystals that are larger than the matrix and are

identifiable with the unaided eye. Rhomb porphyry is an example with large rhomb shaped

phenocrysts embedded in a very fine grained matrix.

Volcanic rocks often have a vesicular texture caused by voids left by volatiles escaping from the

molten lava. Pumice is an example of explosive volcanic eruption. It is so vesicular that it floats

in water.

Naming

Vesicular olivine basalt from La Palma (green phenocrysts are olivine).

Volcanic rocks are named according to both their chemical composition and texture. Basalt is a

very common volcanic rock with low silica content. Rhyolite is a volcanic rock with high silica

content. Rhyolite has silica content similar to that of granite while basalt is compositionally

equal to gabbro. Intermediate volcanic rocks include andesite, dacite, trachyte, and latite.

Pyroclastic rocks are the product of explosive volcanism. They are often felsic (high in silica).

Pyroclastic rocks are often the result of volcanic debris, such as ash, bombs and tephra, and other

volcanic ejecta. Examples of pyroclastic rocks are tuff and ignimbrite.

Shallow intrusions, which possess structure similar to volcanic rather than plutonic rocks are also

considered to be volcanic.

Composition of volcanic rocks

ʻAʻā next to pāhoehoe lava at the Craters of the Moon National Monument and Preserve, Idaho,

United States.

The sub-family of rocks that form from volcanic lava are called igneous volcanic rocks (to

differentiate them from igneous rocks that form from magma below the surface, called igneous

plutonic rocks).

The lavas of different volcanoes, when cooled and hardened, differ much in their appearance and

composition. If a rhyolite lava-stream cools quickly, it can quickly freeze into a black glassy

substance called obsidian. When filled with bubbles of gas, the same lava may form the spongy

mineral pumice. Allowed to cool slowly, it forms a light-colored, uniformly solid rock called

rhyolite.

The lavas, having cooled rapidly in contact with the air or water, are mostly finely crystalline or

have at least fine-grained ground-mass representing that part of the viscous semi-crystalline lava

flow that was still liquid at the moment of eruption. At this time they were exposed only to

atmospheric pressure, and the steam and other gases, which they contained in great quantity were

free to escape; many important modifications arise from this, the most striking being the frequent

presence of numerous steam cavities (vesicular structure) often drawn out to elongated shapes

subsequently filled up with minerals by infiltration (amygdaloidal structure).

As crystallization was going on while the mass was still creeping forward under the surface of

the Earth, the latest formed minerals (in the ground-mass) are commonly arranged in subparallel

winding lines that follow the direction of movement (fluxion or fluidal structure)—and larger

early minerals that previously crystallized may show the same arrangement. Most lavas fall

considerably below their original temperatures before emitted. In their behavior, they present a

close analogy to hot solutions of salts in water, which, when they approach the saturation

temperature, first deposit a crop of large, well-formed crystals (labile stage) and subsequently

precipitate clouds of smaller less perfect crystalline particles (metastable stage).

In igneous rocks the first generation of crystals generally forms before the lava has emerged to

the surface, that is to say, during the ascent from the subterranean depths to the crater of the

volcano. It has frequently been verified by observation that freshly emitted lavas contain large

crystals borne along in a molten, liquid mass. The large, well-formed, early crystals

(phenocrysts) are said to be porphyritic; the smaller crystals of the surrounding matrix or ground-

mass belong to the post-effusion stage. More rarely lavas are completely fused at the moment of

ejection; they may then cool to form a non-porphyritic, finely crystalline rock, or if more rapidly

chilled may in large part be non-crystalline or glassy (vitreous rocks such as obsidian, tachylyte,

pitchstone).

A common feature of glassy rocks is the presence of rounded bodies (spherulites), consisting of

fine divergent fibres radiating from a center; they consist of imperfect crystals of feldspar, mixed

with quartz or tridymite; similar bodies are often produced artificially in glasses that are allowed

to cool slowly. Rarely these spherulites are hollow or consist of concentric shells with spaces

between (lithophysae). Perlitic structure, also common in glasses, consists of the presence of

concentric rounded cracks owing to contraction on cooling.

Volcanic rocks, Porto Moniz, Madeira

The phenocrysts or porphyritic minerals are not only larger than those of the ground-mass; as the

matrix was still liquid when they formed they were free to take perfect crystalline shapes,

without interference by the pressure of adjacent crystals. They seem to have grown rapidly, as

they are often filled with enclosures of glassy or finely crystalline material like that of the

ground-mass . Microscopic examination of the phenocrysts often reveals that they have had a

complex history. Very frequently they show layers of different composition, indicated by

variations in color or other optical properties; thus augite may be green in the center surrounded

by various shades of brown; or they may be pale green centrally and darker green with strong

pleochoism (aegirine) at the periphery.

In the feldspars the center is usually richer in calcium than the surrounding layers, and successive

zones may often be noted, each less calsic than those within it. Phenocrysts of quartz (and of

other minerals), instead of sharp, perfect crystalline faces, may show rounded corroded surfaces,

with the points blunted and irregular tongue-like projections of the matrix into the substance of

the crystal. It is clear that after the mineral had crystallized it was partly again dissolved or

corroded at some period before the matrix solidified.

Corroded phenocrysts of biotite and hornblende are very common in some lavas; they are

surrounded by black rims of magnetite mixed with pale green augite. The hornblende or biotite

substance has proved unstable at a certain stage of consolidation, and has been replaced by a

paramorph of augite and magnetite, which may partially or completely substitute for the original

crystal but still retains its characteristic outlines.

Felsite

Felsite covered with dendritic pyrolusite formations.

Felsite (also called felstone [1]) is a very fine grained volcanic rock that may or may not contain

larger crystals. Felsite is a field term for a light colored rock that typically requires petrographic

examination or chemical analysis for more precise definition. Color is generally white through

light gray, or red to tan and may include any color except dark gray, green or black (the colors of

traprock).[1]

The mass of the rock consists of a fine-grained matrix of felsic materials,

particularly quartz, sodium and potassium feldspar, and may be termed a quartz felsite or quartz

porphyry if the quartz phenocrysts are present. This rock is typically of volcanic origin, and may

be found in association with obsidian and rhyolite. In some cases, it is sufficiently fine-grained

for use in making stone tools.

Rhyolite

Rhyolite

— Igneous Rock —

Composition

Felsic: igneous quartz and alkali feldspar (orthoclase,

sanidine and sodic plagioclase), biotite and hornblende.

This page is about a volcanic rock. For the ghost town see Rhyolite, Nevada, and for the satellite

system, see Rhyolite/Aquacade.

Rhyolite is an igneous, volcanic (extrusive) rock, of felsic (silica-rich) composition (typically >

69% SiO2 — see the TAS classification). It may have any texture from glassy to aphanitic to

porphyritic. The mineral assemblage is usually quartz, alkali feldspar and plagioclase (in a ratio

> 1:2 — see the QAPF diagram). Biotite and hornblende are common accessory minerals.

Rocks from the Bishop tuff, uncompressed with pumice on left; compressed with fiamme on

right.

Geology

Rhyolite can be considered as the extrusive equivalent to the plutonic granite rock, and

consequently, outcrops of rhyolite may bear a resemblance to granite. Due to their high content

of silica and low iron and magnesium contents, rhyolite melts are highly polymerized and form

highly viscous lavas. They can also occur as breccias or in volcanic plugs and dikes. Rhyolites

that cool too quickly to grow crystals form a natural glass or vitrophyre, also called obsidian.

Slower cooling forms microscopic crystals in the lava and results in textures such as flow

foliations, spherulitic, nodular, and lithophysal structures. Some rhyolite is highly vesicular

pumice. Many eruptions of rhyolite are highly explosive and the deposits may consist of fallout

tephra/tuff or of ignimbrites.

History

During the second millennium BC, rhyolite was quarried extensively in what is now eastern

Pennsylvania in the United States. Among the leading quarries was the Carbaugh Run Rhyolite

Quarry Site in Adams County, where as many as fifty small quarry pits are known.[1]

Name

Top stone is obsidian (vitrophyre), below that is pumice and in lower right corner is rhyolite

(light color)

A sample of Rhyolite from the Conical Hill dome at the head Lyttelton Harbour, Banks

Peninsula, New Zealand

The name rhyolite was introduced into science by the German traveler and geologist Ferdinand

von Richthofen after his explorations in the Rocky Mountains in the 1860s.

Obsidian

.

Obsidian

Obsidian from Lake County, Oregon

General

Category Volcanic glass

Chemical formula

70–75% SiO2,

plus MgO, Fe3O4

Identification

Color Black, gray, dark green, red, yellow,

pink

Fracture Conchoidal

Mohs scale

hardness ~ 5 to 5.5

Luster Vitreous

Specific gravity ~ 2.5

Optical properties Translucent

Obsidian is a naturally occurring volcanic glass formed as an extrusive igneous rock.

It is produced when felsic lava extruded from a volcano cools rapidly without crystal growth.

Obsidian is commonly found within the margins of rhyolitic lava flows known as obsidian

flows, where the chemical composition (high silica content) induces a high viscosity and

polymerization degree of the lava. The inhibition of atomic diffusion through this highly viscous

and polymerized lava explains the lack of crystal growth. Because of this lack of crystal

structure, obsidian blade edges can reach almost molecular thinness, leading to its ancient use as

projectile points and blades, and its modern use as surgical scalpel blades.[1][2]

Origin and properties

Pliny's Natural History features volcanic glass called "Obsidianus", so named from its

resemblance to a stone found in Ethiopia by one Obsius.[3]

Obsidian is mineral-like, but not a true mineral because as a glass it is not crystalline; in addition,

its composition is too complex to comprise a single mineral. It is sometimes classified as a

mineraloid. Though obsidian is dark in color similar to mafic rocks such as basalt, obsidian's

composition is extremely felsic. Obsidian consists mainly of SiO2 (silicon dioxide), usually 70%

or more. Crystalline rocks with obsidian's composition include granite and rhyolite. Because

obsidian is metastable at the Earth's surface (over time the glass becomes fine-grained mineral

crystals), no obsidian has been found that is older than Cretaceous age. This breakdown of

obsidian is accelerated by the presence of water. Obsidian has low water content when fresh,

typically less than 1% water by weight,[4]

but becomes progressively hydrated when exposed to

groundwater, forming perlite. Tektites were once thought by many to be obsidian produced by

lunar volcanic eruptions, though few scientists now adhere to this hypothesis.

Pure obsidian is usually dark in appearance, though the color varies depending on the presence of

impurities. Iron and magnesium typically give the obsidian a dark green to brown to black color.

Very few samples are nearly colorless. In some stones, the inclusion of small, white, radially

clustered crystals of cristobalite in the black glass produce a blotchy or snowflake pattern

(snowflake obsidian). It may contain patterns of gas bubbles remaining from the lava flow,

aligned along layers created as the molten rock was flowing before being cooled. These bubbles

can produce interesting effects such as a golden sheen (sheen obsidian) or an iridescent,

rainbow-like sheen (rainbow obsidian).

Glass Mountain, a large

obsidian flow at Medicine

Lake Volcano

Counterclockwise from top:

obsidian, pumice and

rhyolite (light color)

Snowflake

obsidian

Rainbow obsidian

Occurrence

Obsidian can be found in locations which have experienced rhyolitic eruptions. It can be found in

Argentina, Armenia, Canada, Chile, Greece, Guatemala, Iceland, Italy, Japan, Kenya, Mexico,

New Zealand, Peru, Scotland and United States. Obsidian flows which may be hiked on are

found within the calderas of Newberry Volcano and Medicine Lake Volcano in the Cascade

Range of western North America, and at Inyo Craters east of the Sierra Nevada in California.

Yellowstone National Park has a mountainside containing obsidian located between Mammoth

Hot Springs and the Norris Geyser Basin, and deposits can be found in many other western U.S.

states including Arizona, Colorado, New Mexico, Texas, Utah, Washington,[5]

Oregon[6]

and

Idaho. Obsidian can also be found in the eastern U.S. state of Virginia.

Obsidian arrowhead.

Historical use

Obsidian was valued in Stone Age cultures because, like flint, it could be fractured to produce

sharp blades or arrowheads. Like all glass and some other types of naturally occurring rocks,

obsidian breaks with a characteristic conchoidal fracture. It was also polished to create early

mirrors.

Modern archaeologists have developed a relative dating system, obsidian hydration dating, to

calculate the age of obsidian artifacts.

Middle East

In Ubaid in the 5th millennium BC, blades were manufactured from obsidian mined in today's

Turkey.[7]

Obsidian talus at Obsidian Dome, California.

Americas

Lithic analysis can be instrumental in understanding prehispanic groups in Mesoamerica. A

careful analysis of obsidian in a culture or place can be of considerable use to reconstruct

commerce, production, distribution and thereby understand economic, social and political

aspects of a civilization. This is the case in Yaxchilán, a Maya city where even warfare

implications have been studied linked with obsidian use and its debris.[8]

Another example is the

archeological recovery at coastal Chumash sites in California indicating considerable trade with

the distant site of Casa Diablo, California in the Sierra Nevada Mountains.[9]

Pre-Columbian Mesoamericans' use of obsidian was extensive and sophisticated; including

carved and worked obsidian for tools and decorative objects. Mesoamericans also made a type of

sword with obsidian blades mounted in a wooden body. Called a macuahuitl, the weapon was

capable of inflicting terrible injuries, combining the sharp cutting edge of an obsidian blade with

the ragged cut of a serrated weapon.

Native American people traded obsidian throughout the Americas. Each volcano and in some

cases each volcanic eruption produces a distinguishable type of obsidian, making it possible for

archaeologists to trace the origins of a particular artifact. Similar tracing techniques have allowed

obsidian to be identified in Greece also as coming from Melos, Nisyros or Yiali, islands in the

Aegean Sea. Obsidian cores and blades were traded great distances inland from the coast.[citation

needed]

In Chile obsidian tools from Chaitén Volcano have been found as far away as in Chan-Chan

400 km north of the volcano and also in sites 400 km south of it.[10][11]

Pitchstone

Pitchstone ridge: An Sgurr, Isle of Eigg, Scotland

Pitchstone is a dull black glassy volcanic rock formed when viscous lava or magma cools

swiftly. It is similar to but coarser than obsidian. It is a volcanic glass with a conchoidal fracture

(like glass), a resinous lustre, and a variable composition. Its colour may be mottled, streaked, or

uniform brown, red, green, gray, or black. It is an extrusive rock that is very resistant to erosion.

The ridge of An Sgurr on the Isle of Eigg was originally formed as a lava flow in a valley.

Pumice

Specimen of highly porous pumice from Teide volcano on Tenerife, Canary Islands. Density of

specimen approx 0.25 g/cm³. Scale is in centimeters.

Pumice (pronounced /ˈpʌməs/ ) is a textural term for a volcanic rock that is a solidified frothy

lava typically created when super-heated, highly pressurized rock is violently ejected from a

volcano. It can be formed when lava and water are mixed. This unusual formation is due to the

simultaneous actions of rapid cooling and rapid depressurization. The depressurization creates

bubbles by lowering the solubility of gases (including water and CO2) dissolved in the lava, so

that they rapidly exsolve (like the bubbles of CO2 that appear when a carbonated drink is

opened). The simultaneous cooling then freezes the bubbles in the matrix.

Properties

Illustrates the porous nature in detail

Rocks from the Bishop tuff, uncompressed with pumice on left; compressed with fiamme on

right.

A 10 centimeter (6 inch) piece of pumice supported by a rolled-up U.S. 20-dollar bill

demonstrates its very low density.

Pumice is composed of highly microvesicular glass pyroclastic with very thin, translucent bubble

walls of extrusive igneous rock. It is commonly, but not exclusively of silicic or felsic to

intermediate in composition (e.g., rhyolitic, dacitic, andesite, pantellerite, phonolite, trachyte),

but basaltic and other compositions are known. Pumice is commonly pale in color, ranging from

white, cream, blue or grey, to green-brown or black. It forms when volcanic gases exsolving

from viscous magma nucleate bubbles which cannot readily decouple from the viscous magma

prior to chilling to glass. Pumice is a common product of explosive eruptions (plinian and

ignimbrite-forming) and commonly forms zones in upper parts of silicic lavas. Pumice has an

average porosity of 90%, and initially floats on water.

Scoria differs from pumice in being denser, with larger vesicles and thicker vesicle walls; it sinks

rapidly. The difference is the result of the lower viscosity of the magma that forms scoria. When

larger amounts of gas are present, the result is a finer-grained variety of pumice known as

pumicite. Pumice is considered a glass because it has no crystal structure. Pumice varies in

density according to the thickness of the solid material between the bubbles; many samples float

in water. After the explosion of Krakatoa, rafts of pumice drifted through the Pacific Ocean for

up to 20 years, with tree trunks floating among them.[1]

In fact, pumice rafts disperse and support

several marine species.[2]

In 1979, 1984 and 2006, underwater volcanic eruptions near Tonga

created large pumice rafts, some as large as 30 kilometres that floated hundreds of kilometres to

Fiji.[3]

There are two main forms of vesicles. Most pumice contains tubular microvesicles that can

impart a silky or fibrous fabric. The elongation of the microvesicles occurs due to ductile

elongation in the volcanic conduit or, in the case of pumiceous lavas, during flow. The other

form of vesicles are subspherical to spherical and result from high vapor pressure during

eruption.

Scoria

Scoria of various hues exists on Mount Tarawera, from its 1886 eruption.

Scoria

Holocene scoria-producing volcano near Veyo, Utah

Tuff moai with red scoria pukao on its head

Fresh scoria sometimes has a blue sheen to its surface.

Photomicrograph of a volcanic lithic fragment (sand grain) derived from scoria; upper picture is

plane-polarized light, bottom picture is cross-polarized light, scale box at left-center is 0.25

millimeter.

Scoria is a volcanic rock containing many holes or vesicules. It is most generally dark in color

(generally dark brown, black or red), and basaltic or andesitic in composition. Scoria is relatively

light as a result of its numerous macroscopic ellipsoidal vesicles, but in contrast to pumice, all

scoria has a specific gravity greater than 1, and sinks in water. The holes or vesicules form when

gases that were dissolved in the magma come out of solution as it erupts, creating bubbles in the

molten rock, some of which are frozen in place as the rock chills and solidifies. Scoria may form

as part of a lava flow, typically near its surface, or as fragmental ejecta (lapilli, blocks and

bombs), for instance in Strombolian eruptions that form steep-sided scoria cones. Most scoria is

composed of glassy fragments, and may contain phenocrysts.

The word scoria comes from the Greek σκωρία, skōria, rust. An old name for scoria is cinder.

Comparisons

Scoria differs from pumice, another vesicular volcanic rock, in having larger vesicles and thicker

vesicle walls, and hence is denser. The difference is probably the result of lower magma

viscosity, allowing rapid volatile diffusion, bubble growth, coalescence, and bursting.

Formation

As rising magma encounters lower pressures, dissolved gases are able to exsolve and form

vesicles. Some of the vesicles are trapped when the magma chills and solidifies. Vesicles are

usually small, spheroidal and do not impinge upon one another; instead they open into one

another with little distortion.

Volcanic cones of scoria can be left behind after eruptions, usually forming mountains with a

crater at the summit. An example is Mount Wellington, Auckland in New Zealand, which like

the Three Kings in the south of the same city has been extensively quarried. Quincan, a unique

form of Scoria, is quarried at Mount Quincan in Far North Queensland, Australia.

Trachyte

A sample of trachyte

Trachyte is an igneous, volcanic rock with an aphanitic to porphyritic texture. The mineral

assemblage consists of essential alkali feldspar; relatively minor plagioclase and quartz or a

feldspathoid such as nepheline may also be present. (See the QAPF diagram). Biotite,

clinopyroxene and olivine are common accessory minerals.

Chemically, trachyte contains less SiO2 than rhyolite and more (Na2O plus K2O) than dacite.

These chemical differences are consistent with the position of trachyte in the TAS classification,

and they account for the feldspar-rich mineralogy of the rock type.

Trachytes usually consist mainly of sanidine feldspar. Very often they have minute irregular

steam cavities which make the broken surfaces of specimens of these rocks rough and irregular,

and from this character they have derived their name. It was first given to certain rocks of this

class from Auvergne, and was long used in a much wider sense than that defined above, in fact it

included quartz-trachytes (now known as liparites and rhyolites) and oligoclase-trachytes, which

are now more properly assigned to andesites. The trachytes are often described as being the

volcanic equivalents of the plutonic syenites. Their dominant mineral, sanidine feldspar, very

commonly occurs in two generations, i.e. both as large well-shaped porphyritic crystals and in

smaller imperfect rods or laths forming a finely crystalline groundmass. With this there is

practically always a smaller amount of plagioclase, usually oligoclase; but the potash felspar

(sanidine) often contains a considerable proportion of the sodium feldspar (albite), and has rather

the characteristics of anorthoclase or cryptoperthite than of pure sanidine. Rhomb porphyry is an

example with usually large porphyritic rhomb shaped phenocrysts embedded in a very fine

grained matrix.

Quartz is typically rare in trachyte, but tridymite (which likewise consists of silica) is by no

means uncommon. It is rarely in crystals large enough to, be visible without the aid of the

microscope, but in thin sections it may appear as small hexagonal plates, which overlap and form

dense aggregates, like a mosaic or like the tiles on a roof. They often cover the surfaces of the

larger feldspars or line the steam cavities of the rock, where they may be mingled with

amorphous opal or fibrous chalcedony. In the older trachytes, secondary quartz is not rare, and

probably sometimes results from the recrystallization of tridymite.

Of the mafic minerals present, augite is the most common. It is usually of pale green color, and

its small crystals are often very perfect in form. Brown hornblende and biotite occur also, and are

usually surrounded by black corrosion borders composed of magnetite and pyroxene; Sometimes

the replacement is complete and no bornblende or biotite is left, though the outlines of the cluster

of magnetite and augite may clearly indicate from which of these minerals it was derived.

Olivine is unusual, though found in some trachytes, like those of the Arso in Isthia. Basic

varieties of plagioclase, such as labradorite, are known also as phenocrysts in some Italian

trachytes. Dark brown varieties of augite and rhombic pyroxene (hypersthene or bronzite) have

been observed but are not common. Apatite, zircon and magnetite are practically always present

as accessory minerals.

The trachytes being very rich in potash feldspar, necessarily contain considerable amounts of

alkali; in this character they approach the phonolites. Occasionally minerals of the feldspathoid

group, such as nepheline, sodalite and leucite, occur, and rocks of this kind are known as

phonolitic trachytes. The sodium-bearing amphiboles and pyroxenes so characteristic of the

phonolites may also be found in some trachytes; thus aegirine or aegirine augite forms

outgrowths on diopside crystals, and riebeckite may be present in spongy growths among the

feldspars of the groundmass (as in the trachyte of Berkum on the Rhine). Trachytic rocks are

typically porphyritic, and some of the best known examples, such as the trachyte of Drachenfels

on the Rhine, show this character excellently, having large sanidine crystals of tabular form an

inch or two in length scattered through their fine-grained groundmass. In many trachytes,

however, the phenocrysts are few and small, and the groundmass comparatively coarse. The

ferromagnesian minerals rarely occur in large crystals, and are usually not conspicuous in hand

specimens of these rocks. Two types of groundmass are generally recognized: the trachytic,

composed mainly of long, narrow, subparallel rods of sanidine, and the orthophyric, consisting

of small, squarish or rectangular prisms of the same mineral. Sometimes granular augite or

spongy riebeckite occurs in the groundmass, but as a rule this part of the rock is highly

feldspathic. Glassy forms of trachyte (obsidian) occur, as in Iceland, and pumiceous varieties are

known (in Teneriffe and elsewhere), but these rocks as contrasted with the rhyolites have a

remarkably strong tendency to crystallize, and are rarely to any considerable extent vitreous.

A polished opal on trachyte

Trachytes are well represented among the Tertiary and recent volcanic rocks of Europe. In

Britain they occur in Skye as lava flows and as dikes or intrusions, but they are much more

common on the continent of Europe, as in the Rhine district and the Eifel, also in Auvergne,

Bohemia and the Euganean Hills. In the neighborhoord of Rome, Naples and the island of Ischia

trachytic lavas and tuffs are of common occurrence. In the United States trachytes are less

frequent, being known in South Dakota (Black Hills). In Iceland, the Azores, Teneriffe and

Ascension there are recent trachytic lavas, and rocks of this kind occur also in New South Wales

(Cambewarra range), East Africa, Madagascar, Aden and in many other districts.

Among the older volcanic rocks trachytes also are not scarce, though they have often been

described under the names orthophyre and orthoclase-porphyry, while trachyte was reserved for

Tertiary and recent rocks of similar composition. In England there are Permian trachytes in the

Exeter district, and Carboniferous trachytes are found in many parts of the central valley of

Scotland. The latter differ in no essential respect from their modern representatives in Italy and

the Rhine valley, but their augite and biotite are often repiaced by chlorite and other secondary

products. Permian trachytes occur also in Thuringia and the Saar district in Germany.

Closely allied to the trachytes are the keratophyres, which occur mainly in Palaeozoic strata in

the Harz (Germany), in the Southern Uplands of Scotland, in Cornwall, etc. They are usually

porphyritic and fluidal; and consist mainly of alkali feldspar (anorthoclase principally, but also

albite and orthoclase), with a small quantity of chlorite and iron oxides.

Phonolite

Aegirine phonolite. Dark prismatic minerals are aegirine phenocrysts.

Phonolite is a rare igneous, volcanic (extrusive) rock of intermediate (between felsic and mafic)

composition, with aphanitic to porphyritic texture.

The name phonolite comes from the Greek meaning (more or less) "sounding stone" because of

the metallic sound it produces if an unfractured plate is hit, hence the English name clinckstone.

Genesis

Phonolite is unusual in that it forms from a highly silica undersaturated melt by low degrees of

partial melting (less than 10%) of highly aluminous lower crustal rocks such as tonalite,

monzonite and metamorphic rocks. Melting of such rocks to a very low degree promotes the

liberation of aluminium, potassium, sodium and calcium via melting of feldspar, with some

involvement of mafic minerals. The melt formed is silica undersaturated (ie; quartz is absent

from the melts or solidified rocks), with feldspathoid species dominating over feldspar species in

the melt.

Phonolite occurrences are associated with a few geological processes and tectonic events, which

can lead to the melting of appropriate precursor lithologies. These include intracontinental

hotspot volcanism, such as may form above mantle plumes covered by thick continental crust. A-

type granites and alkaline igneous provinces are usually associated with phonolites. Phonolites

may also be produced by low degree partial melting of underplates of granitic material in

collisional orogenic belts.

Mineralogy

Phonolites, as they are products of low degree partial melts, are silica undersaturated, and have

feldspathoids in their normative mineralogy.

Mineral assemblages in phonolite occurrences are usually abundant feldspathoids (nepheline,

sodalite, hauyne, leucite and analcite) and alkali feldspar (sanidine, anorthoclase or orthoclase),

and rare sodic plagioclase. Biotite, sodium rich amphiboles and pyroxenes along with iron rich

olivine are common minor minerals. Accessory phases include titanite, apatite, corundum.

zircon, magnetite and ilmenite.[1]

Phonolites are silica under-saturated, as illustrated by the

position of phonolite in the TAS classification and QAPF diagrams.

Phonolite is a fine-grained equivalent of nepheline syenite, and the genesis of such magmas is

discussed in the treatment of that rock type.

Occurrence

Nepheline syenites and phonolites occur widely distributed throughout the world[2]

in Canada,

Norway, Greenland, Sweden, the Ural Mountains, the Pyrenees, Italy, Brazil,the Transvaal

region, and Magnet Cove igneous complex of Arkansas, as well as on oceanic islands (eg;

Canary Islands[3]

).

Nepheline-normative rocks occur in close association with the Bushveld Igneous Complex,

possibly formed from partial melting of the wall rocks adjacent to that large ultramafic layered

intrusion.

Examples

Devil's Tower, a striking example of columnar jointed phonolite.

Dunedin, New Zealand[4]

Hoodoo Mountain, northwestern British Columbia, Canada.

Jebel Nefusa, Libya[5]

Teide, a stratovolcano on the island of Tenerife.[6]

Mont Gerbier de Jonc South East France

Latite

Latite from Boxberg, High-Eifel, Germany

Photomicrograph of thin section of latite (in plane polarised light)

Photomicrograph of thin section of latite (in cross polarised light)

Latite is an igneous, volcanic (extrusive) rock, with aphanitic-aphyric to aphyric-porphyritic

texture. Its mineral assemblage is usually alkali feldspar and plagioclase (in a ratio < 1:4).

Quartz is absent in a feldspathoid-bearing latite, and olivine is absent in a quartz-bearing

latite. Biotite, hornblende, pyroxene and scarce olivine or quartz are common accessory

minerals.Rhomb porphyries are an unusual variety with gray-white porphyritic rhomb shaped

phenocrysts embedded in a very fine grained red-brown matrix. The composition of rhomb

porphyry places it in the trachyte - latite classification of the QAPF diagram.

Quartz latite

A quartz latite is a latite with a phenocryst modal composition containing 5-20% quartz. Above

20% quartz, the rock would be classified as a rhyolite.

Pegmatite

Pegmatite with blue corundum crystals

Pegmatite containing lepidolite, tourmaline, and quartz from the White Elephant Mine in the

Black Hills, South Dakota.

A pegmatite is a very coarse-grained, intrusive igneous rock composed of interlocking grains

usually larger than 2.5 cm in size;[1]

such rocks are referred to as pegmatitic.

Most pegmatites are composed of quartz, feldspar and mica; in essence a granite. Rarer

intermediate composition and mafic pegmatites containing amphibole, Ca-plagioclase feldspar,

pyroxene and other minerals are known, found in recrystallised zones and apophyses associated

with large layered intrusions.

Crystal size is the most striking feature of pegmatites, with crystals usually over 5 cm in size.

Individual crystals over 10 meters across have been found, and the world's largest crystal was

found within a pegmatite.[citation needed]

Similarly, crystal texture and form within pegmatitic rock may be taken to extreme size and

perfection. Feldspar within a pegmatite may display exaggerated and perfect twinning,

exsolution lamellae, and when affected by hydrous crystallization, macroscale graphic texture is

known, with feldspar and quartz intergrown. Perthite feldspar within a pegmatite often shows

gigantic perthitic texture visible to the naked eye.

Petrology

Crystal growth rates in pegmatite must be incredibly fast to allow gigantic crystals to grow

within the confines and pressures of the Earth's crust. For this reason, the consensus on

pegmatitic growth mechanisms involves a combination of the following processes;

Low rates of nucleation of crystals coupled with high diffusivity to force growth of a few

large crystals instead of many smaller crystals

High vapor and water pressure, to assist in the enhancement of conditions of diffusivity

High concentrations of fluxing elements such as boron and lithium which lower the

temperature of solidification within the magma or vapor

Low thermal gradients coupled with a high wall rock temperature, explaining the

preponderance for pegmatite to occur only within greenschist metamorphic terranes

Despite this consensus on likely chemical, thermal and compositional conditions required to

promote pegmatite growth there are three main theories behind pegmatite formation;

1. Metamorphic; pegmatite fluids are created by devolatilisation (dewatering) of

metamorphic rocks, particularly felsic gneiss, to liberate the right constituents and water,

at the right temperature

2. Magmatic; pegmatites tend to occur in the aureoles of granites in most cases, and are

usually granitic in character, often closely matching the compositions of nearby granites.

Pegmatites thus represent exsolved granitic material which crystallises in the country

rocks

3. Metasomatic; pegmatite, in a few cases, could be explained by the action of hot alteration

fluids upon a rock mass, with bulk chemical and textural change.

Metasomatism is currently not well favored as a mechanism for pegmatite formation and it is

likely that metamorphism and magmatism are both contributors toward the conditions necessary

for pegmatite genesis.

Mineralogy

Pegmatitic granite, Rock Creek Canyon, eastern Sierra Nevada, California. Note pink potassium

feldspars and cumulate-filled chamber.

The mineralogy of a pegmatite is in all cases dominated by some form of feldspar, often with

mica and usually with quartz, being altogether "granitic" in character. Beyond that, pegmatite

may include most minerals associated with granite and granite-associated hydrothermal systems,

granite-associated mineralisation styles, for example greisens, and somewhat with skarn

associated mineralisation.

It is however impossible to quantify the mineralogy of pegmatite in simple terms because of their

varied mineralogy and difficulty in estimating the modal abundance of mineral species which are

of only a trace amount. This is because of the difficulty in counting and sampling mineral grains

in a rock which may have crystals from centimeters to meters across.

Garnet, commonly almandine or spessartine, is a common mineral within pegmatites intruding

mafic and carbonate-bearing sequences. Pegmatites associated with granitic domes within the

Archaean Yilgarn Craton intruding ultramafic and mafic rocks contain red, orange and brown

almandine garnet.

Tantalum and niobium minerals (columbite, tantalite, niobite) are found in association with

spodumene, lepidolite, tourmaline, cassiterite in the massive Greenbushes Pegmatite in the

Yilgarn Craton of Western Australia, considered a typical metamorphic pegmatite unassociated

with granite.

Geochemistry

Pegmatite is difficult to sample representatively due to the large size of the constituent mineral

crystals. Often, bulk samples of some 50–60 kg of rock must be crushed to obtain a meaningful

and repeatable result. Hence, pegmatite is often characterised by sampling the individual

minerals which comprise the pegmatite, and comparisons are made according to mineral

chemistry.

Geochemically, pegmatites typically have major element compositions approximating "granite",

however, when found in association with granitic plutons it is likely that a pegmatite dike will

have a different trace element composition with greater enrichment in large-ion lithophile

(incompatible) elements, boron, beryllium, aluminium, potassium and lithium, uranium, thorium,

cesium, et cetera.

Occasionally, enrichment in the unusual trace elements will result in crystallisation of equally

unusual and rare minerals such as beryl, tourmaline, columbite, tantalite, zinnwaldite and so

forth. In most cases, there is no particular genetic significance to the presence of rare mineralogy

within a pegmatite, however it is possible to see some causative and genetic links between, say,

tourmaline-bearing granite dikes and tourmaline-bearing pegmatites within the area of influence

of a composite granite intrusion (Mount Isa Inlier, Queensland, Australia).

Sedimentary rock

Middle Triassic marginal marine sequence of siltstones (below) and limestones (above), Virgin

Formation, southwestern Utah, USA

Sedimentary rock is a type of rock that is formed by sedimentation of material at the Earth's

surface and within bodies of water. Sedimentation is the collective name for processes that cause

mineral and/or organic particles (detritus) to settle and accumulate or minerals to precipitate

from a solution. Particles that form a sedimentary rock by accumulating are called sediment.

Before being deposited, sediment was formed by weathering and erosion in a source area, and

then transported to the place of deposition by water, wind, mass movement or glaciers which are

called agents of denudation.

The sedimentary rock cover of the continents of the Earth's crust is extensive, but the total

contribution of sedimentary rocks is estimated to be only 5% of the total volume of the crust.

Sedimentary rocks are only a thin veneer over a crust consisting mainly of igneous and

metamorphic rocks.

Sedimentary rocks are deposited in layers as strata, forming a structure called bedding. The study

of sedimentary rocks and rock strata provides information about the subsurface that is useful for

civil engineering, for example in the construction of roads, houses, tunnels, canals or other

constructions. Sedimentary rocks are also important sources of natural resources like coal, fossil

fuels, drinking water or ores.

The study of the sequence of sedimentary rock strata is the main source for scientific knowledge

about the Earth's history, including palaeogeography, paleoclimatology and the history of life.

The scientific discipline that studies the properties and origin of sedimentary rocks is called

sedimentology. Sedimentology is both part of geology and physical geography and overlaps

partly with other disciplines in the Earth sciences, such as pedology, geomorphology,

geochemistry or structural geology.

Genetic classification schemes

Based on the processes responsible for their formation, sedimentary rocks can be subdivided into

four groups: clastic sedimentary rocks, biochemical (or biogenic) sedimentary rocks, chemical

sedimentary rocks and a fourth category for "other" sedimentary rocks formed by impacts,

volcanism, and other minor processes.

Clastic sedimentary rocks

Main article: Clastic rock

Claystone deposited in Glacial Lake Missoula, Montana, USA. Note very fine and flat bedding, common

for distal lacustrine deposition.

Clastic sedimentary rocks are composed of silicate minerals and rock fragments that were

transported by moving fluids (as bed load, suspended load, or by sediment gravity flows) and

were deposited when these fluids came to rest. Clastic rocks are composed largely of quartz,

feldspar, rock (lithic) fragments, clay minerals, and mica; numerous other minerals may be

present as accessories and may be important locally.

Clastic sediment, and thus clastic sedimentary rocks, are subdivided according to the dominant

particle size (diameter). Most geologists use the Udden-Wentworth grain size scale and divide

unconsolidated sediment into three fractions: gravel (>2 mm diameter), sand (1/16 to 2 mm

diameter), and mud (clay is <1/256 mm and silt is between 1/16 and 1/256 mm). The

classification of clastic sedimentary rocks parallels this scheme; conglomerates and breccias are

made mostly of gravel, sandstones are made mostly of sand, and mudrocks are made mostly of

mud. This tripartite subdivision is mirrored by the broad categories of rudites, arenites, and

lutites, respectively, in older literature.

Subdivision of these three broad categories is based on differences in clast shape (conglomerates

and breccias), composition (sandstones), grain size and/or texture (mudrocks).

[edit] Conglomerates and breccias

Conglomerates are dominantly composed of rounded gravel and breccias are composed of

dominantly angular gravel.

Sandstones

Sandstone classification schemes vary widely, but most geologists have adopted the Dott

scheme,[1]

which uses the relative abundance of quartz, feldspar, and lithic framework grains and

the abundance of muddy matrix between these larger grains.

Composition of framework grains

The relative abundance of sand-sized framework grains determines the first word in a sandstone

name. For naming purposes, the abundance of framework grains is normalized to quartz,

feldspar, and lithic fragments formed from other rocks. These are the three most abundant

components of sandstones; all other minerals are considered accessories and not used in the

naming of the rock, regardless of abundance.

Quartz sandstones have >90% quartz grains Feldspathic sandstones have <90% quartz grains and more feldspar grains than lithic

grains Lithic sandstones have <90% quartz grains and more lithic grains than feldspar grains

Abundance of muddy matrix between sand grains

When sand-sized particles are deposited, the space between the sand grains either remains

open or is filled with mud (silt and/or clay sized particle).

"Clean" sandstones with open pore space (that may later be filled with cement) are called arenites

Muddy sandstones with abundant (>10%) muddy matrix are called wackes.

Six sandstone names are possible using descriptors for grain composition (quartz-, feldspathic-,

and lithic-) and amount of matrix (wacke or arenite). For example, a quartz arenite would be

composed of mostly (>90%) quartz grains and have little/no clayey matrix between the grains, a

lithic wacke would have abundant lithic grains (<90% quartz, remainder would have more lithics

than feldspar) and abundant muddy matrix, etc.

Although the Dott classification scheme[1]

is widely used by sedimentologists, common names

like greywacke, arkose, and quartz sandstone are still widely used by nonspecialists and in

popular literature.

Mudrocks

Lower Antelope Canyon was carved out of the surrounding sandstone by both mechanical weathering

and chemical weathering. Wind, sand, and water from flash flooding are the primary weathering agents.

Mudrocks are sedimentary rocks composed of at least 50% silt- and clay-sized particles. These

relatively fine-grained particles are commonly transported as suspended particles by turbulent

flow in water or air, and deposited as the flow calms and the particles settle out of suspension.

Most authors presently use the term "mudrock" to refer to all rocks composed dominantly of

mud.[2][3][4][5]

Mudrocks can be divided into siltstones (composed dominantly of silt-sized

particles), mudstones (subequal mixture of silt- and clay-sized particles), and claystones

(composed mostly of clay-sized particles).[2][3]

Most authors use "shale" is a term for a fissile

mudrock (regardless of grain size), although some older literature uses the term "shale" as a

synonym for mudrock.

Biochemical sedimentary rocks

Outcrop of Ordovician oil shale (kukersite), northern Estonia

Biochemical sedimentary rocks are created when organisms use materials dissolved in air or

water to build their tissue. Examples include:

Most types of limestone are formed from the calcareous skeletons of organisms such as corals, mollusks, and foraminifera.

Coal which forms as plants remove carbon from the atmosphere and combine with other elements to build their tissue.

Deposits of chert formed from the accumulation of siliceous skeletons from microscopic organisms such as radiolaria and diatoms.

Chemical sedimentary rocks

Chemical sedimentary rock forms when mineral constituents in solution become supersaturated

and inorganically precipitate. Common chemical sedimentary rocks include oolitic limestone and

rocks composed of evaporite minerals such as halite (rock salt), sylvite, barite and gypsum.

"Other" sedimentary rocks

This fourth miscellaneous category includes rocks formed by Pyroclastic flows, impact breccias,

volcanic breccias, and other relatively uncommon processes.

Compositional classification schemes

Alternately, sedimentary rocks can be subdivided into compositional groups based on their

mineralogy:

Siliciclastic sedimentary rocks, as described above, are dominantly composed of silicate minerals. The sediment that makes up these rocks was transported as bed load, suspended load, or by sediment gravity flows. Siliciclastic sedimentary rocks are subdivided into conglomerates and breccias, sandstone, and mudrocks.

Carbonate sedimentary rocks are composed of calcite (rhombohedral CaCO3), aragonite (orthorhombic CaCO3), dolomite (CaMg(CO3)2), and other carbonate minerals based on the CO3

2- ion. Common examples include limestone and dolostone.

Evaporite sedimentary rocks are composed of minerals formed from the evaporation of water. The most common evaporite minerals are carbonates (calcite and others based on CO3

2-), chlorides (halite and others built on Cl-), and sulfates (gypsum and others built on SO4

2-). Evaporite rocks commonly include abundant halite (rock salt), gypsum, and anhydrite.

Organic-rich sedimentary rocks have significant amounts of organic material, generally in excess of 5% total organic carbon. Common examples include coal, oil shale, and other sedimentary rocks that act as reservoirs for liquid hydrocarbons and natural gas.

Siliceous sedimentary rocks are almost entirely composed of silica (SiO2), typically as chert, opal, chalcedony or other microcrystalline forms.

Iron-rich sedimentary rocks are composed of >15% iron; the most common forms are banded iron formations and ironstones[3]

Phosphatic sedimentary rocks are composed of phosphate minerals and contain more than 6.5% phosphorus; examples include deposits of phosphate nodules, bone beds, and phosphatic mudrocks[4]

Creation of sedimentary rocks

Sediment transport and deposition

Cross-bedding and scour in a fine sandstone; the Logan Formation (Mississippian) of Jackson County,

Ohio.

Sedimentary rocks are formed when sediment is deposited out of air, ice, wind, gravity, or water

flows carrying the particles in suspension. This sediment is often formed when weathering and

erosion break down a rock into loose material in a source area. The material is then transported

from the source area to the deposition area. The type of sediment transported depends on the

geology of the hinterland (the source area of the sediment). However, some sedimentary rocks,

like evaporites, are composed of material that formed at the place of deposition. The nature of a

sedimentary rock therefore not only depends on sediment supply, but also on the sedimentary

depositional environment in which it formed.

Diagenesis

Pressure solution at work in a clastic rock. While material dissolves at places where grains are in contact,

material crystallizes from the solution (as cement) in open pore spaces. This means there is a net flow of

material from areas under high stress to those under low stress. As a result, the rock becomes more

compact and harder. Loose sand can become sandstone in this way.

Main article: diagenesis

The term diagenesis is used to describe all the chemical, physical, and biological changes,

including cementation, undergone by a sediment after its initial deposition, exclusive of surface

weathering. Some of these processes cause the sediment to consolidate: a compact, solid

substance forms out of loose material. Young sedimentary rocks, especially those of Quaternary

age (the most recent period of the geologic time scale) are often still unconsolidated. As

sediment deposition builds up, the overburden (or lithostatic) pressure rises and a process known

as lithification takes place.

Sedimentary rocks are often saturated with seawater or groundwater, in which minerals can

dissolve or from which minerals can precipitate. Precipitating minerals reduce the pore space in a

rock, a process called cementation. Due to the decrease in pore space, the original connate fluids

are expelled. The precipitated minerals form a cement and make the rock more compact and

competent. In this way, loose clasts in a sedimentary rock can become "glued" together.

When sedimentation continues, an older rock layer becomes buried deeper as a result. The

lithostatic pressure in the rock increases due to the weight of the overlying sediment. This causes

compaction, a process in which grains mechanical reorganize. Compaction is, for example, an

important diagenetic process in clay, which can initially consist of 60% water. During

compaction, this interstitial water is pressed out of pore spaces. Compacation can also be the

result of dissolution of grains by pressure solution. The dissolved material precipitates again in

open pore spaces, which means there is a nett flow of material into the pores. However, in some

cases a certain mineral dissolves and not precipitate again. This process is called leaching and

increases pore space in the rock.

Some biochemical processes, like the activity of bacteria, can affect minerals in a rock and are

therefore seen as part of diagenesis. Fungi and plants (by their roots) and various other organisms

that live beneath the surface can also influence diagenesis.

Burial of rocks due to ongoing sedimentation leads to increased pressure and temperature, which

stimulates certain chemical reactions. An example is the reactions by which organic material

becomes lignite or coal. When temperature and pressure increase still further, the realm of

diagenesis makes way for metamorphism, the process that forms metamorphic rock.

Properties

A piece of a banded iron formation, a type of rock that consists of alternating layers with iron(III) oxide

(red) and iron(II) oxide (grey). BIFs were mostly formed during the Precambrian, when the atmosphere

wasn't yet rich in oxygen. Moories Group, Barberton Greenstone Belt, South Africa.

Color

The color of a sedimentary rock is often mostly determined by iron, an element with two major

oxides: iron(II) oxide and iron(III) oxide. Iron(II) oxide only forms under anoxic circumstances

and gives the rock a grey or greenish colour. Iron(III) oxide is often in the form of the mineral

hematite and gives the rock a reddish to brownish colour. In arid continental climates rocks are

in direct contact with the atmosphere, and oxidation is an important process, giving the rock a

red or orange colour. Thick sequences of red sedimentary rocks formed in arid climates are

called red beds. However, a red colour does not necessarily mean the rock formed in a

continental environment or arid climate.[6]

The presence of organic material can colour a rock black or grey. Organic material is in nature

formed from dead organisms, mostly plants. Normally, such material eventually decays by

oxidation or bacterial activity. Under anoxic circumstances, however, organic material cannot

decay and becomes a dark sediment, rich in organic material. This, can for example, occur at the

bottom of deep seas and lakes. There is little water current in such environments, so oxygen from

surface water is not brought down, and the deposited sediment is normally a fine dark clay. Dark

rocks rich in organic material are therefore often shales.[7]

Texture

Diagram showing the difference between well-sorted (left) and poorly sorted (right) clastic rocks.

The size, form and orientation of clasts or minerals in a rock is called its texture. The texture is a

small-scale property of a rock, but determined many of its large-scale properties, such as the

density, porosity or permeabililty.[8]

Clastic rocks have a 'clastic texture', which means they consist of clasts. The 3D orientation of

these clasts is called the fabric of the rock. Between the clasts the rock can be composed of a

matrix or a cement (the latter can consist of crystals of one or more precipitated minerals). The

size and form of clasts can be used to determine the velocity and direction of current in the

sedimentary environment where the rock was formed; fine, calcareous mud only settles in quiet

water, while gravel and larger clasts are only deposited by rapidly moving water.[9]

The grain

size of a rock is usually expressed with the Wentworth scale, though alternative scales are used

sometimes. The grain size can be expressed as a diameter or a volume, and is always an average

value - a rock is composed of clasts with different sizes. The statistical distribution of grain sizes

is different for different rock types and is described in a property called the sorting of the rock.

When all clasts are more or less of the same size, the rock is called 'well-sorted', when there is a

large spread in grain size, the rock is called 'poorly sorted'.[10]

Diagram showing the influence of rounding and sphericity.

The form of clasts can reflect the origin of the rock.

Coquina, a rock composed of clasts of broken shells, can only form in energetic water. The form

of a clast can be described by using four parameters:[11]

Surface texture describes the amount of small-scale relief of the surface of a grain that is too small to influence the general shape.

rounding describes the general smoothness of the shape of a grain. 'Sphericity' describes the degree to which the grain approaches a sphere.

'Grain form' describes the three dimensional shape of the grain.

Chemical sedimentary rocks have a non-clastic texture, consisting entirely of crystals. To

describe such a texture only the average size of the crystals and the fabric are necessary.

Mineralogy

Most sedimentary rocks contain either quartz (especially siliciclastic rocks) or calcite (especially

carbonate rocks). In contrast with igneous and metamorphic rocks, a sedimentary rocks usually

contains very few different major minerals. However, the origin of the minerals in a sedimentary

rock is often more complex than those in an igneous rock. Minerals in a sedimentary rock can

have formed by precipitation during sedimentation or diagenesis. In the second case, the mineral

precipitate can have grown over an older generation of cement.[12]

A complex diagenetic history

can be studied by optical mineralogy, using a petrographic microscope.

Carbonate rocks dominantly consist of carbonate minerals like calcite, aragonite or dolomite.

Both cement and clasts (including fossils and ooids) of a carbonate rock can consist of carbonate

minerals. The mineralogy of a clastic rock is determined by the supplied material from the source

area, the manner of transport to the place of deposition and the stability of a particular mineral.

The stability of the major rock forming minerals (their resistance to weathering) is expressed by

Bowen's reaction series. In this series, quartz is most stable, followed by feldspar, micas, and

other less stable minerals that are only present when little weathering has occurred.[13]

The

amount of weathering depends mainly on the distance to the source area, the local climate and

the time it took for the sediment to be transported there. In most sedimentary rocks, mica,

feldspar and less stable minerals have reacted to clay minerals like kaolinite, illite or smectite.

Fossils

Fossil-rich layers in a sedimentary rock, Año Nuevo State Reserve, California.

Main articles: fossil and fossilisation

Sedimentary rocks are the only type of rock that can contain fossils, the remains or imprints of

dead organisms. In nature, dead organisms are usually quickly removed by scavengers, bacteria,

rotting and erosion. In some exceptional circumstances a carcass is fossilized because these

natural processes are unable to work. The chance of fossilisation is higher when the

sedimentation rate is high (so that a carcass is quickly buried), in anoxic environments (where

little bacterial activity exists) or when the organism had a particularly hard skeleton. Larger,

well-preserved fossils are relatively rare. Most sedimentary rocks contains fossils, though with

many the fact only becomes apparent when studied under a microscope (microfossils) or with a

loupe.

Burrows in a turbidite, made by crustaceans. San Vincente Formation (early Eocene) of the Ainsa Basin,

southern foreland of the Pyrenees.

Fossils can both be the direct remains or imprints of organisms and their skeletons. Most

commonly preserved are the harder parts of organisms such as bones, shells, woody tissue of

plants. Soft tissue has a much smaller chance of being preserved and fossilized and soft tissue of

animals older than 40 million years is very rare.[14]

Imprints of organisms made while still alive

are called trace fossils. Examples are burrows, foot prints, etc.

Being part of a sedimentary rock, fossils undergo the same diagenetic processes as the rock. A

shell consisting of calcite can for example dissolve, while a cement of silica then fills the cavity.

In the same way, precipitating minerals can fill cavities formerly occupied by blood vessels,

vascular tissue or other soft tissues. This preserves the form of the organism but changes the

chemical composition, a process called permineralisation.[15]

The most common minerals in

permineralisation cements are carbonates (especially calcite), forms of amorphous silica

(chalcedony, flint, chert) and pyrite. In the case of silica cements, the process is called

lithification.

At high pressure and temperature, the organic material of a dead organism undergoes chemical

reactions in which volatiles like water and carbon dioxide are expulsed. The fossil, in the end,

consists of a thin layer of pure carbon or its mineralized form, graphite. This form of fossilisation

is called carbonisation. It is particularly important for plant fossils.[16]

The same process is

responsible for the formation of fossil fuels like lignite or coal.

Primary sedimentary structures

Cross-bedding in a fluviatile sandstone, Middle Old Red Sandstone (Devonian) on Bressay, Shetland

Islands.

A flute cast, a type of sole marking, from the Book Cliffs of Utah.

Ripple marks formed by a current in a sandstone that was later tilted. Location: Haßberge, Bavaria.

Structures in sedimentary rocks can be divided in 'primary' structures (formed during deposition)

and 'secondary' structures (formed after deposition). Unlike textures, structures are always large-

scale features that can easily be studied in the field. Sedimentary structures can tell something

about the sedimentary environment or can serve to tell which side originally faced up where

tectonics have tilted or overturned sedimentary layers.

Sedimentary rocks are laid down in layers called beds or strata. A bed is defined as a layer of

rock that has a uniform lithology and texture. Beds form by the deposition of layers of sediment

on top of each other. The sequence of beds that characterizes sedimentary rocks is called

bedding.[17]

Single beds can be a couple of centimetres to several meters thick. Finer, less

pronounced layers are called laminae and the structure it forms in a rock is called lamination.

Laminae are usually less than a few centimetres thick.[18]

Though bedding and lamination are

often originally horizontal in nature, this is not always the case. In some environments, beds are

deposited at a (usually small) angle. Sometimes multiple sets of layers with different orientations

exist in the same rock, a structure called cross-bedding.[19]

Cross-bedding forms when small-

scale erosion occurs during deposition, cutting off part of the beds. Newer beds then form at an

angle to older ones.

The opposite of cross-bedding is parallel lamination, where all sedimentary layering is

parallel.[20]

With laminations, differences are generally caused by cyclic changes in the sediment

supply, caused for example by seasonal changes in rainfall, temperature or biochemical activity.

Laminae that represent seasonal changes (like tree rings) are called varves. Some rocks have no

lamination at all, their structural character is called massive bedding.

Graded bedding is a structure where beds with a smaller grain size occur on top of beds with

larger grains. This structure forms when fast flowing water stops flowing. Larger, heavier clasts

in suspension settle first, then smaller clasts. Though graded bedding can form in many different

environments, it is characteristic for turbidity currents.[21]

The bedform (the surface of a particular bed) can be indicative for a particular sedimentary

environment too. Examples of bed forms include dunes and ripple marks. Sole markings, such as

tool marks and flute casts, are groves dug into a sedimentary layer that are preserved. These are

often elongated structures and can be used to establish the direction of the flow during

deposition.[22]

Ripple marks also form in flowing water. There are two types: asymmetric wave ripples and

symmetric current ripples. Environments where the current is in one direction, such as rivers,

produce asymmetric ripples. The longer flank of such ripples is oriented opposite to the direction

of the current.[23]

Wave ripples occur in environments where currents occur in all directions, such

as tidal flats.

Another type of bed form are mud cracks, caused by the dehydration of sediment that

occasionally comes above the water surface. Such structures are commonly found at tidal flats or

point bars along rivers.

Secondary sedimentary structures

Secondary sedimentary structures are structures in sedimentary rocks which formed after

deposition. Such structures form by chemical, physical and biological processes inside the

sediment. They can be indicators for circumstances after deposition. Some can be used as way up

criteria.

Organic presence in a sediment can leave more traces than just fossils. Preserved tracks and

burrows are examples of trace fossils (also called ichnofossils).[24]

Some trace fossils such as

paw prints of dinosaurs or early humans can capture human imagination, but such traces are

relatively rare. Most trace fossils are burrows of molluscs or arthropods. This burrowing is called

bioturbation by sedimentologists. It can be a valuable indicator of the biological and ecological

environment after the sediment was deposited. On the other hand, the burrowing activity of

organisms can destroy other (primary) structures in the sediment, making a reconstruction more

difficult.

Chert concretions in chalk, Middle Lefkara Formation (upper Paleocene to middle Eocene), Cyprus.

Secondary structures can also have been formed by diagenesis or the formation of a soil

(pedogenesis) when a sediment is exposed above the water level. An example of a diagenetic

structure common in carbonate rocks is a stylolite.[25]

Stylolites are irregular planes were material

was dissolved into the pore fluids in the rock. The result of precipitation of a certain chemical

species can be colouring and staining of the rock, or the formation of concretions. Concretions

are roughly concentric bodies with a different composition from the host rock. Their formation

can be the result of localized precipitation due to small differences in composition or porosity of

the host rock, such as around fossils, inside burrows or around plant roots.[26]

In carbonate rocks

such as limestone or chalk, chert or flint concretions are common, while terrestrial sandstones

can have iron concretions. Calcite concretions in clay are called septarian concretions.

After deposition, physical processes can deform the sediment, forming a third class of secondary

structures. Density contrasts between different sedimentary layers, such as between sand and

clay, can result in flame structures or load casts, formed by inverted diapirism.[27]

The diapirism

causes the denser upper layer to sink into the other layer. Sometimes, density contrast can result

or grow when one of the lithologies dehydrates. Clay can be easily compressed as a result of

dehydration, while sand retains the same volume and becomes relatively less dense. On the other

hand, when the pore fluid pressure in a sand layer surpasses a critical point the sand can flow

through overlying clay layers, forming discordant bodies of sedimentary rock called sedimentary

dykes (the same process can form mud volcanoes on the surface).

A sedimentary dyke can also be formed in a cold climate where the soil is permanently frozen

during a large part of the year. Frost weathering can form cracks in the soil that fill with rubble

from above. Such structures can be used as climate indicators as well as way up structures.[28]

Density contrasts can also cause small-scale faulting, even while sedimentation goes on (syn-

sedimentary faulting).[29]

Such faulting can also occur when large masses of non-lithified

sediment are deposited on a slope, such as at the front side of a delta or the continental slope.

Instabilities in such sediments can result in slumping. The resulting structures in the rock are

syn-sedimentary folds and faults, which can be difficult to distinguish from folds and faults

formed by tectonic forces in lithified rocks.

Sedimentary environments

The setting in which a sedimentary rock forms is called the sedimentary environment. Every

environment has a characteristic combination of geologic processes and circumstances. The type

of sediment that is deposited is not only dependent on the sediment that is transported to a place,

but also on the environment itself.[30]

A marine environment means the rock was formed in a sea or ocean. Often, a distinction is made

between deep and shallow marine environments. Deep marine usually refers to environments

more than 200 m below the water surface. Shallow marine environments exist adjacent to

coastlines and can extend out to the boundaries of the continental shelf. The water in such

environments has a generally higher energy than that in deep environments, because of wave

activity. This means coarser sediment particles can be transported and the deposited sediment

can be coarser than in deep environments. When the available sediment is transported from the

continent, an alternation of sand, clay and silt is deposited. When the continent is far away, the

amount of such sediment brought in may be small, and biochemical processes dominate the type

of rock that forms. Especially in warm climates, shallow marine environments far offshore

mainly see deposition of carbonate rocks. The shallow, warm water is an ideal habitat for many

small organisms that build carbonate skeletons. When these organisms die their skeletons sink to

the bottom, forming a thick layer of calcareous mud that may lithify into limestone. Warm

shallow marine environments also are ideal environments for coral reefs, where the sediment

consists mainly of the calcareous skeletons of larger organisms.[31]

In deep marine environments, the water current over the sea bottom is small. Only fine particles

can be transported to such places. Typically sediments depositing on the ocean floor are fine clay

or small skeletons of micro-organisms. At 4 km depth, the solubility of carbonates increases

dramatically (the depth zone where this happens is called the lysocline). Calcareous sediment

that sinks below the lysocline dissolve, so no limestone can be formed below this depth.

Skeletons of micro-organisms formed of silica (such as radiolarians) still deposit though. An

example of a rock formed out of silica skeletons is radiolarite. When the bottom of the sea has a

small inclination, for example at the continental slopes, the sedimentary cover can become

unstable, causing turbidity currents. Turbidity currents are sudden disturbances of the normally

quite deep marine environment and can cause the geologically speaking instantaneous deposition

of large amounts of sediment, such as sand and silt. The rock sequence formed by a turbidity

current is called a turbidite.[32]

The coast is an environment dominated by wave action. At the beach, dominantly coarse

sediment like sand or gravel is deposited, often mingled with shell fragments. Tidal flats and

shoals are places that sometimes dry out because of the tide. They are often cross-cut by gullies,

where the current is strong and the grain size of the deposited sediment is larger. Where along a

coast (either the coast of a sea or a lake) rivers enter the body of water, deltas can form. These

are large accumulations of sediment transported from the continent to places in front of the

mouth of the river. Deltas are dominantly composed of clastic sediment.

A sedimentary rock formed on the land has a continental sedimentary environment. Examples of

continental environments are lagoons, lakes, swamps, floodplains and alluvial fans. In the quiet

water of swamps, lakes and lagoons, fine sediment is deposited, mingled with organic material

from dead plants and animals. In rivers, the energy of the water is much higher and the

transported material consists of clastic sediment. Besides transport by water, sediment can in

continental environments also be transported by wind or glaciers. Sediment transported by wind

is called aeolian and is always very well sorted, while sediment transported by a glacier is called

glacial and is characterized by very poor sorting.[33]

Sedimentary facies

Sedimentary environments usually exist alongside each other in certain natural successions. A

beach, where sand and gravel is deposited, is usually bounded by a deeper marine environment a

little offshore, where finer sediments are deposited at the same time. Behind the beach, there can

be dunes (where the dominant deposition is well sorted sand) or a lagoon (where fine clay and

organic material is deposited). Every sedimentary environment has its own characteristic

deposits. The typical rock formed in a certain environment is called its sedimentary facies. When

sedimentary strata accumulate through time, the environment can shift, forming a change in

facies in the subsurface at one location. On the other hand, when a rock layer with a certain age

is followed laterally, the lithology (the type of rock) and facies eventually change.[34]

Shifting sedimentary facies in the case of transgression (above) and regression of the sea (below).

Facies can be distinguished in a number of ways: the most common ways are by the lithology

(for example: limestone, siltstone or sandstone) or by fossil content. Coral for example only lives

in warm and shallow marine environments and fossils of coral are thus typical for shallow

marine facies. Facies determined by lithology are called lithofacies; facies determined by fossils

are biofacies.[35]

Sedimentary environments can shift their geographical positions through time. Coastlines can

shift in the direction of the sea when the sea level drops, when the surface rises due to tectonic

forces in the Earth's crust or when a river forms a large delta. In the subsurface, such geographic

shifts of sedimentary environments of the past are recorded in shifts in sedimentary facies. This

means that sedimentary facies can change either parallel or perpendicular to an imaginary layer

of rock with a fixed age, a phenomenon described by Walther's facies rule.[36]

The situation in which coastlines move in the direction of the continent is called transgression. In

the case of transgression, deeper marine facies are deposited over shallower facies, a succession

called onlap. Regression is the situation in which a coastline moves in the direction of the sea.

With regression, shallower facies are deposited on top of deeper facies, a situation called

offlap.[37]

The facies of all rocks of a certain age can be plotted on a map to give an overview of the

palaeogeography. A sequence of maps for different ages can give an insight in the development

of the regional geography.

Sedimentary basins

Main article: sedimentary basin

Places where large-scale sedimentation takes place are called sedimentary basins. The amount of

sediment that can be deposited in a basin depends on the depth of the basin, the so called

accommodation space. Depth, shape and size of a basin depend on tectonics, movements within

the Earth's lithosphere. Where the lithosphere moves upward (tectonic uplift), land eventually

rises above sea level, so that and erosion removes material, and the area becomes a source for

new sediment. Where the lithosphere moves downward (tectonic subsidence), a basin forms and

sedimentation can take place. When the lithosphere keeps subsiding, new accommodation space

keeps being created.

A type of basin formed by the moving apart of two pieces of a continent is called a rift basin.

Rift basins are elongated, narrow and deep basins. Due to divergent movement, the lithosphere is

stretched and thinned, so that the hot asthenosphere rises and heats the overlying rift basin. Apart

from continental sediments, rift basins normally also have part of their infill consisting of

volcanic deposits. When the basin grows due to continued stretching of the lithosphere, the rift

grows and the sea can enter, forming marine deposits.

When a piece of lithosphere that was heated and stretched cools again, its density rises, causing

isostatic subsidence. If this subsidence continues long enough the basin is called a sag basin.

Examples of sag basins are the regions along passive continental margins, but sag basins can also

be found in the interior of continents. In sag basins, the extra weight of the newly deposited

sediments is enough to keep the subsidence going in a vicious circle. The total thickness of the

sedimentary infill in a sag basins can thus exceed 10 km.

A third type of basin exists along convergent plate boundaries - places where one tectonic plate

moves under another into the asthenosphere. The subducting plate bends and forms a fore-arc

basin in front of the overriding plate—an elongated, deep asymmetric basin. Fore-arc basins are

filled with deep marine deposits and thick sequences of turbidites. Such infill is called flysch.

When the convergent movement of the two plates results in continental collision, the basin

becomes shallower and develops into a foreland basin. At the same time, tectonic uplift forms a

mountain belt in the overriding plate, from which large amounts of material are eroded and

transported to the basin. Such erosional material of a growing mountain chain is called molasse

and has either a shallow marine or a continental facies.

At the same time, the growing weight of the mountain belt can cause isostatic subsidence in the

area of the overriding plate on the other side to the mountain belt. The basin type resulting from

this subsidence is called a back-arc basin and is usually filled by shallow marine deposits and

molasse.[38]

Cyclic alternation of competent and less competent beds in the Blue Lias at Lyme Regis, southern

England.

Influence of astronomical cycles

In many cases facies changes and other lithological features in sequences of sedimentary rock

have a cyclic nature. This cyclic nature was caused by cyclic changes in sediment supply and the

sedimentary environment. Most of these cyclic changes are caused by astronomic cycles. Short

astronomic cycles can be the difference between the tides or the spring tide every two weeks. On

a larger time-scale, cyclic changes in climate and sea level are caused by Milankovitch cycles:

cyclic changes in the orientation and/or position of the Earth's rotational axis and orbit around

the Sun. There are a number of Milankovitch cycles known, lasting between 10,000 and 200,000

years.[39]

Relatively small changes in the orientation of the Earth's axis or length of the seasons can be a

major influence on the Earth's climate. An example are the ice ages of the past 2.6 million years

(the Quaternary period), which are assumed to have been caused by astronomic cycles.[40]

Climate change can influence the global sea level (and thus the amount of accommodation space

in sedimentary basins) and sediment supply from a certain region. Eventually, small changes in

astronomic parameters can cause large changes in sedimentary environment and sedimentation.

Sedimentation rates

The rate at which sediment is deposited differs depending on the location. A channel in a tidal

flat can see the deposition of a few metres of sediment in one day, while on the deep ocean floor

each year only a few millimetres of sediment accumulate. A distinction can be made between

normal sedimentation and sedimentation caused by catastrophic processes. The latter category

includes all kinds of sudden exceptional processes like mass movements, rock slides or flooding.

Catastrophic processes can see the sudden deposition of a large amount of sediment at once. In

some sedimentary environments, most of the total column of sedimentary rock was formed by

catastrophic processes, even though the environment is usually a quiet place. Other sedimentary

environments are dominated by normal, ongoing sedimentation.[41]

In some sedimentary environments, sedimentation only occurs in some places. In a desert, for

example, the wind deposits siliciclastic material (sand or silt) in some spots, or catastrophic

flooding of a wadi may cause sudden deposis of large quantities of detrital material, but in most

places eolian erosion dominates. The amount of sedimentary rock that forms is not only

dependent on the amount of supplied material, but also on how well the material consolidates.

Erosion removes most deposited sediment shortly after deposition.[41]

Stratigraphy

The Permian through Jurassic stratigraphy of the Colorado Plateau area of southeastern Utah that

makes up much of the famous prominent rock formations in protected areas such as Capitol Reef

National Park and Canyonlands National Park. From top to bottom: Rounded tan domes of the Navajo

Sandstone, layered red Kayenta Formation, cliff-forming, vertically jointed, red Wingate Sandstone,

slope-forming, purplish Chinle Formation, layered, lighter-red Moenkopi Formation, and white, layered

Cutler Formation sandstone. Picture from Glen Canyon National Recreation Area, Utah.

That new rock layers are above older rock layers is stated in the principle of superposition. There

are usually some gaps in the sequence called unconformities. These represent periods where no

new sediments were laid down, or when earlier sedimentary layers raised above sea level and

eroded away.

Sedimentary rocks contain important information about the history of the Earth. They contain

fossils, the preserved remains of ancient plants and animals. Coal is considered a type of

sedimentary rock. The composition of sediments provides us with clues as to the original rock.

Differences between successive layers indicate changes to the environment over time.

Sedimentary rocks can contain fossils because, unlike most igneous and metamorphic rocks, they

form at temperatures and pressures that do not destroy fossil remains.

Conglomerate (geology)

conglomerate

— Sedimentary Rock —

Boulder of conglomerate with cobble-sized clasts. Rock

hammer for scale.

Carmelo Formation (Conglomerate) at Point Lobos

Rock climbing hold on a Conglomerate rock in Margalef, Spain.

A conglomerate (pronounced /kəŋˈɡlɒmərɨt/) is a rock consisting of individual clasts within a

finer-grained matrix that have become cemented together. Conglomerates are sedimentary rocks

consisting of rounded fragments and are thus differentiated from breccias, which consist of

angular clasts.[1]

Both conglomerates and breccias are characterized by clasts larger than sand

(>2 mm).

Classification

In addition to the factors described in this section, conglomerates are classified in terms of both

their rounding and sorting.

Texture

Paraconglomerates consist of a matrix-supported rock that contains at least 15% sand-sized or

smaller grains (<2 mm), the rest being larger grains of varying sizes.[2]

Orthoconglomerates consist of a clast-supported rock with less than 15% matrix of sand and

finer particles.[3]

Metamorphic alteration transforms conglomerate into metaconglomerate.

A conglomerate at the base of the Cambrian in the Black Hills, South Dakota.

Section of polymict conglomerate from offshore rock core, Alaska, approximate depth 10,000 ft.

Clast composition

Conglomerates are classified for the lithologies of the clasts[4]

Monomict - clasts with only a single lithology

Oligomict - clasts of only a few different lithologies

Polymict - clasts of many different lithologies

Intraformational - clasts derived from the same formation in which they are found

Extraformational - clasts derived older rocks than the formation in which they are found

Clast size

Conglomerates are also classified by the dominant clast size.

Granule conglomerate 2–4 mm

Pebble conglomerate 4–64 mm

Cobble conglomerate 64–256 mm

Boulder conglomerate >256 mm

Sedimentary environments

Conglomerates are deposited in a variety of sedimentary environments.

Deepwater marine

In turbidites, the basal part of a bed is typically coarse-grained and sometimes conglomeratic. In

this setting, conglomerates are normally very well sorted, well-rounded and often with a strong

A-axis type imbrication of the clasts.[5]

Shallow marine

Conglomerates are normally present at the base of sequences laid down during marine

trangressions above an unconformity, and are known as basal conglomerates. They represent the

position of the shoreline at a particular time and will be diachronous.[6]

Fluvial

Conglomerates deposited in fluvial environments are typically well-rounded and well-sorted.

Clasts of this size are carried as bedload and only at times of high flow-rate. The maximum clast

size decreases as the clasts are transported further due to attrition, so conglomerates are more

characteristic of immature river systems. In the sediments deposited by mature rivers,

conglomerates are generally confined to the basal part of a channel fill where they are known as

pebble lags.[7]

Conglomerates deposited in a fluvial environment often have an AB-plane type

imbrication.

Alluvial

Alluvial deposits are formed in areas of high relief and are typically coarse-grained. At mountain

fronts individual alluvial fans merge together to form braidplains and these two environments are

associated with the thickest deposits of conglomerates. The bulk of conglomerates deposited in

this setting are clast-supported with a strong AB-plane imbrication. Some matrix-supported

conglomerates are present, a result of debris-flow deposition on some alluvial fans.[5]

Glacial

Glaciers carry a lot of coarse-grained material and many glacial deposits are conglomeratic.

Tillites, the sediments deposited directly by a glacier, are typically poorly-sorted, matrix-

supported conglomerates. The matrix is generally fine-grained, consisting of finely milled rock

fragments. Waterlain deposits associated with glaciers are often conglomeratic, forming

structures such as eskers.[7]

Examples

A spectacular example of conglomerate can be seen at Montserrat, near Barcelona. Here erosion

has created vertical channels giving the characteristic jagged shapes for which the mountain is

named (Montserrat literally means "jagged mountain"). The rock is strong enough to be used as a

building material - see Montserrat abbey front at full resolution for detail of the rock structure.

Another spectacular example of conglomerate, the Crestone Conglomerate may be viewed in and

near the town of Crestone, at the foot of the Sangre de Cristo Range in Colorado's San Luis

Valley. The Crestone Conglomerate is a metamorphic rock stratum and consists of tiny to quite

large rocks that appear to have been tumbled in an ancient river. Some of the rocks have hues of

red and green.

Conglomerate may also be seen in the domed hills of Kata Tjuta, in Australia's Northern

Territory.

In the nineteenth century a thick layer of Pottsville conglomerate was recognized to underlie

anthracite coal measures in Pennsylvania.[8]

Fanglomerate

Fanglomerate

When a series of conglomerates accumulates into an alluvial fan, in rapidly eroding (e.g. desert)

environments, the resulting rock unit is often called a fanglomerate. These form the basis of a

number of large oil fields, e.g. the Tiffany and Brae fields in the North Sea. These fanglomerates

were actually deposited into a deep marine environment but against a rapidly moving fault line,

which supplied an intermittent stream of debris into the conglomerate pile. The sediment fans are

several kilometers deep at the fault line and the sedimentation moved focus repeatedly, as

different sectors of the fault moved.

Sandstone

Sandstone

— Sedimentary Rock —

Prepared sample of sandstone

Composition

Typically quartz and/or feldspar (on earth); lithic fragments

are also common. Other minerals may be found in

particularly immature sandstone.

Alcove in the Navajo Sandstone

This article is about the geological rock type. For other uses, see Sandstone (disambiguation).

Sandstone (sometimes known as arenite) is a sedimentary rock composed mainly of sand-sized

minerals or rock grains. Most sandstone is composed of quartz and/or feldspar because these are

the most common minerals in the Earth's crust. Like sand, sandstone may be any color, but the

most common colors are tan, brown, yellow, red, gray and pink, white. Since sandstone beds

often form highly visible cliffs and other topographic features, certain colors of sandstone have

been strongly identified with certain regions.

Some sandstones are resistant to weathering, yet are easy to work. This makes sandstone a

common building and paving material. However, some that have been used in the past, such as

the Collyhurst sandstone used in North West England, have been found less resistant,

necessitating repair and replacement in older buildings.[1]

Because of the hardness of the

individual grains, uniformity of grain size and friability of their structure, some types of

sandstone are excellent materials from which to make grindstones, for sharpening blades and

other implements. Non-friable sandstone can be used to make grindstones for grinding grain,

e.g., gritstone.

Rock formations that are primarily composed of sandstone usually allow percolation of water

and other fluids and are porous enough to store large quantities, making them valuable aquifers

and petroleum reservoirs. Fine-grained aquifers, such as sandstones, are more apt to filter out

pollutants from the surface than are rocks with cracks and crevices, such as limestone or other

rocks fractured by seismic activity.

Origins

Sand from Coral Pink Sand Dunes State Park, Utah. These are grains of quartz with a hematite

coating providing the orange color. Scale bar is 1.0 mm.

Millet-Seed sandstone macro (size: ~4 cm or ~1.6 in).

Sandstones are clastic in origin (as opposed to either organic, like chalk and coal, or chemical,

like gypsum and jasper).[2]

They are formed from cemented grains that may either be fragments

of a pre-existing rock or be mono-minerallic crystals. The cements binding these grains together

are typically calcite, clays and silica. Grain sizes in sands are defined (in geology) within the

range of 0.0625 mm to 2 mm (0.002-0.079 inches). Clays and sediments with smaller grain sizes

not visible with the naked eye, including siltstones and shales, are typically called argillaceous

sediments; rocks with larger grain sizes, including breccias and conglomerates are termed

rudaceous sediments.

Red sandstone interior of Lower Antelope Canyon, Arizona, worn smooth by erosion from flash

flooding over thousands of years.

The formation of sandstone involves two principal stages. First, a layer or layers of sand

accumulates as the result of sedimentation, either from water (as in a river, lake, or sea) or from

air (as in a desert). Typically, sedimentation occurs by the sand settling out from suspension; i.e.,

ceasing to be rolled or bounced along the bottom of a body of water (e.g., seas or rivers) or

ground surface (e.g., in a desert or erg). Finally, once it has accumulated, the sand becomes

sandstone when it is compacted by pressure of overlying deposits and cemented by the

precipitation of minerals within the pore spaces between sand grains.

The most common cementing materials are silica and calcium carbonate, which are often derived

either from dissolution or from alteration of the sand after it was buried. Colors will usually be

tan or yellow (from a blend of the clear quartz with the dark amber feldspar content of the sand).

A predominant additional colorant in the southwestern United States is iron oxide, which imparts

reddish tints ranging from pink to dark red (terracotta), with additional manganese imparting a

purplish hue. Red sandstones are also seen in the Southwest and West of England and Wales, as

well as central Europe and Mongolia. The regularity of the latter favors use as a source for

masonry, either as a primary building material or as a facing stone, over other construction.

The environment where it is deposited is crucial in determining the characteristics of the

resulting sandstone, which, in finer detail, include its grain size, sorting and composition and, in

more general detail, include the rock geometry and sedimentary structures. Principal

environments of deposition may be split between terrestrial and marine, as illustrated by the

following broad groupings:

Terrestrial environments

Sandstone near Stadtroda, Germany.

1. Rivers (levees, point bars, channel sands)

2. Alluvial fans

3. Glacial outwash

4. Lakes

5. Deserts (sand dunes and ergs)

Marine environments

1. Deltas

2. Beach and shoreface sands

3. Tidal flats

4. Offshore bars and sand waves

5. Storm deposits (tempestites)

6. Turbidites (submarine channels and fans)

Types

Sandstone composed mainly of quartz grains

Photomicrograph of a volcanic sand grain; upper picture is plane-polarized light, bottom picture

is cross-polarized light, scale box at left-center is 0.25 millimeter. This type of grain would be a

main component of a lithic sandstone.

Arkose

Arkosic sand in the Llano Uplift, Texas, USA with granite outcrops.

Arkose (pronounced /ˈɑrkoʊz/) is a detrital sedimentary rock, specifically a type of sandstone

containing at least 25% feldspar.[1]

,[2]

Arkosic sand is sand that is similarly rich in feldspar, and

thus the potential precursor of arkose. The other mineral components may vary, but quartz is

commonly dominant, and some mica is often present. Apart from the mineral content, rock

fragments may also be a significant component. Arkose usually contains small amounts of calcite

cement, which causes it to 'fizz' slightly in dilute hydrochloric acid; sometimes the cement also

contains iron oxide. Arkose is typically grey to reddish in colour. The sand grains making up an

arkose may range from fine to very coarse, but tends toward the coarser end of the scale. Fossils

are rare in arkose, due to the depositional processes that form it, although bedding is frequently

visible.

Arkose sandstone found in Slovakia

Arkose is generally formed from the weathering of feldspar-rich igneous or metamorphic, most

commonly granitic rocks, which are primarily composed of quartz and feldspar. These sediments

must be deposited rapidly and/or in a cold or arid environment such that the feldspar does not

undergo significant chemical weathering and decomposition; therefore arkose is designated a

texturally immature sedimentary rock. Arkose is often associated with conglomerate deposits

sourced from granitic terrain and is often found above unconformities over such granitic terrain.

The famous central Australian monolith Uluru (Ayers Rock) is composed of late

Neoproterozoic/Cambrian arkose, deposited in the Amadeus Basin.[3]

Sandstones fall into several major groups based on their mineralogy and texture. Below is a

partial list of common sandstone types.

quartz arenites are made up almost entirely of quartz grains, usually well sorted and

rounded. These pure quartz sands result from extensive weathering that occurred before

and during transport and removed everything but quartz, the most stable mineral. They

are common in beach environments.

arkoses are more than 25 percent feldspar.[2]

The grains tend to be poorly rounded and

less well sorted than those of pure quartz sandstones. These feldspar-rich sandstones

come from rapidly eroding granitic and metamorphic terrains where chemical weathering

is subordinate to physical weathering.

lithic sandstones contain many lithic fragments derived from fine-grained rocks, mostly

shales, volcanic rocks, and fine-grained metamorphic rocks.

graywacke is a heterogeneous mixture of lithic fragments and angular grains of quartz

and feldspar, and/or grains surrounded by a fine-grained clay matrix. Much of this matrix

is formed by relatively soft fragments, such as shale and some volcanic rocks, that are

chemically altered and physically compacted after deep burial of the sandstone formation.

Eolianite is a term used for a rock which is composed of sand grains that show signs of

significant transportation by wind. These have usually been deposited in desert

environments. They are commonly extremely well sorted and rich in quartz.

Oolite is more a limestone than a sandstone, but is made of sand-sized carbonate ooids,

and is common in saline beaches with gentle wave action.

Sandstone composition is (generally) based on the make up of the framework, or sand-sized

grains in the sandstone. This is typically done by point-counting a thin section of the sandstone

using a method like the Gazzi-Dickinson Method. The composition of a sandstone can have

important information regarding the genesis of the sediment when used with QFL diagrams.

Shale

Shale

— Sedimentary Rock —

Shale

Composition

Clay minerals and quartz

Shale is a fine-grained, clastic sedimentary rock composed of mud that is a mix of flakes of clay

minerals and tiny fragments (silt-sized particles) of other minerals, especially quartz and calcite.

The ratio of clay to other minerals is variable.[1]

Shale is characterized by breaks along thin

laminae or parallel layering or bedding less than one centimeter in thickness, called fissility.[1]

Mudstones, on the other hand, are similar in composition but do not show the fissility.

Historical mining terminology

Before the mid 19th century, the terms slate, shale and schist were not sharply distinguished.

[2] In

the context of underground coal mining, shale was frequently referred to as slate well into the

20th

century.[3]

Texture

Shale typically exhibits varying degrees of fissility breaking into thin layers, often splintery and

usually parallel to the otherwise indistinguishable bedding plane because of parallel orientation

of clay mineral flakes.[1]

Non-fissile rocks of similar composition but made of particles smaller

than 0.06 mm are described as mudstones (1/3 to 2/3 silt particles) or claystone (less than 1/3

silt). Rocks with similar particle sizes but with less clay (greater than 2/3 silt) and therefore

grittier are siltstones.[1]

Shale is the most common sedimentary rock.[4]

Sample of drill cuttings of shale while drilling an oil well in Louisiana. Sand grain = 2 mm. in dia.

Composition and color

Shales are typically composed of variable amounts of clay minerals and quartz grains and the

typical color is gray. Addition of variable amounts of minor constituents alters the color of the

rock. Black shale results from the presence of greater than one percent carbonaceous material

and indicates a reducing environment.[1]

Black shale can also be referred to as black metal.[5]

Red, brown and green colors are indicative of ferric oxide (hematite - reds), iron hydroxide

(goethite - browns and limonite - yellow), or micaceous minerals (chlorite, biotite and illite -

greens).[1]

Clays are the major constituent of shales and other mudrocks. The clay minerals represented are

largely kaolinite, montmorillonite and illite. Clay minerals of Late Tertiary mudstones are

expandable smectites whereas in older rocks especially in mid to early Paleozoic shales illites

predominate. The transformation of smectite to illite produces silica, sodium, calcium,

magnesium, iron and water. These released elements form authigenic quartz, chert, calcite,

dolomite, ankerite, hematite and albite, all trace to minor (except quartz) minerals found in

shales and other mudrocks.[1]

Shales and mudrocks contain roughly 95 percent of the organic matter in all sedimentary rocks.

However, this amounts to less than one percent by mass in an average shale. Black shales which

form in anoxic conditions contain reduced free carbon along with ferrous iron (Fe2+

) and sulfur

(S2-

). Pyrite and amorphous iron sulfide along with carbon produce the black coloration.[1]

Formation

Limey shale overlaid by limestone, Cumberland Plateau, Tennessee

The process in the rock cycle which forms shale is compaction. The fine particles that compose

shale can remain suspended in water long after the larger and denser particles of sand have

deposited. Shales are typically deposited in very slow moving water and are often found in lakes

and lagoonal deposits, in river deltas, on floodplains and offshore from beach sands. They can

also be deposited on the continental shelf, in relatively deep, quiet water.

'Black shales' are dark, as a result of being especially rich in unoxidized carbon. Common in

some Paleozoic and Mesozoic strata, black shales were deposited in anoxic, reducing

environments, such as in stagnant water columns. Some black shales contain abundant heavy

metals such as molybdenum, uranium, vanadium, and zinc.[6][7][8]

The enriched values are of

controversial origin, having been alternatively attributed to input from hydrothermal fluids

during or after sedimentation or to slow accumulation from sea water over long periods of

sedimentation.[7][9][10]

Splitting the shale with a large knife to reveal fossils.

Fossils, animal tracks/burrows and even raindrop impact craters are sometimes preserved on

shale bedding surfaces. Shales may also contain concretions consisting of pyrite, apatite, or

various carbonate minerals.

Shales that are subject to heat and pressure of metamorphism alter into a hard, fissile,

metamorphic rock known as slate. With continued increase in metamorphic grade the sequence is

phyllite, then schist and finally to gneiss.

Weathering shale at a road cut in southeastern Kentucky

When shale is hit against other rock it can emit sparks of various colors like blue, green,

purple,yellow, and white. dependent on the type of rock it is hit against.

Limestone

Limestone

— Sedimentary Rock —

Limestone in Waitomo District, New Zealand

Composition

Calcium carbonate: inorganic crystalline calcite and/or

organic calcareous material.

Limestone is a sedimentary rock composed largely of the minerals calcite and/or aragonite,

which are different crystal forms of calcium carbonate (CaCO3).

Like most other sedimentary rocks, limestones are composed of grains; however, most grains in

limestone are skeletal fragments of marine organisms such as coral or foraminifera. Other

carbonate grains comprising limestones are ooids, peloids, intraclasts, and extraclasts. Some

limestones do not consist of grains at all and are formed completely by the chemical precipitation

of calcite or aragonite. i.e. travertine.

The solubility of limestone in water and weak acid solutions leads to karst landscapes. Regions

overlying limestone bedrock tend to have fewer visible groundwater sources (ponds and

streams), as surface water easily drains downward through joints in the limestone. While

draining, water and organic acid from the soil slowly (over thousands or millions of years)

enlarges these cracks; dissolving the calcium-carbonate and carrying it away in solution. Most

cave systems are through limestone bedrock.

Description

Limestone quarry at Cedar Creek, Virginia, USA.

La Zaplaz formations in Piatra Craiului Mountains

Limestone often contains variable amounts of silica in the form of chert (aka chalcedony, flint,

jasper, etc.) or siliceous skeletal fragment (sponge spicules, diatoms, radiolarians), as well as

varying amounts of clay, silt and sand sized terrestrial detritus carried in by rivers. The primary

source of the calcite in limestone is most commonly marine organisms. These organisms secrete

shells made of aragonite or calcite and leave these shells behind after the organism dies. Some of

these organisms can construct mounds of rock known as reefs, building upon past generations.

Below about 3,000 meters, water pressure and temperature causes the dissolution of calcite to

increase non-linearly so that limestone typically does not form in deeper waters (see lysocline).

Secondary calcite may also be deposited by supersaturated meteoric waters (groundwater that

precipitates the material in caves). This produces speleothems such as stalagmites and stalactites.

Another form taken by calcite is that of oolites (oolitic limestone) which can be recognized by its

granular appearance.

Limestone makes up about 10% of the total volume of all sedimentary rocks.[1][2]

Limestones

may also form in both lacustrine and evaporite depositional environments.[3][4]

Calcite can be either dissolved by groundwater or precipitated by groundwater, depending on

several factors including the water temperature, pH, and dissolved ion concentrations. Calcite

exhibits an unusual characteristic called retrograde solubility in which it becomes less soluble in

water as the temperature increases.

When conditions are right for precipitation, calcite forms mineral coatings that cement the

existing rock grains together or it can fill fractures.

Karst topography and caves develop in carbonate rocks due to their solubility in dilute acidic

groundwater. Cooling groundwater or mixing of different groundwaters will also create

conditions suitable for cave formation.

Coastal limestones are often eroded by organisms which bore into the rock by various means.

This process is known as bioerosion. It is most common in the tropics, and it is known

throughout the fossil record (see Taylor and Wilson, 2003).

Because of impurities, such as clay, sand, organic remains, iron oxide and other materials, many

limestones exhibit different colors, especially on weathered surfaces. Limestone may be

crystalline, clastic, granular, or massive, depending on the method of formation. Crystals of

calcite, quartz, dolomite or barite may line small cavities in the rock. Folk and Dunham

classifications are used to describe limestones more precisely.

Travertine is a banded, compact variety of limestone formed along streams, particularly where

there are waterfalls and around hot or cold springs. Calcium carbonate is deposited where

evaporation of the water leaves a solution that is supersaturated with chemical constituents of

calcite. Tufa, a porous or cellular variety of travertine, is found near waterfalls. Coquina is a

poorly consolidated limestone composed of pieces of coral or shells.

During regional metamorphism that occurs during the mountain building process (orogeny)

limestone recrystallizes into marble.

Limestone is a parent material of Mollisol soil group.

Types

Main article: List of types of limestone

Limestone landscape

Main article: Karst topography

The Cudgel of Hercules, a tall limestone rock and Pieskowa Skała Castle in the background.

Limestone is partially soluble, especially in acid, and therefore forms many erosional landforms.

These include limestone pavements, pot holes, cenotes, caves and gorges. Such erosion

landscapes are known as karsts. Limestone is less resistant than most igneous rocks, but more

resistant than most other sedimentary rocks. Limestone is therefore usually associated with hills

and downland and occurs in regions with other sedimentary rocks, typically clays.

Bands of limestone emerge from the Earth's surface in often spectacular rocky outcrops and

islands. Examples include the Burren in Co. Clare, Ireland; the Verdon Gorge in France; Malham

Cove in North Yorkshire and the Isle of Wight,[5]

England; on Fårö near the Swedish island of

Gotland, the Niagara Escarpment in Canada/United States, Notch Peak in Utah, the Ha Long Bay

National Park in Vietnam and the hills around the Lijiang River and Guilin city in China.

The Florida Keys, islands off the south coast of Florida, are composed mainly of oolitic

limestone (the Lower Keys) and the carbonate skeletons of coral reefs (the Upper Keys), which

thrived in the area during interglacial periods when sea level was higher than at present.

Unique habitats are found on alvars, extremely level expanses of limestone with thin soil

mantles. The largest such expanse in Europe is the Stora Alvaret on the island of Öland, Sweden.

Another area with large quantities of limestone is the island of Gotland, Sweden. Huge quarries

in northwestern Europe, such as those of Mount Saint Peter (Belgium/Netherlands), extend for

more than a hundred kilometers.

The world's largest limestone quarry is at Michigan Limestone and Chemical Company in

Rogers City, Michigan.[6]

The Great Pyramid of Giza. One of the Seven Wonders of the Ancient World, the structure is made

entirely from limestone.

Courthouse built of limestone in Manhattan, Kansas

Gallery

Limestone

cropping out at

São Pedro de

Moel beach,

Marinha Grande,

Portugal.

A stratigraphic section of

Ordovician limestone

exposed in central

Tennessee, U.S. The less-

resistant and thinner beds

are composed of shale.

Vertical lines are drill holes

for explosives used during

road construction.

Thin-section view of a

Middle Jurassic limestone

in southern Utah. The

round grains are ooids; the

largest is 1.2 mm in

diameter. This limestone is

an oosparite.

Photo and Etched

section of a sample of

fossiliferous limestone

from the Kope

Formation near

Cincinnati, Oh

Karst topography

The Kravice waterfall on the Trebižat river in Bosnia and Herzegovina has karst geology.

A karst landscape in Minerve, Hérault, France.

The karst hills of The Burren on the west coast of Ireland

Karst topography is a landscape shaped by the dissolution of a layer or layers of soluble

bedrock, usually carbonate rock such as limestone or dolomite.[1]

Due to subterranean drainage, there may be very limited surface water, even to the absence of all

rivers and lakes. Many karst regions display distinctive surface features, with sinkholes or

dolines being the most common. However, distinctive karst surface features may be completely

absent where the soluble rock is mantled, such as by glacial debris, or confined by a

superimposed non-soluble rock strata. Some karst regions include thousands of caves, even

though evidence of caves that are big enough for human exploration is not a required

characteristic of karst.

Various karst landforms have been found on all continents except Antarctica (see below: Notable

karst areas).

Background

Karst topography is characterized by subterranean limestone caverns, carved by groundwater.

The geographer Jovan Cvijić (1865–1927) was born in western Serbia and studied widely in the

Dinaric Kras region. His publication of Das Karstphänomen (1893) established that rock

dissolution was the key process and that it created most types of dolines, "the diagnostic karst

landforms". The Dinaric Kras thus became the type area for dissolutional landforms and

aquifers; the regional name kras, Germanicised as "karst", is now applied to modern and paleo-

dissolutional phenomena worldwide. Cvijić related the complex behaviour of karstic aquifers to

development of solutional conduit networks and linked it to a cycle of landform evolution. After

Cvijić, two main kinds of karstic areas exist: holokarst i.e. karst developed at whole as it is

Dinaric region along eastern Adriatic coast comprises deep in the inland of Balkan Peninsula and

merokarst developed imperfectly with some karstic forms as it is in eastern Serbia. He is

recognized as "the father of karst geomorphology".

The international community has settled on karst, the German name for Kras, a region in

Slovenia partially extending into Italy, where it is called "Carso" and where the first scientific

research of a karst topography was made. The name has an Indo-European origin (from karra

meaning "stone")[2]

, and in antiquity it was called "Carusardius" in Latin. The Slovene form

grast is attested since 1177, and the Croatian kras since 1230.[citation needed]

. "Krš" - "Krsh"

meaning in Serbo-Croatian "barren land" which is typical feature in the Northern Dinaric

limestone mountains could also be a origin to the word Karst.

Chemistry

Karst lake (Doberdò del Lago, Italy), from underground water springing into a depression. This lake has

no surface inlet or outlet.

Karst landforms are generally the result of mildly acidic water acting on soluble bedrock such as

limestone or dolostone. The carbonic acid that causes these features is formed as rain passes

through the atmosphere picking up CO2, which dissolves in the water. Once the rain reaches the

ground, it may pass through soil that may provide further CO2 to form a weak carbonic acid

solution: H2O + CO2 → H2CO3 (the acid).

Recent studies of sulfates, in karst waters, suggests sulfuric acid and hydrosulfuric acid may also

play an important role in karst formation. One such study, in the Frasassi Caves of Italy, showed

the oxidation of sulfuric acid is one of the leading corrosion factors in karst formation. As water

seeps into karst caves it brings in oxygen which reacts with H2S to form sulfuric acid (H2SO4)

and a hydronium (H3O). The H3O reacts with the limestone causing increased erosion within the

formation.

2O2 + H2S → H2SO4

H2SO4 + 2H2O → 2H3O + SO22-

H2O + CaCO3 → Ca2+ + HCO- + H2O

As a result of this reaction the mineral gypsum forms as a replacement mineral since it provides

many similar structures to the dissolution and redeposition of calcium carbonate.[3]

This mildly acidic water begins to dissolve the surface along fractures or bedding planes in the

limestone bedrock. Over time, these fractures enlarge as the bedrock continues to dissolve.

Openings in the rock increase in size, and an underground drainage system begins to develop,

allowing more water to pass through the area, and accelerating the formation of underground

karst features.[4]

Morphology

Limestone pavement in Dent de Crolles, France

The karstification of a landscape may result in a variety of large or small scale features both on

the surface and beneath. On exposed surfaces, small features may include flutes, runnels, clints

and grikes, collectively called karren or lapiez. Medium-sized surface features may include

sinkholes or cenotes (closed basins), vertical shafts, foibe (inverted funnel shaped sinkholes),

disappearing streams, and reappearing springs. Large-scale features may include limestone

pavements, poljes and blind valleys. Mature karst landscapes, where more bedrock has been

removed than remains, may result in karst towers, or haystack/eggbox landscapes. Beneath the

surface, complex underground drainage systems (such as karst aquifers) and extensive caves and

cavern systems may form.

The Witch's Finger stalagmite in Carlsbad Caverns, USA

Erosion along limestone shores, notably in the tropics, produces karst topography that includes a

sharp makatea surface above the normal reach of the sea and undercuts that are mostly the result

of biological activity or bioerosion at or a little above mean sea level. Some of the most dramatic

of these formations can be seen in Thailand's Phangnga Bay and Halong Bay in Vietnam.

Calcium carbonate dissolved into water may precipitate out where the water discharges some of

its dissolved carbon dioxide. Rivers which emerge from springs may produce tufa terraces,

consisting of layers of calcite deposited over extended periods of time. In caves, a variety of

features collectively called speleothems are formed by deposition of calcium carbonate and other

dissolved minerals.

Hydrology

A karst spring in the Jura mountains near Ouhans in eastern France at the source of the river Loue

Farming in karst areas must take into account the lack of surface water. The soils may be fertile

enough, and rainfall may be adequate, but rainwater quickly moves through the crevices into the

ground, sometimes leaving the surface soil parched between rains.

A karst fenster is where an underground stream emerges onto the surface between layers of rock,

cascades some feet, and then disappears back down, often into a sinkhole. Rivers in karst areas

may disappear underground a number of times and spring up again in different places, usually

under a different name (like Ljubljanica, the river of seven names). An example of this is the

Popo Agie River in Fremont County, Wyoming. At a site simply named "The Sinks" in Sinks

Canyon State Park, the river flows into a cave in a formation known as the Madison Limestone,

and then rises again a half-mile down the canyon in a placid pool. A Turlach is a unique type of

seasonal lake found in Irish karst areas which are formed through the annual welling-up of water

from the underground water system.

Water supplies from wells in karst topography may be unsafe, as the water may have run

unimpeded from a sinkhole in a cattle pasture, through a cave and to the well, bypassing the

normal filtering that occurs in a porous aquifer. Karst formations are cavernous and therefore

have high rates of permeability, resulting in reduced opportunity for contaminants to be filtered

out.

Groundwater in karst areas is just as easily polluted as surface streams. Sinkholes have often

been used as farmstead or community trash dumps. Overloaded or malfunctioning septic tanks in

karst landscapes may dump raw sewage directly into underground channels.

The karst topography itself also poses difficulties for human inhabitants. Sinkholes can develop

gradually as surface openings enlarge, but quite often progressive erosion is unseen and the roof

of an underground cavern suddenly collapses. Such events have swallowed homes, cattle, cars,

and farm machinery.

The Driftless Area National Wildlife Refuge in Iowa protects Discus macclintocki, a species of

ice age snail surviving in air chilled by flowing over buried karst ice formations.

Pseudokarst

Pseudokarsts are similar in form or appearance to karst features, but are created by different

mechanisms. Examples include lava caves and granite tors—for example, Labertouche Cave in

Victoria, Australia and paleocollapse features.

Notable karst areas

Africa

Madagascar

Anjajavy Forest, western Madagascar Ankarana Reserve, Madagascar Madagascar dry deciduous forests, western

Madagascar Tsingy de Bemaraha Strict Nature Reserve,

Madagascar

Asia

Poland

Kraków-Częstochowa Upland (Jura Krakowsko-Częstochowska)

Holy Cross Mountains (Góry Świętokrzyskie) with the Jaskinia Raj (Raj Cave)

Tatra Mountains including the Jaskinia Wielka Śnieżna (Great Snowy Cave)— the longest cave in Poland

Romania

Apuseni Mountains, Romania Bucegi Mountains, Romania

Serbia

Dinaric Alps region

Phong Nha Cave in Phong Nha-Ke Bang, Vietnam

China

Area around Guilin and Yangshuo Jiuzhaigou Valley and Huanglong National

Park, (UNESCO World Heritage Site) South China Karst, World Heritage Site Stone Forest Zhangjiajie National Forest park, forming

part of the Wulingyuan scenic area, World Heritage Site

Dunnieh mountains, North Lebanon

merokarst of eastern Serbia

Scotland

Assynt, southeast Skye and near Kentallen in Scotland, United Kingdom[9]

Slovakia

Slovak Paradise, Slovak Karst and Muránska planina, Slovakia

Slovenia

Region of Inner Carniola, Goriška, Upper Carniola and Lower Carniola

Kras (German: Karst), a plateau in southwestern Slovenia and northeastern Italy

Spain

El Torcal (Antequera - Spain)

Chalk

Chalk

— Sedimentary Rock —

The Needles, situated on the Isle of Wight, are part of the

extensive Southern England Chalk Formation.

Composition

calcite (calcium carbonate)

Chalk (pronounced /ˈtʃɔːk/) is a soft, white, porous sedimentary rock, a form of limestone

composed of the mineral calcite. Calcite is calcium carbonate or CaCO3. It forms under relatively

deep marine conditions from the gradual accumulation of minute calcite plates (coccoliths) shed

from micro-organisms called coccolithophores. It is common to find chert or flint nodules

embedded in chalk. Chalk can also refer to other compounds including magnesium silicate and

calcium sulfate.

Chalk is resistant to weathering and slumping compared to the clays with which it is usually

associated, thus forming tall steep cliffs where chalk ridges meet the sea. Chalk hills, known as

chalk downland, usually form where bands of chalk reach the surface at an angle, so forming a

scarp slope. Because chalk is porous it can hold a large volume of ground water, providing a

natural reservoir that releases water slowly through dry seasons.

Deposits

The Chalk Group is a European stratigraphic unit deposited during the late Cretaceous Period. It

forms the famous White Cliffs of Dover in Kent, England, as well as their counterparts of the

Cap Blanc Nez on the other side of the Dover Strait. The Champagne region of France is mostly

underlain by chalk deposits, which contain artificial caves used for wine storage. Some of the

highest chalk cliffs in the world occur at Møns Klint in Denmark.

Formation

Ninety million years ago the chalk downland of Northern Europe was ooze at the bottom of a

great sea. Protozoans such as foraminifera lived on the marine debris that showered down from

the upper layers of the ocean. Their bodies were made of chalk extracted from the rich sea-water.

As they died a deep layer built up and slowly became consolidated into rock. At a later date the

sea-bed became dry land, as earth movements thrust it upward.

Composition

Chalk is composed mostly of calcium carbonate with minor amounts of silt and clay. It is

normally formed underwater, commonly on the sea bed, then consolidated and compressed

during diagenesis into the form commonly seen today. During diagenesis silica accumulates to

form chert or flint nodules within the carbonate rock.

Coquina

Coquina outcropping on the beach at Washington Oaks State Gardens, Florida

Coquina (Spanish, "cockle"; pronounced /koʊˈkiːnə/) is an incompletely consolidated

sedimentary rock. Coquina was formed in association with marine reefs and is a variety of "coral

rag", technically a subset of limestone.

Composition and distribution

Coquina is mainly composed of mineral calcite, often including some phosphate, in the form of

seashells or coral. It is found in surface exposures along the east coast of Florida from St. Johns

County to Palm Beach County. It may occur up to 20 miles inland from the coast in the sub-

surface. It is found as far north as Fort Fisher, North Carolina. It has also been formed in the

South Island of New Zealand, where it outcrops in a disused quarry near Oamaru. The Oligocene

deposits here are composed primarily of very well preserved brachiopod shells, in a matrix of

brachiopod, echinoid, and bryozoan detritus and foraminifera.

History and use

Coquina from Florida.

Close-up of coquina from Florida. The scale bar is 10 mm.

Occasionally quarried or mined and used as a building stone in Florida for over 400 years,

coquina forms the walls of the Castillo de San Marcos, Saint Augustine. The stone makes a very

good material for forts, particularly those built during the period of heavy cannon use. Because

of coquina's softness, cannon balls would sink into, rather than shatter or puncture, the walls of

the Castillo de San Marcos.

When first quarried, coquina is extremely soft. This softness makes it very easy to remove from

the quarry and cut into shape. However, the stone is also at first much too soft to be used for

building. In order to be used as a building material, the stone is left out to dry for approximately

one to three years, which causes the stone to harden into a usable, but still comparatively soft,

form.

Coquina has also been used as a source of paving material. It is usually poorly cemented and

easily breaks into component shell or coral fragments, which can be substituted for gravel or

crushed harder rocks. Large pieces of coquina of unusual shape are sometimes used as landscape

decoration.

Because coquina often includes a component of phosphate, it is sometimes mined for use as

fertilizer.

Notable exposures of coquina

Washington Oaks State Gardens, Flagler County, Florida House of Refuge, Hutchinson Island, Martin County, Florida Blowing Rocks Preserve (and along Country Club Road), Palm Beach County, Florida North Carolina Aquarium at Fort Fisher, features a "Coquina Outcrop Touch Pool"

Lithographic Limestone

A lithographic limestone printing plate after use to print a map. Note the uniform fine texture of the

stone.

Lithographic limestone is hard limestone that is sufficiently fine-grained, homogeneous and

defect free to be used for lithography. Geologists use the term lithographic texture to refer to a

grain size under 1/250 mm.[1]

The term sublithographic is sometimes used for homogeneous

fine-grained limestone with a somewhat coarser texture.[2]

Origin

The generally accepted theory for the origin of lithographic and sublithographic limestones is

that they were formed in shallow stagnant hypersaline and anoxic lagoons. The combination of

mild hypersalinity and low oxygen content is believed to have inhibited the formation of

microbial mats and prevented the invasion of bottom dwelling organisms. Microbial mats and

bottom dwelling organisms would have left fossils, and bottom dwelling organisms would have

churned the accumulating sediment, producing a less homogeneous rock. Stagnancy was

required to avoid churning or sculpting of the sediment by currents or wave action.[3][4]

Distribution

Europe

The original source for lithographic limestone was the Solnhofen Limestone named after the

quarries of Solnhofen where it was first found. This is a late Jurassic deposit, part of a deposit of

plattenkalk (a very fine-grained limestone that splits into thin plates, usually Micrite) that

extends through the Swabian Alb and Franconian Alb in Southern Germany.[5]

Only a small

fraction of plattenkalk is suitable for lithography.[4]

For many years, the Solnhofen deposits were the only source of lithographic limestone. French

lithographic limestone from quarries near Montdardier, about 6 km south of le Vigan, Gard was

exhibited at the Great Exhibition of 1851, where it earned an honorable mention. This stone is

from the upper Lias Group, from the early Jurassic.[6][7]

The largest lithographic printing stone

ever quarried came from Le Vigan, 230x150cm (90x59 in).[8]

Théophile Steinlen used a

comparable stone for some of his posters.[9]

Several quarries are visible today on the chalky

plateau above Montdardier, between 2 km north ( 43°56′54.46″N 3°35′3.44″E / 43.9484611°N

3.5842889°E), and 2 km west ( 43°55′52.61″N 3°33′38.78″E / 43.9312806°N 3.5607722°E) of

the town.

Shortly before 1867, a second lithographic limestone quarry was opened in France near Cerin

and Crey, Isère ( 45°46′45.77″N 5°33′14.06″E / 45.7793806°N 5.5539056°E).[10]

The

lithographic limestones of Cerin are from the Kimmeridgian stage of the Upper Jurassic, and as

with the Solnhofen deposits, they preserve numerous interesting fossils.[11]

Lithographic limestone from the Lower Cretaceous has been quarried near Santa Maria de Meià

on the south flank of the Serra del Montsec in Spain. In 1902, L. M. Vidal, a mining engineer,

recognized the importance of the fossils found there.[12]

The Americas

The American Lithographic Stone Company was organized in Louisville, Kentucky in late 1868.

It initially focused its operation on quarries in Overton County, Tennessee,[13]

but shortly before

1900, it opened a quarry at Brandenburg, Kentucky. This quarry was the only commercial source

of lithographic stone in the United States at the turn of the 20th century. Unlike the Solnhofen

stone, Kentucky lithographic limestone was slightly dolomitic, and it was judged to be

competitive with Solnhofen stone for some purposes, but not for the highest quality work.[14][15]

This stone source was sub-Carboniferous (Mississipian).[16]

In 1917, the Brandenburg quarry was

judged the most important source of Lithographic stone in the United States.[17]

Prior to 1916, the

output of the Brandenburg quarry was small, but in 1916, as World War I cut off access to

Solnhofen stone, the quarry produced 20 tons of finished lithographic stone.[18]

The Remains of

the Brandenburg Lithograph Quarry are located along the Buttermilk Falls Historic Walking

Trail ( 38°0′3.54″N 86°9′34.74″W / 38.0009833°N 86.15965°W).[19]

In 1903, Clement L. Webster discovered a bed of lithographic limestone about 2 miles southwest

of Orchard, Iowa. His company, the Interstate Investment & Development Company platted a

town named Lithograph City nearby and opened a quarry ( 43°11′38.2″N 92°48′59.52″W /

43.193944°N 92.8165333°W).[20][21]

The Lithograph City Formation of the Cedar Valley Group

straddles the border between the Middle and Late Devonian and was named for its exposure in

this quarry. Outcrops of this formation extend from near Cedar Falls, Iowa north into

Minnesota.[22]

The suitability of Lithograph City limestone for lithography was tested by A. B.

Hoen who reported that stone from two layers in the Lithograph City quarry was excellent for

lithography and finer grained than the finest Solnhofen stone.[23]

Lithograph City was an

important source of lithographic stone in the United States during World War I, but the quarries

closed as metal printing plates replaced stone. In 1918, the Devonian Products Company took

over the operation, focusing on the production of crushed rock and renaming the town

Devonia.[24]

By 1938, the town had disappeared.[25]

Oolite

Modern ooids from a beach on Joulter's Cay, The Bahamas

Ooids on the surface of a limestone; Carmel Formation (Middle Jurassic) of southern Utah

Thin-section of calcitic ooids from an oolite within the Carmel Formation (Middle Jurassic) of southern

Utah

Oolite (egg stone) is a sedimentary rock formed from ooids, spherical grains composed of

concentric layers. The name derives from the Hellenic word òoion for egg. Strictly, oolites

consist of ooids of diameter 0.25–2 mm; rocks composed of ooids larger than 2 mm are called

pisolites. The term oolith can refer to oolite or individual ooids.

Composition

Ooids are most commonly composed of calcium carbonate (calcite or aragonite), but can be

composed of phosphate, chert, dolomite or iron minerals, including hematite. Dolomitic and

chert ooids are most likely the result of the replacement of the original texture in limestone.

Oolitic hematite occurs at Red Mountain near Birmingham, Alabama, along with oolitic

limestone.

Oolites are often used in the home aquarium industry because their small grain size (0.2 to 1.22

mm) is ideal for shallow static beds and bottom covering of up to 1" in depth. Also known as

"oolitic" sand, the sugar-sized round grains of this sand pass easily through the gills of gobies

and other sand-sifting organisms. Importantly, this incredibly smooth sand promotes the growth

of bacteria, which are important biofilters in home aquaria. Because of its extremely small grain

size, oolitic sand has a lot of surface area, which promotes high bacterial growth.

Occurrence

Some exemplar oolitic limestone, a common term for an oolite, was formed in England during

the Jurassic period, and forms the Cotswold Hills on the Isle of Portland with its famous Portland

Stone,[1]

and part of the North Yorkshire Moors. A particular type, Bath Stone, gives the

buildings of the World Heritage City of Bath their distinctive appearance.

The islands of the Lower Keys in the Florida Keys, as well as some barrier islands east of Miami

bordering Biscayne Bay, are mainly oolitic limestone, which was formed by deposition when

shallow seas covered the area between periods of glaciation. The material consolidated and

eroded during later exposure above the ocean surface.

This type of limestone is also found in Indiana in the United States. The town of Oolitic, Indiana,

was founded for the trade of limestone and bears its name. Quarries in Oolitic, Bedford, Indiana,

and Bloomington, Indiana contributed the materials for such iconic US landmarks as the Empire

State Building and The Pentagon. The Soldiers' and Sailors' Monument in downtown

Indianapolis is built mainly of grey oolitic limestone.

The 1979 movie Breaking Away, centers around the sons of quarry workers in Bloomington, the

home of Indiana University. Many of the buildings on the Indiana University campus are built

with native oolitic limestone material.

Roggenstein is a term describing a specific type of oolite in which the cementing matter is

argillaceous

Travertine

Travertine terraces at Mammoth Hot Springs, Yellowstone National Park

Calcium-carbonate-encrusted, yet growing moss, early stage of porous travertine formation.

Travertine is a form of limestone deposited by mineral springs, especially hot springs.

Travertine often has a fibrous or concentric appearance and exists in white, tan and cream-

colored varieties. It is formed by a process of rapid precipitation of calcium carbonate, often at

the mouth of a hot spring or in a limestone cave. In the latter it can form stalactites, stalagmites

and other speleothems. It is frequently used in Italy and elsewhere as a building material.

Travertine is a terrestrial sedimentary rock, formed by the precipitation of carbonate minerals

from solution in ground and surface waters, and/or geothermally heated hot-springs.[1][2]

Similar

(but softer and extremely porous) deposits formed from ambient-temperature water are known as

tufa.

Features

Travertine forms the stalactites and stalagmites of limestone caves, and the filling of some veins

and hot spring conduits. Travertine forms from geothermal springs and is often linked to

siliceous systems which form siliceous sinter. [Macrophyte]]s, bryophytes, algae, cyanobacteria

and other organisms often colonise the surface of travertine and are preserved, giving travertine

its distinctive porosity.

Some springs have temperatures high enough to exclude macrophytes and bryophytes from the

deposits, consequently, deposits are generally less porous than tufa. Thermophilic microbes are

important in these environments and stromatolitic fabrics are common. When deposits are

apparently devoid of any biological component, they are often referred to as calcareous sinter.

Geochemistry

Modern travertine is formed from geothermally heated supersaturated alkaline waters, with

raised pCO2 (see partial pressure). On emergence, waters degas CO2 due to the lower

atmospheric pCO2, resulting in an increase in pH. Since carbonate solubility decreases with

increased pH,[3]

precipitation is induced. Supersaturation may be enhanced by factors leading to

a reduction in pCO2, for example increased air-water interactions at waterfalls may be

important,[4]

as may photosynthesis.[5]

Precipitation may also be enhanced by evaporation in

some springs.

Both calcite and aragonite are found in hot spring travertines; aragonite is preferentially

precipitated when temperatures are hot, while calcite dominates when temperatures are

cooler.[6][7]

When pure and fine, travertine is white, but often it is brown to yellow due to

impurities.

Travertine forming at Jupiter Terrace, Fountain Geyser Pool, Yellowstone National Park. Photo by Ansel

Adams, 1941.

Travertine may precipitate out directly onto rock and other inert materials as in Pamukkale or

Yellowstone for example. Travertine may also precipitate out onto growing moss as in Plitvice

Lakes.

Occurrence

Travertine waterfalls exist not only in the U.S. in Oklahoma and Texas but most famously in

Italy, in Tivoli and Guidonia Montecelio where we can find most important quarries since

Ancient Roman like the old quarry Bernini in Guidonia. The latter has a major historic value,

because it was one of the quarries that Gian Lorenzo Bernini selected material from to build the

famous (colonnato di Piazza S.Pietro ) The Colonnade of St. Peter's Square in Rome in 1500.

Travertine derives its name from the former town, known as Tibur in ancient Roman times. The

ancient name for the stone was lapis tiburtinus, meaning tibur stone, which was gradually

corrupted to travertine. Detailed studies of the Tivoli and Guidonia travertine deposits revealed

diurnal and annual rhythmic banding and laminae, which have potential use in geochronology.[8]

In Central Europe's last post-glacial palaeoclimatic optimum (Atlantic Period, 8000-5000 B.C.),

huge deposits of tufa formed from karst springs. Important geotopes are found at the Swabian

Alb, mainly in valleys at the foremost northwest ridge of the cuesta; in many valleys of the

eroded periphery of the karstic Franconian Jura; at the northern Alpine foothills; and the northern

Karst Alps. On a smaller scale, these karst processes are still working. Travertine has been an

important building material since the Middle Ages.

Travertine has formed sixteen huge, natural dams in a valley in Croatia known as Plitvice Lakes

National Park. Clinging to moss and rocks in the water, the travertine has built up over several

millennia to form waterfalls up to 70 metres (230 ft) in height.[9]

Cascades of natural lakes formed behind travertine dams can be seen in Mahallat, Abbass Abad,

and Atash Kooh in Iran; Pamukkale, Turkey; Band-i-Amir, Afghanistan; HuangLong Valley,

Sichuan, China; and Semuc Champey, Guatemala.

In the U.S., the most well-known place for travertine formation is Yellowstone National Park,

where the geothermal areas are rich in travertine deposits. Oklahoma has two parks are dedicated

to this natural wonder. Turner Falls, the tallest waterfall in Oklahoma, is a 77 feet (23 m) cascade

of spring water flowing over a travertine cave. Honey Creek feeds this waterfall and creates

miles of travertine shelves both up and downstream. Many small waterfalls upstream in the

dense woods repeat the travertine-formation effect. The city of Davis now owns thousands of

acres of this land and has made it a tourist attraction. Another travertine resource is in Sulphur,

Oklahoma, 10 miles (16 km) east of Turner Falls. Travertine Creek flows through a spring-water

nature preserve within the boundaries of the Chickasaw National Recreation Area.

In Texas, the city of Austin and its surrounding "Hillcountry" to the south is built on limestone.

The area has many travertine formations, such as those found at Gorman Falls within Colorado

Bend State Park, the nature preserve known as Hamilton Pool, the West Cave Preserve, and

Krause Springs in Spicewood.

Tufa

Tufa columns at Mono Lake, California.

Tufa is a variety of limestone, formed by the precipitation of carbonate minerals from ambient

temperature water bodies. Geothermally heated hot-springs sometimes produce similar (but less

porous) carbonate deposits known as travertine. Tufa is sometimes referred to as (meteogene)

travertine;[1]

care must be taken when searching through literature to prevent confusion with hot

spring (thermogene) travertine. Calcareous tufa should not be confused with tuff, a porous

volcanic rock with parallel etymological origins.

Classification and features

Modern and fossil tufa deposits abound with wetland plants;[2]

as such many tufa deposits are

characterised by their large macrobiological component and are highly porous. Tufa forms either

in fluvial channels or in lacustrine settings. Ford and Pedley (1996)[3]

provide a review of tufa

systems worldwide.

Barrage Tufa at Cwm Nash, South Wales

Fluvial deposits

Deposits can be classified by their depositional environment (or otherwise by vegetation or

petrographically). Pedley (1990)[4]

provides an extensive classification system, which includes

the following classes of fluvial tufa:-

Spring - Deposits form on emergence from a spring/seep. Morphology can vary from mineratrophic wetlands to spring aprons (see calcareous sinter)

Braided channel - Deposits form within a fluvial channel, dominated by oncoids (see oncolite) Cascade - Deposits form at waterfalls, deposition is focussed here due to accelerated flow (see

Geochemistry) Barrage - Deposits form as a series of phytoherm barrages across a channel, which may grow up

to several metres in height. Barrages often contain a significant detrital component, composed of organic material (leaf litter, branches etc.).

Lacustrine deposits

Tufa deposits and columns and their flora on the lakeside of Mono Lake.

Lacustrine tufas are generally formed at the periphery of lakes and build up phytoherms

(freshwater reefs) and stromatolites. Oncoids are also common in these environments.

Other deposits

While fluvial and lacustrine systems make up the bulk of tufa systems worldwide, there are

several other important tufa environments.

Calcareous Sinter

Although sometimes regarded as a distinct carbonate deposit, calcareous sinter formed from

ambient temperature water can be considered a sub-type of tufa.

Speleothems

Calcareous speleothems may be regarded as a form of calcareous sinter. They lack any

significant macrophyte component due to the absence of light, for this reason they are often

morphologically closer to travertine or calcareous sinter.

Tufa columns

Tufa columns are an unusual form of tufa typically associated with saline lakes. They are distinct

from most tufa deposits in that they lack any significant macrophyte component; this is due to

the salinity excluding mesophilic organisms.[3]

Some tufa columns may actually form from hot-

springs and therefore actually be a form of travertine. It is generally thought that such features

form from CaCO3 precipitated when carbonate rich source waters emerge into alkaline soda

lakes. They have also been found in marine settings.[5]

Biology

Tufa deposits form an important habitat for a diverse flora. Bryophytes (non-vascular land

plants) and diatoms are well represented. The porosity of the deposits creates a wet habitat ideal

for these plants.

Geochemistry

Modern tufa is formed from supersaturated alkaline waters, with raised pCO2. On emergence,

waters degas CO2 due to the lower atmospheric pCO2 (see partial pressure), resulting in an

increase in pH. Since carbonate solubility decreases with increased pH,[6]

precipitation is

induced. Supersaturation may be enhanced by factors leading to a reduction in pCO2, for

example increased air-water interactions at waterfalls may be important,[7]

as may

photosynthesis.[8]

Recently it has been demonstrated that microbially induced precipitation may be more important

than physico-chemical precipitation. Pedley et al. (2009)[9]

showed with flume experiments that

precipitation does not occur unless a biofilm is present, despite supersaturation.

Calcite is the dominant mineral precipitate found; however, the polymorph aragonite is also

found.

Occurrence

Tufa is common in many parts of the world. Some notable deposits include:-

Pyramid Lake, Nevada, USA - tufa formations Mono Lake, California, USA- tufa columns Trona Pinnacles, California, USA - tufa columns North Dock Tufa, United Kingdom

Some sources suggest that "tufa" was used as the primary building material for most of the

châteaux of the Loire Valley, France. This results from a mis-translation of the terms "tuffeau

jaune" and "tuffeau blanc", which are porous varieties of the Late Cretaceous marine limestone

known as chalk.[10][11]

Dolomite

Dolomite

Dolomite and magnesite - Spain

General

Category Carbonate mineral

Chemical formula CaMg(CO3)2

Identification

Color white, gray to pink

Crystal habit

tabular crystals, often with curved

faces, also columnar, stalactitic,

granular, massive.

Crystal system trigonal - rhombohedral, bar3

Twinning common as simple contact twins

Cleavage rhombohedral cleavage (3 planes)

Fracture brittle - conchoidal

Mohs scale

hardness 3.5 to 4

Luster vitreous to pearly

Streak white

Specific gravity 2.84–2.86

Optical properties Uniaxial (-)

Refractive index nω = 1.679–1.681 nε = 1.500

Birefringence δ = 0.179–0.181

Solubility

Poorly soluble in dilute HCl unless

powdered.

Other

characteristics

May fluoresce white to pink under UV;

triboluminescent.

References [1][2][3][4]

Dolomite druse with calcite crystals from Lawrence County, Arkansas, USA (size: 17.0 x 6.3 x 2.8 cm)

Dolomite.

Dolomite (pronounced /ˈdɒləmaɪt/) is the name of a sedimentary carbonate rock and a mineral,

both composed of calcium magnesium carbonate CaMg(CO3)2 found in crystals.

Dolomite rock (also dolostone) is composed predominantly of the mineral dolomite. Limestone

that is partially replaced by dolomite is referred to as dolomitic limestone, or in old U.S. geologic

literature as magnesian limestone. Dolomite was first described in 1791 as the rock by the French

naturalist and geologist, Déodat Gratet de Dolomieu (1750–1801) for exposures in what are now

known as the Dolomite Alps of northern Italy.

Properties

The mineral dolomite crystallizes in the trigonal-rhombohedral system. It forms white, gray to

pink, commonly curved crystals, although it is usually massive. It has physical properties similar

to those of the mineral calcite, but does not rapidly dissolve or effervesce (fizz) in dilute

hydrochloric acid unless it is scratched or in powdered form. The Mohs hardness is 3.5 to 4 and

the specific gravity is 2.85. Refractive index values are nω = 1.679 - 1.681 and nε = 1.500.

Crystal twinning is common. A solid solution series exists between dolomite and iron rich

ankerite. Small amounts of iron in the structure give the crystals a yellow to brown tint.

Manganese substitutes in the structure also up to about three percent MnO. A high manganese

content gives the crystals a rosy pink color noted in the image above. A series with the

manganese rich kutnohorite may exist. Lead and zinc also substitute in the structure for

magnesium.

Formation

Dolomite bedrock underneath a Bristlecone Pine, White Mountains, California.

Vast deposits are present in the geological record, but the mineral is relatively rare in modern

environments. Laboratory synthesis of stoichiometric dolomite has been carried out only at

temperatures of greater than 100 degrees Celsius (conditions typical of burial in sedimentary

basins), even though much dolomite in the rock record appears to have formed in low-

temperature conditions. The high temperature is likely to speed up the movement of calcium and

magnesium ions so that they can find their places in the ordered structure within a reasonable

amount of time. This suggests that the lack of dolomite that is being formed today is likely due to

kinetic factors. I.e. due to the lack of kinetic energy or temperature.

Modern dolomite does occur as a precipitating mineral in specialized environments on the

surface of the earth today. In the 1950s and 60s, dolomite was found to be forming in highly

saline lakes in the Coorong region of South Australia. Dolomite crystals also occur in deep-sea

sediments, where organic matter content is high. This dolomite is termed "organogenic"

dolomite.

Recent research has found modern dolomite formation under anaerobic conditions in

supersaturated saline lagoons along the Rio de Janeiro coast of Brazil, namely, Lagoa Vermelha

and Brejo do Espinho. One interesting reported case was the formation of dolomite in the

kidneys of a Dalmatian dog.[citation needed]

This was believed to be due to chemical processes

triggered by bacteria. Dolomite has been speculated to develop under these conditions with the

help of sulfate-reducing bacteria.[citation needed]

The actual role of bacteria in the low-temperature formation of dolomite remains to be

demonstrated. The specific mechanism of dolomitization, involving sulfate-reducing bacteria,

has not yet been demonstrated.[5]

Dolomite appears to form in many different types of environment and can have varying

structural, textural and chemical characteristics. Some researchers have stated "there are

dolomites and dolomites", meaning that there may not be one single mechanism by which

dolomite can form. Much modern dolomite differs significantly from the bulk of the dolomite

found in the rock record, leading researchers to speculate that environments where dolomite

formed in the geologic past differ significantly from those where it forms today.

Reproducible laboratory syntheses of dolomite (and magnesite) leads first to the initial

precipitation of a metastable "precursor" (such as magnesium calcite), to be changed gradually

into more and more of the stable phase (such as dolomite or magnesite) during periodical

intervals of dissolution and reprecipitation. The general principle governing the course of this

irreversible geochemical reaction has been coined Ostwald's step rule.

For a very long time scientists had difficulties synthesizing dolomite. However, in a 1999 study,

through a process of dissolution alternating with intervals of precipitation, measurable levels of

dolomite were synthesized at low temperatures and pressures.[6]

Coral atolls

Dolomitization of calcite also occurs at certain depths of coral atolls where water is

undersaturated in calcium carbonate but saturated in dolomite. Convection created by tides and

sea currents enhance this change. Hydrothermal currents created by volcanoes under the atoll

may also play an important role.

Magnesite

Magnesite

General

Category Carbonate mineral

Chemical formula MgCO3

Identification

Color Colorless, white, pale yellow, pale

brown, faintly pink, lilac-rose

Crystal habit

Usually massive, rarely as

rhombohedrons or hexagonal prisms

Crystal system

Trigonal - Hexagonal Scalenohedral H-M

Symbol 32/m Space Group: R3c

Cleavage [1011] perfect

Fracture Conchoidal

Tenacity Brittle

Mohs scale

hardness 3.5 - 4.5

Luster Vitreous

Streak White

Diaphaneity Transparent to translucent

Specific gravity 3.0 - 3.2

Optical

properties Uniaxial (-)

Refractive index nω=1.508 - 1.510 nε=1.700

Fusibility Infusible

Solubility Effervesces in hot HCl

Other

characteristics

May exhibit pale green to pale blue

fluorescence and phosphorescence

under UV; triboluminescent

References [1][2][3][4]

Magnesite is magnesium carbonate, MgCO3. Iron (as Fe2+

) substitutes for magnesium (Mg) with

a complete solution series with siderite, FeCO3. Calcium, manganese, cobalt, and nickel may

also occur in small amounts. Dolomite, (Mg,Ca)CO3, is almost indistinguishable from magnesite.

Occurrence

Magnesite occurs as veins in and an alteration product of ultramafic rocks, serpentinite and other

magnesium rich rock types in both contact and regional metamorphic terranes. These magnesites

often are cryptocrystalline and contain silica as opal or chert.

Magnesite is also present within the regolith above ultramafic rocks as a secondary carbonate

within soil and subsoil, where it is deposited as a consequence of dissolution of magnesium-

bearing minerals by carbon dioxide within groundwaters.

Formation

Magnesite can be formed via talc carbonate metasomatism of peridotite and other ultrabasic

rocks. Magnesite is formed via carbonation of olivine in the presence of water and carbon

dioxide, and is favored at moderate temperatures and pressures typical of greenschist facies;

Magnesite can also be formed via the carbonation of magnesian serpentine (lizardite) via the

following reaction:

Serpentine + carbon dioxide → Talc + magnesite + Water

2Mg3Si2O5(OH)4 + 3CO2 → Mg3Si4O10(OH)2 + 3MgCO3 + H2O

Forsterite magnesia-rich olivine compositions favor production of magnesite from peridotite.

Fayalitic (iron-rich) olivine favors production of magnetite-magnesite-silica compositions.

Magnesite can also be formed from metasomatism in skarn deposits, in dolomitic limestones,

associated with wollastonite, periclase, and talc.

Magnesite is also found in a number of Precambrian carbonate hosted sediments, and is thought

to have formed as an evaporite.

Geyserite

Geyserite from Iceland

Geyserite is a form of opaline silica that is often found around hot springs and geysers.

Botryoidal geyserite is known as fiorite. It is sometimes referred to as sinter.

Diatomaceous earth

A sample of diatomaceous earth

Diatomaceous earth (pronounced /ˌdaɪ.ətəˌmeɪʃəs ˈɜrθ/) also known as diatomite or kieselgur,

is a naturally occurring, soft, siliceous sedimentary rock that is easily crumbled into a fine white

to off-white powder. It has a particle size ranging from less than 1 micron to more than 1

millimeter, but typically 10 to 200 microns.[1]

This powder has an abrasive feel, similar to

pumice powder, and is very light, due to its high porosity. The typical chemical composition of

oven dried diatomaceous earth is 80 to 90% silica, with 2 to 4% alumina (attributed mostly to

clay minerals) and 0.5 to 2% iron oxide.[1]

Diatomaceous earth consists of fossilized remains of diatoms, a type of hard-shelled algae. It is

used as a filtration aid, as a mild abrasive, as a mechanical insecticide, as an absorbent for

liquids, as cat litter, as an activator in blood clotting studies, and as a component of dynamite. As

it is also heat-resistant, it can be used as a thermal insulator.

Geology and occurrence

Diatomaceous earth as viewed under bright field illumination on a light microscope. Diatomaceous earth

is made up of the cell walls/shells of single cell diatoms and readily crumbles to a fine powder. Diatom

cell walls are made up of biogenic silica; silica synthesised in the diatom cell by the polymerisation of

silicic acid. This image of diatomaceous earth particles in water is at a scale of 6.236 pixels/μm, the

entire image covers a region of approximately 1.13 by 0.69 mm.

Formation

Diatomite forms by the accumulation of the amorphous silica (opal, SiO2·nH2O) remains of dead

diatoms (microscopic single-celled algae) in lacustrine or marine sediments. The fossil remains

consist of a pair of symmetrical shells or frustules.[1]

Discovery

In 1836 or 1837, the peasant and goods waggoner, Peter Kasten,[2]

discovered kieselgur when

sinking a well on the northern slopes of the Haußelberg hill, in the Lüneburg Heath in north

Germany. Initially, it was thought that limestone had been found, which could be used as

fertiliser. Alfred Nobel used the properties of kieselgur in the manufacture of dynamite. The

Celle engineer, Wilhelm Berkefeld, recognised its ability to filter, and developed 'filter candles'

fired from kieselgur.[3]

During the cholera epidemic in Hamburg in 1892, these Berkefeld filters

were used successfully.

Extraction and storage sites in the Lüneburg Heath

Neuohe - Abbau from 1863 to 1994 Wiechel from 1871 to 1978 Hützel from 1876 to 1969 Hösseringen from ca. 1880 to 1894 Hammerstorf from ca. 1880 to 1920 Oberohe from 1884 to 1970 Schmarbeck from 1896 to ca. 1925 Steinbeck from 1897 to 1928 Breloh from 1907 to 1975 Schwindebeck from 1913 to 1975 Hetendorf from 1970 to 1994

The deposits are up to 28 metres thick and are all of freshwater kieselgur.

ca. 1900–1910

Kieselgur pit at Neuohe

ca.1900–1910 a drying area: one

firing pile is being prepared;

another is under way

1913 Staff at the Neuohe factory,

with workers and a female cook in

front of a drying shed

Until the First World War almost the entire worldwide production of kieselgur was from this

region.

Other deposits

In Germany kieselgur was also extracted at Altenschlirf [4]

on the Vogelsberg (Upper Hesse) and

at Klieken [5]

(Saxony-Anhalt).

There is a layer of kieselgur up to 4 metres thick in the nature reserve of Soos in the Czech

Republic.

In Colorado and in Clark, Nevada (USA), there are deposits that are up to several hundred metres

thick in places.

Sometimes kieselgur is found on the surface in deserts. Research has shown that the erosion of

kieselgur in such areas (such as the Bodélé Depression in the Sahara) is one of the most

important sources of climate-affecting dust in the atmosphere.

The commercial deposits of diatomite are restricted to Tertiary or Quaternary periods. Older

deposits from as early as the Cretaceous Period are known, but are of low quality.[6]

Marine

deposits have been worked in the Sisquoc Formation in Santa Barbara County, California near

Lompoc and along the southern California coast. Additional marine deposits have been worked

in Maryland, Virginia, Algeria and the MoClay of Denmark. Fresh water lake deposits occur in

Nevada, Oregon, Washington and California. Lake deposits also occur in interglacial lakes in the

eastern US and Canada and in Europe in Germany, France, Denmark and the Czech Republic.

The worldwide association of diatomite deposits and volcanic deposits suggests that the

availability of silica from volcanic ash may be needed for thick diatomite deposits.[6]

Evaporite

Cobble encrusted with halite evaporated from the Dead Sea, Israel.

Evaporites (pronounced /ɨˈvæpəraɪt/) are water-soluble mineral sediments that result from the

evaporation of bodies of surficial water. Evaporites are considered sedimentary rocks.

Formation of evaporite rocks

Although all water bodies on the surface and in aquifers contain dissolved salts, the water must

evaporate into the atmosphere for the minerals to precipitate. For this to happen the water body

must enter a restricted environment where water input into this environment remains below the

net rate of evaporation. This is usually an arid environment with a small basin fed by a limited

input of water. When evaporation occurs, the remaining water is enriched in salts, and they

precipitate when the water becomes oversaturated.

Evaporite depositional environments

Evaporite depositional environments which meet the above conditions include;

Graben areas and half-grabens within continental rift environments fed by limited riverine drainage, usually in subtropical or tropical environments

o Example environments at the present which match this is the Denakil Depression, Ethiopia; Death Valley, California

Graben environments in oceanic rift environments fed by limited oceanic input, leading to eventual isolation and evaporation

o Examples include the Red Sea, and the Dead Sea in Jordan and Israel Internal drainage basins in arid to semi-arid temperate to tropical environments fed by

ephemeral drainage o Example environments at the present include the Simpson Desert, Western Australia,

the Great Salt Lake in Utah Non-basin areas fed exclusively by groundwater seepage from artesian waters

o Example environments include the seep-mounds of the Victoria Desert, fed by the Great Artesian Basin, Australia

Restricted coastal plains in regressive sea environments o Examples include the sabkha deposits of Iran, Saudi Arabia and the Red Sea; the

Garabogazköl of the Caspian Sea Drainage basins feeding into extremely arid environments

o Examples include the Chilean deserts, certain parts of the Sahara and the Namib

The most significant known evaporite depositions happened during the Messinian salinity crisis

in the basin of the Mediterranean.

Evaporitic formations

Hopper crystal cast of halite in a Jurassic rock, Carmel Formation, southwestern Utah.

Evaporite formations need not be composed entirely of halite salt. In fact, most evaporite

formations do not contain more than a few percent of evaporite minerals, the remainder being

composed of the more typical detrital clastic rocks and carbonates.

For a formation to be recognised as evaporitic it may simply require recognition of halite

pseudomorphs, sequences composed of some proportion of evaporite minerals, and recognition

of mud crack textures or other textures.

Economic importance of evaporites

Evaporites are important economically because of their mineralogy, their physical properties in-

situ and their behaviour within the subsurface.

Evaporite minerals, especially nitrate minerals, are economically important in Peru and Chile.

Nitrate minerals are often mined for use in the production on fertilizer and explosives.

Thick halite deposits are expected to become an important location for the disposal of nuclear

waste because of their geologic stability, predictable engineering and physical behaviour and

imperviousness to groundwater.

Halite formations are famous for their ability to form diapirs which produce ideal locations for

trapping petroleum deposits.

Major groups of evaporite minerals

Halides: halite, sylvite (KCl), and fluorite Sulfates: such as gypsum, barite, and anhydrite Nitrates: nitratine (soda niter) and niter Borates: typically found in arid-salt-lake deposits plentiful in the southwestern US. A common

borate is borax, which has been used in soaps as a surfactant. Carbonates: such as trona, formed in inland brine lakes.

Evaporite minerals start to precipitate when their concentration in water reaches such a level that

they can no longer exist as solutes.

The minerals precipitate out of solution in the reverse order of their solubilities, such that the

order of precipitation from sea water is

1. Calcite (CaCO3) and dolomite (CaMg(CO3)2) 2. Gypsum (CaSO4-2H2O) and anhydrite (CaSO4). 3. Halite (i.e. common salt, NaCl) 4. Potassium and magnesium salts

The abundance of rocks formed by seawater precipitation is in the same order as the precipitation

given above. Thus, limestone (calcite) and dolomite are more common than gypsum, which is

more common than halite, which is more common than potassium and magnesium salts.

Metamorphic rock

Quartzite, a form of metamorphic rock, from the Museum of Geology at University of Tartu collection.

Metamorphic rock is the transformation of an existing rock type, the protolith, in a process

called metamorphism, which means "change in form". The protolith is subjected to heat and

pressure (temperatures greater than 150 to 200 °C and pressures of 1500 bars[1]

) causing

profound physical and/or chemical change. The protolith may be sedimentary rock, igneous rock

or another older metamorphic rock. Metamorphic rocks make up a large part of the Earth's crust

and are classified by texture and by chemical and mineral assemblage (metamorphic facies).

They may be formed simply by being deep beneath the Earth's surface, subjected to high

temperatures and the great pressure of the rock layers above it. They can form from tectonic

processes such as continental collisions, which cause horizontal pressure, friction and distortion.

They are also formed when rock is heated up by the intrusion of hot molten rock called magma

from the Earth's interior. The study of metamorphic rocks (now exposed at the Earth's surface

following erosion and uplift) provides us with information about the temperatures and pressures

that occur at great depths within the Earth's crust.

Some examples of metamorphic rocks are gneiss, slate, marble, schist, and quartzite.

Metamorphic minerals

Metamorphic minerals are those that form only at the high temperatures and pressures associated

with the process of metamorphism. These minerals, known as index minerals, include

sillimanite, kyanite, staurolite, andalusite, and some garnet.

Other minerals, such as olivines, pyroxenes, amphiboles, micas, feldspars, and quartz, may be

found in metamorphic rocks, but are not necessarily the result of the process of metamorphism.

These minerals formed during the crystallization of igneous rocks. They are stable at high

temperatures and pressures and may remain chemically unchanged during the metamorphic

process. However, all minerals are stable only within certain limits, and the presence of some

minerals in metamorphic rocks indicates the approximate temperatures and pressures at which

they formed.

The change in the particle size of the rock during the process of metamorphism is called

recrystallization. For instance, the small calcite crystals in the sedimentary rock limestone

change into larger crystals in the metamorphic rock marble, or in metamorphosed sandstone,

recrystallization of the original quartz sand grains results in very compact quartzite, in which the

often larger quartz crystals are interlocked. Both high temperatures and pressures contribute to

recrystallization. High temperatures allow the atoms and ions in solid crystals to migrate, thus

reorganizing the crystals, while high pressures cause solution of the crystals within the rock at

their point of contact.

Foliation

Folded foliation in a metamorphic rock from near Geirangerfjord, Norway

The layering within metamorphic rocks is called foliation (derived from the Latin word folia,

meaning "leaves"), and it occurs when a rock is being shortened along one axis during

recrystallization. This causes the platy or elongated crystals of minerals, such as mica and

chlorite, to become rotated such that their long axes are perpendicular to the orientation of

shortening. This results in a banded, or foliated, rock, with the bands showing the colors of the

minerals that formed them.

Textures are separated into foliated and non-foliated categories. Foliated rock is a product of

differential stress that deforms the rock in one plane, sometimes creating a plane of cleavage. For

example, slate is a foliated metamorphic rock, originating from shale. Non-foliated rock does not

have planar patterns of strain.

Rocks that were subjected to uniform pressure from all sides, or those that lack minerals with

distinctive growth habits, will not be foliated. Slate is an example of a very fine-grained, foliated

metamorphic rock, while phyllite is medium, schist coarse, and gneiss very coarse-grained.

Marble is generally not foliated, which allows its use as a material for sculpture and architecture.

Another important mechanism of metamorphism is that of chemical reactions that occur between

minerals without them melting. In the process atoms are exchanged between the minerals, and

thus new minerals are formed. Many complex high-temperature reactions may take place, and

each mineral assemblage produced provides us with a clue as to the temperatures and pressures

at the time of metamorphism.

Metasomatism is the drastic change in the bulk chemical composition of a rock that often occurs

during the processes of metamorphism. It is due to the introduction of chemicals from other

surrounding rocks. Water may transport these chemicals rapidly over great distances. Because of

the role played by water, metamorphic rocks generally contain many elements absent from the

original rock, and lack some that originally were present. Still, the introduction of new chemicals

is not necessary for recrystallization to occur.

Types of metamorphism

Contact metamorphism

A contact metamorphic rock made of interlayered calcite and serpentine from the Precambrian of

Canada. Once thought to be a fossil called Eozoön canadense. Scale in mm.

Contact metamorphism is the name given to the changes that take place when magma is

injected into the surrounding solid rock (country rock). The changes that occur are greatest

wherever the magma comes into contact with the rock because the temperatures are highest at

this boundary and decrease with distance from it. Around the igneous rock that forms from the

cooling magma is a metamorphosed zone called a contact metamorphism aureole. Aureoles may

show all degrees of metamorphism from the contact area to unmetamorphosed (unchanged)

country rock some distance away. The formation of important ore minerals may occur by the

process of metasomatism at or near the contact zone.

When a rock is contact altered by an igneous intrusion it very frequently becomes more

indurated, and more coarsely crystalline. Many altered rocks of this type were formerly called

hornstones, and the term hornfels is often used by geologists to signify those fine grained,

compact, non-foliated products of contact metamorphism. A shale may become a dark

argillaceous hornfels, full of tiny plates of brownish biotite; a marl or impure limestone may

change to a grey, yellow or greenish lime-silicate-hornfels or siliceous marble, tough and

splintery, with abundant augite, garnet, wollastonite and other minerals in which calcite is an

important component. A diabase or andesite may become a diabase hornfels or andesite hornfels

with development of new hornblende and biotite and a partial recrystallization of the original

feldspar. Chert or flint may become a finely crystalline quartz rock; sandstones lose their clastic

structure and are converted into a mosaic of small close-fitting grains of quartz in a metamorphic

rock called quartzite.

If the rock was originally banded or foliated (as, for example, a laminated sandstone or a foliated

calc-schist) this character may not be obliterated, and a banded hornfels is the product; fossils

even may have their shapes preserved, though entirely recrystallized, and in many contact-altered

lavas the vesicles are still visible, though their contents have usually entered into new

combinations to form minerals that were not originally present. The minute structures, however,

disappear, often completely, if the thermal alteration is very profound; thus small grains of quartz

in a shale are lost or blend with the surrounding particles of clay, and the fine ground-mass of

lavas is entirely reconstructed.

By recrystallization in this manner peculiar rocks of very distinct types are often produced. Thus

shales may pass into cordierite rocks, or may show large crystals of andalusite (and chiastolite),

staurolite, garnet, kyanite and sillimanite, all derived from the aluminous content of the original

shale. A considerable amount of mica (both muscovite and biotite) is often simultaneously

formed, and the resulting product has a close resemblance to many kinds of schist. Limestones, if

pure, are often turned into coarsely crystalline marbles; but if there was an admixture of clay or

sand in the original rock such minerals as garnet, epidote, idocrase, wollastonite, will be present.

Sandstones when greatly heated may change into coarse quartzites composed of large clear

grains of quartz. These more intense stages of alteration are not so commonly seen in igneous

rocks, because their minerals, being formed at high temperatures, are not so easily transformed

or recrystallized.

In a few cases rocks are fused and in the dark glassy product minute crystals of spinel, sillimanite

and cordierite may separate out. Shales are occasionally thus altered by basalt dikes, and

feldspathic sandstones may be completely vitrified. Similar changes may be induced in shales by

the burning of coal seams or even by an ordinary furnace.

There is also a tendency for metasomatism between the igneous magma and sedimentary country

rock, whereby the chemicals in each are exchanged or introduced into the other. Granites may

absorb fragments of shale or pieces of basalt. In that case, hybrid rocks called skarn arise, which

don't have the characteristics of normal igneous or sedimentary rocks. Sometimes an invading

granite magma permeates the rocks around, filling their joints and planes of bedding, etc., with

threads of quartz and feldspar. This is very exceptional but instances of it are known and it may

take place on a large scale.[2]

Regional metamorphism

Mississippian marble in Big Cottonwood Canyon, Wasatch Mountains, Utah.

Regional metamorphism is the name given to changes in great masses of rock over a wide area.

Rocks can be metamorphosed simply by being at great depths below the Earth's surface,

subjected to high temperatures and the great pressure caused by the immense weight of the rock

layers above. Much of the lower continental crust is metamorphic, except for recent igneous

intrusions. Horizontal tectonic movements such as the collision of continents create orogenic

belts, and cause high temperatures, pressures and deformation in the rocks along these belts. If

the metamorphosed rocks are later uplifted and exposed by erosion, they may occur in long belts

or other large areas at the surface. The process of metamorphism may have destroyed the original

features that could have revealed the rock's previous history. Recrystallization of the rock will

destroy the textures and fossils present in sedimentary rocks. Metasomatism will change the

original composition.

Regional metamorphism tends to make the rock more indurated and at the same time to give it a

foliated, shistose or gneissic texture, consisting of a planar arrangement of the minerals, so that

platy or prismatic minerals like mica and hornblende have their longest axes arranged parallel to

one another. For that reason many of these rocks split readily in one direction along mica-bearing

zones (schists). In gneisses, minerals also tend to be segregated into bands; thus there are seams

of quartz and of mica in a mica schist, very thin, but consisting essentially of one mineral. Along

the mineral layers composed of soft or fissile minerals the rocks will split most readily, and the

freshly split specimens will appear to be faced or coated with this mineral; for example, a piece

of mica schist looked at facewise might be supposed to consist entirely of shining scales of mica.

On the edge of the specimens, however, the white folia of granular quartz will be visible. In

gneisses these alternating folia are sometimes thicker and less regular than in schists, but most

importantly less micaceous; they may be lenticular, dying out rapidly. Gneisses also, as a rule,

contain more feldspar than schists do, and they are tougher and less fissile. Contortion or

crumbling of the foliation is by no means uncommon, and then the splitting faces are undulose or

puckered. Schistosity and gneissic banding (the two main types of foliation) are formed by

directed pressure at elevated temperature, and to interstitial movement, or internal flow arranging

the mineral particles while they are crystallizing in that directed pressure field.

Rocks that were originally sedimentary and rocks that were undoubtedly igneous convert into

schists and gneisses. If originally of similar composition they may be very difficult to distinguish

from one another if the metamorphism has been great. A quartz-porphyry, for example, and a

fine feldspathic sandstone, may both the converted into a grey or pink mica-schist.[2]

They are

made by heat and pressure.

Metamorphic rock textures

The five basic metamorphic textures with typical rock types are slaty (includes slate and phyllite)

(the foliation is called "slaty cleavage"), "schistose" (includes schist) (the foliation is called

("schistosity"), gneissose (gneiss) (the foliation is called "gneissosity"), granoblastic (includes

granulite, some marbles and quartzite), and hornfelsic (includes hornfels and skarn).

Gneiss

Gneiss rock

Augen gneiss from Rio de Janeiro, Brazil

Granitic gneiss from Enfield, New York

Study of Gneiss Rock, Glenfinlas, the Trossachs, Scotland. A pen and ink study by John Ruskin, 1853, is

now in the Ashmolean Museum, Oxford.

Gneiss (pronounced /ˈnaɪs/ "nice") is a common and widely distributed type of rock formed by

high-grade regional metamorphic processes from pre-existing formations that were originally

either igneous or sedimentary rocks.

Etymology

The etymology of the word "gneiss" is disputed. Some sources say it comes from the Middle

High German verb gneist (to spark; so called because the rock glitters) and has occurred in

English at least since 1757.[1]

Other sources claim the root to be an old Saxon mining term that

seems to have meant decayed, rotten, or possibly worthless material.[citation needed]

Composition

Gneissic rocks are usually medium- to coarse-foliated and largely recrystallized but do not carry

large quantities of micas, chlorite or other platy minerals. Gneisses that are metamorphosed

igneous rocks or their equivalent are termed granite gneisses, diorite gneisses, etc. Depending on

their composition, they may also be called garnet gneiss, biotite gneiss, albite gneiss, etc.

Gneiss displays compositional banding where the minerals are arranged into bands of more

mafic minerals and more felsic minerals. This is developed under high temperature and pressure

conditions.

Types of gneiss

Orthogneiss designates a gneiss derived from an igneous rock, and paragneiss is one from a

sedimentary rock. Gneissose is used to describe rocks with properties similar to gneiss.

Lewisian

Most of the Outer Hebrides of Scotland have a bedrock formed from Lewisian gneiss. These are

amongst the oldest rocks in Europe and some of the oldest in the world, having been formed in

the Precambrian "super-eon", up to 3 billion years ago. In addition to the Outer Hebrides, they

form basement deposits on the Scottish mainland west of the Moine Thrust and on the islands of

Coll and Tiree.[2]

These rocks are largely igneous in origin, mixed with metamorphosed marble,

quartzite and mica schist and intruded by later basaltic dykes and granite magma.[3]

The gneiss's

delicate pink colours are exposed throughout the islands and it is sometimes referred to by

geologists as "The Old Boy".[4]

Augen gneiss

Augen gneiss, from the German Augen [ˈaʊɡən], meaning "eye", is a coarse-grained gneiss,

interpreted as resulting from metamorphism of granite, which contains characteristic elliptic or

lenticular shear bound feldspar porphyroclasts, normally microcline, within the layering of the

quartz, biotite and magnetite bands.

Schist

Schist

The schists constitute a group of medium-grade metamorphic rocks, chiefly notable for the

preponderance of lamellar minerals such as micas, chlorite, talc, hornblende, graphite, and

others. Quartz often occurs in drawn-out grains to such an extent that a particular form called

quartz schist is produced. By definition, schist contains more than 50% platy and elongated

minerals, often finely interleaved with quartz and feldspar. Schist is often garnetiferous.

The individual mineral grains in schist, drawn out into flaky scales by heat and pressure, can be

seen by the naked eye. Schist is characteristically foliated, meaning the individual mineral grains

split off easily into flakes or slabs. The word schist is derived from the Greek word σχίζειν

meaning "to split", which is a reference to the ease with which schists can be split along the

plane in which the platy minerals lie.

Most schists have been derived from clays and muds which have passed through a series of

metamorphic processes involving the production of shales, slates and phyllites as intermediate

steps. Certain schists have been derived from fine-grained igneous rocks such as basalts and

tuffs. Most schists are mica schists, but graphite and chlorite schists are also common.

Thin section of Garnet-Mica-Schist

View of cut Garnet-Mica-Schist

Schists are named for their prominent or perhaps unusual mineral constituents, such as garnet

schist, tourmaline schist, glaucophane schist, etc.

Schists are frequently used as dimension stone.

Historical mining terminology

Before the mid 19th century, the terms slate, shale and schist were not sharply distinguished.

[1] In

the context of underground coal mining, shale was frequently referred to as slate well into the

20th

century.[2]

Formation

During metamorphism, rocks which were originally sedimentary or igneous are converted into

schists and gneisses. If the composition of the rocks was originally similar, they may be very

difficult to distinguish from one another if the metamorphism has been great. A quartz-porphyry,

for example, and a fine grained feldspathic sandstone, may both be converted into a grey or pink

mica-schist. Usually, however, it is possible to distinguish between sedimentary and igneous

schists and gneisses. If the whole district, for example, occupied by these rocks have traces of

bedding, clastic structure, or unconformability then it may be a sign that the original rock was

sedimentary. In other cases intrusive junctions, chilled edges, contact alteration or porphyritic

structure may prove that in its original condition a metamorphic gneiss was an igneous rock. The

last appeal is often to the chemistry, for there are certain rock types which occur only as

sediments, while others are found only among igneous masses, and however advanced the

metamorphism may be, it rarely modifies the chemical composition of the mass very greatly.

Such rocks, for example, as limestones, dolomites, quartzites and aluminous shales have very

definite chemical characters which distinguish them even when completely recrystallized.

Manhattan schist, from Southeastern New York

Chlorite schist forms from shale or mudstone.

The schists are classified principally according to the minerals they consist of and on their

chemical composition. For example, many metamorphic limestones, marbles, and calc-schists,

with crystalline dolomites, contain silicate minerals such as mica, tremolite, diopside, scapolite,

quartz and feldspar. They are derived from calcareous sediments of different degrees of purity.

Another group is rich in quartz (quartzites, quartz schists and quartzose gneisses), with variable

amounts of white and black mica, garnet, feldspar, zoisite and hornblende. These were once

sandstones and arenaceous rocks. The graphitic schists may readily be believed to represent

sediments once containing coal or plant remains; there are also schistose ironstones (hematite-

schists), but metamorphic beds of salt or gypsum are exceedingly uncommon. Among schists of

igneous origin there are the silky calc-schists, the foliated serpentines (once ultramafic masses

rich in olivine), and the white mica-schists, porphyroids and banded halleflintas, which have

been derived from rhyolites, quartz-porphyries and felsic tuffs. The majority of mica-schists,

however, are altered claystones and shales, and pass into the normal sedimentary rocks through

various types of phyllite and mica-slates. They are among the most common metamorphic rocks;

some of them are graphitic and others calcareous. The diversity in appearance and composition is

very great, but they form a well-defined group not difficult to recognize, from the abundance of

black and white micas and their thin, foliated, schistose character. A subgroup is the andalusite,

staurolite, kyanite and sillimanite-schists which usually make their appearance in the vicinity of

gneissose granites, and have presumably been affected by contact metamorphism.[3]

Quartzite

Quartzite

Swan Peak Quartzite (Ordovician) exposed just north of Tony Grove Lake, Cache County, Utah.

The quartzite of the Prospect Mountain Formation on top of Wheeler Peak, the highest peak within

Nevada, United States.

Quartzite (from German Quarzit[1]

) is a hard metamorphic rock which was originally

sandstone.[2]

Sandstone is converted into quartzite through heating and pressure usually related to

tectonic compression within orogenic belts. Pure quartzite is usually white to grey, though

quartzites often occur in various shades of pink and red due to varying amounts of iron oxide

(Fe2O3). Other colors, such as yellow and orange, are due to other mineral impurities.

When sandstone is metamorphosed to quartzite, the individual quartz grains recrystallize along

with the former cementing material to form an interlocking mosaic of quartz crystals. Most or all

of the original texture and sedimentary structures of the sandstone are erased by the

metamorphism. Minor amounts of former cementing materials, iron oxide, carbonate and clay,

often migrate during recrystallization and metamorphosis. This causes streaks and lenses to form

within the quartzite.

Orthoquartzite is a very pure quartz sandstone composed of usually well rounded quartz grains

cemented by silica. Orthoquartzite is often 99% SiO2 with only very minor amounts of iron oxide

and trace resistant minerals such as zircon, rutile and magnetite. Although few fossils are

normally present, the original texture and sedimentary structures are preserved. The term is often

misused, and should be used for only tightly-cemented metamorphic quartzites, not quartz-

cemented quartz arenites[3]

. The typical distinction between the two (since each is a gradation

into the other) is a proper quartzite is so highly cemented, diagentically altered, and

metamorphosed that it will fracture and break across grain boundaries, not around them.

Quartzite is very resistant to chemical weathering and often forms ridges and resistant hilltops.

The nearly pure silica content of the rock provides little to form soil from and therefore the

quartzite ridges are often bare or covered only with a very thin layer of soil and little vegetation.

Abandoned quartzite mine in Kakwa Provincial Park, British Columbia, Canada

Biface in quartzite - Stellenbosch, South Africa

Occurrences

In the United States, formations of quartzite can be found in some parts of Pennsylvania, eastern

South Dakota, Central Texas,[6]

southwest Minnesota,[7]

Devil's Lake State Park in the Baraboo

Hills in Wisconsin,[8]

the Wasatch Range in Utah,[9]

near Salt Lake City, Utah and as resistant

ridges in the Appalachians[10]

and other mountain regions. Quartzite is also found in the Morenci

Copper Mine in Arizona.[11]

The town of Quartzsite in western Arizona derives its name from the

quartzites in the nearby mountains in both Arizona and Southeastern California. A glassy

vitreous quartzite has been described from the Belt Supergroup in the Coeur d’Alene district of

northern Idaho.[12]

In the United Kingdom, a craggy ridge of quartzite called the Stiperstones (early Ordovician -

Arenig Epoch, 500 Ma) runs parallel with the Pontesford-Linley fault, 6 km north-west of the

Long Mynd in south Shropshire. Also to be found in England are the Cambrian "Wrekin

quartizite" (in Shropshire), and the Cambrian "Hartshill quartzite" (Nuneaton area).[13]

In Wales,

Holyhead mountain and most of Holy island off Anglesey sport excellent Precambrian quartzite

crags and cliffs. In the Scottish Highlands, several mountains (e.g. Foinaven, Arkle) composed of

Cambrian quartzite can be found in the far north-west Moine Thrust Belt running in a narrow

band from Loch Eriboll in a south-westerly direction to Skye.[14]

In Canada, the La Cloche Mountains in Ontario are composed primarily of white quartzite.

Slate

Slate

— Metamorphic Rock —

Slate

Composition

Primary quartz, muscovite/illite

Secondary biotite, chlorite, hematite, pyrite

Slate Macro (~ 6 cm long and ~ 4 cm high)

Slate is a fine-grained, foliated, homogeneous metamorphic rock derived from an original shale-

type sedimentary rock composed of clay or volcanic ash through low-grade regional

metamorphism. The result is a foliated rock in which the foliation may not correspond to the

original sedimentary layering. When expertly "cut" by striking with a specialized tool in the

quarry, many slates will form smooth flat sheets of stone which have long been used for roofing

and floor tiles and other purposes. Slate is frequently grey in color, especially when seen in

masse covering roofs. However, slate occurs in a variety of colors even from a single locality; for

example, slate from North Wales can be found in many shades of grey, from pale to dark, and

may also be purple, green or cyan. Slate is not to be confused with shale, from which it may be

formed, or schist. Ninety percent of Europe's natural slate used for roofing originates from

Spain.[1]

The word "slate" is also used for some objects made from slate. It may mean a single roofing

slate, or a writing slate, traditionally a small piece of slate, often framed in wood, used with

chalk as a notepad or noticeboard etc., and especially for recording charges in pubs and inns. The

phrase "clean slate" or "blank slate" comes from this use.

Historical mining terminology

Before the mid-19th century, the terms slate, shale and schist were not sharply distinguished.[2]

In the context of underground coal mining, the term slate was commonly used to refer to shale

well into the 20th century.[3]

For example, roof slate refers to shale above a coal seam, and draw

slate refers to roof slate that falls from the mine roof as the coal is removed.[4]

Mineral composition

Slate is mainly composed of quartz and muscovite or illite, often along with biotite, chlorite,

hematite, and pyrite and, less frequently, apatite, graphite, kaolin, magnetite, tourmaline, or

zircon as well as feldspar. Occasionally, as in the purple slates of North Wales, ferrous reduction

spheres form around iron nuclei, leaving a light green spotted texture. These spheres are

sometimes deformed by a subsequent applied stress field to ovoids, which appear as ellipses

when viewed on a cleavage plane of the specimen.

Slate extraction

Exhibition Slate mine Fell.

Historical Pit Vogelsberg 1 at Fell

Main article: Slate industry

In Eurasia

Slate-producing regions in Europe include Wales (see slate industry in Wales), Cornwall

(famously the village of Delabole), Cumbria (see Burlington Slate Quarries, Honister Slate Mine

and Skiddaw Slate) in the United Kingdom; parts of France (Anjou, Ardennes, Brittany, Savoie);

Belgium (Ardenne); Liguria in northern Italy, especially between the town of Lavagna (which

means chalkboard in Italian) and Fontanabuona valley; Portugal especially around Valongo in

the north of the country; Germany's (Moselle River-region, Hunsrück, Eifel, Westerwald,

Thuringia and north Bavaria); Alta, Norway (actually schist not a true slate) and Galicia. Some

of the slate from Wales and Cumbria is colored slate (non-blue): purple and formerly green in

Wales and green in Cumbria. China has vast slate deposits; in recent years its export of finished

and unfinished slate has increased, it has slate in various colors.

In the Americas

Slate is abundant in Brazil (the second-biggest producer of slate) around Papagaios in Minas

Gerais (responsible for 95% of the extraction of slate in Brazil). An independent report by

Consultant Geologist J. A. Walsh describes how certain products originating from Brazil on sale

in the UK, are not entitled to bear the CE mark.[7]

Other areas known for slate production are the east coast of Newfoundland, the Slate Belt of

Eastern Pennsylvania, Buckingham County Virginia (Buckingham Slate), and the Slate Valley of

Vermont and New York, where colored slate is mined in the Granville, New York area.

A major slating operation existed in Monson, Maine, during the late 19th- and early-20th

centuries. The slate found in Monson is usually a dark purple to blackish color, and many local

structures are still roofed with slate tiles. The roof of St. Patrick's Cathedral was made of roofing

slate from Monson, as was the headstone of John F. Kennedy.[citation needed]

Slate is also found in the Arctic and was used by the Inuit to make the blades for ulus.

Fossils

Shale can metamorphose into slate; sometimes the fossils may remain intact.

Because slate was formed in low heat and pressure, compared to a number of other metamorphic

rocks, some fossils can be found in slate; sometimes even microscopic remains of delicate

organisms.[8]

Marble

Marble.

Folded and weathered marble at General Carrera Lake, Chile.

The Taj Mahal is made of marble.

Ancient greek statue of Venus de Milo, sculpted from marble.

Natural patterns on the polished surface of Breccia or "landscape marble" can resemble a city skyline or

even trees, and were used as inlays for furniture etc.

Marble is a metamorphic rock composed of recrystallized carbonate minerals, most commonly

calcite or dolomite.

Geologists use the term "marble" to refer to metamorphosed limestone; however stonemasons

use the term more broadly to encompass unmetamorphosed limestone.[1]

Marble is commonly used for sculpture and as a building material.

Etymology

The word "marble" derives from the Greek "μάρμαρον" (mármaron),[2]

from "μάρμαρος"

(mármaros), "crystalline rock", "shining stone",[3][4]

perhaps from the verb "μαρμαίρω"

(marmaírō), "to flash, sparkle, gleam".[5]

This stem is also the basis for the English word

marmoreal, meaning "marble-like."

Whilst the English term resembles the French marbre, most other european languages (eg

Spanish mármol, Italian marmo, Portuguese mármore, German and Swedish marmor, Dutch

marmer, Polish marmur, Czech mramor and Russian мрáмор ) follow the original Greek.

Physical Origins

Marble is a rock resulting from metamorphism of sedimentary carbonate rocks, (most commonly

limestone or dolomite rock). Metamorphism causes variable recrystallization of the original

carbonate mineral grains.

The resulting marble rock is typically composed of an interlocking mosaic of carbonate crystals.

Primary sedimentary textures and structures of the original carbonate rock (protolith) have

typically been modified or destroyed.

Pure white marble is the result of metamorphism of a very pure (silicate-poor) limestone or

dolomite protolith. The characteristic swirls and veins of many colored marble varieties are

usually due to various mineral impurities such as clay, silt, sand, iron oxides, or chert which were

originally present as grains or layers in the limestone.

Green coloration is often due to serpentine resulting from originally high magnesium limestone

or dolostone with silica impurities. These various impurities have been mobilized and

recrystallized by the intense pressure and heat of the metamorphism.

Types

Examples of historically notable marble varieties and locations:

Marble name Color Location Country

Bucova

Marble

white, gray Băuţar, Caraş-Severin County (applied in

Ulpia Traiana Sarmizegetusa) Romania

Carrara

marble

white or blue-gray Carrara Italy

Macael

marble

white Macael, Almeria Spain

Makrana

Marble

white Makrana India

Murphy

Marble

white Pickens and Gilmer Counties, Georgia United States

Parian marble pure-white, fine-grained Island of Paros Greece

Pentelic

marble

pure-white, fine-grained

semitranslucent Penteliko Mountain, Athens Greece

Phrygian

Marble

purple Phrygia Turkey

Ruskeala

Marble

white near Ruskeala, Karelia Russia

Sienese

Marble

yellow, yellowish-white near Sovicille, Tuscany Italia

Bianco Sivec white near Prilep Republic of

Macedonia

Sylacauga

marble

white Talladega County, Alabama United States

Tennessee

marble

pale pink to cedar-red Knox, Blount and Hawkins Counties,

Tennessee United States

Vermont

Marble

white Proctor, Vermont United States

Yule Marble uniform pure white near Marble, Colorado United States

Wunsiedel

Marble

white Wunsiedel, Bavaria Germany

Serpentinite

A sample of serpentinite rock, partially made up of chrysotile

Boulder of Serpentinite at Soldiers Delight Natural Environmental Area, Maryland

Serpentinite is a rock composed of one or more serpentine group minerals. Minerals in this

group are formed by serpentinization, a hydration and metamorphic transformation of

ultramafic rock from the Earth's mantle. The alteration is particularly important at the sea floor at

tectonic plate boundaries.

Formation

Serpentinization is a geological low-temperature metamorphic process involving heat and water

in which low-silica mafic and ultramafic rocks are oxidized (anaerobic oxidation of Fe2+

by the

protons of water leading to the formation of H2) and hydrolyzed with water into serpentinite.

Peridotite, including dunite, at and near the seafloor and in mountain belts is converted to

serpentine, brucite, magnetite, and other minerals — some rare, such as awaruite (Ni3Fe), and

even native iron. In the process large amounts of water are absorbed into the rock increasing the

volume and destroying the structure.[1]

The density changes from 3.3 to 2.7 g/cm3 with a concurrent volume increase of about 40%. The

reaction is exothermic and large amounts of heat energy are produced in the process.[1]

Rock temperatures can be raised by about 260 °C,[1]

providing an energy source for formation of

non-volcanic hydrothermal vents. The magnetite-forming chemical reactions produce hydrogen

gas under anaerobic conditions prevailing deep in the mantle, far from the Earth atmosphere.

Carbonates and sulfates are subsequently reduced by hydrogen and form methane and hydrogen

sulfide. The hydrogen, methane, and hydrogen sulfide provide energy sources for deep sea

chemotroph microorganisms.[1]

Serpentinite reactions

Serpentinite is formed from olivine via several reactions, some of which are complementary.

Olivine is a solid solution between the magnesium-endmember forsterite and the iron-

endmember fayalite. Serpentinite reactions 1a and 1b, below, exchange silica between forsterite

and fayalite to form serpentine group minerals and magnetite. These are highly exothermic

reactions.

Reaction 1a:

Fayalite + water → magnetite + aqueous silica + hydrogen

3Fe2SiO4 + 2H2O → 2Fe3O4 + 3SiO2 + 2H2

Reaction 1b:

Forsterite + aqueous silica → serpentine

3Mg2SiO4 + SiO2 + 4H2O → 2Mg3Si2O5(OH)4

Reaction 1c:

Forsterite + water → serpentine + brucite

2Mg2SiO4 + 3H2O → Mg3Si2O5(OH)4 + Mg(OH)2

Reaction 1c describes the hydration of olivine with water only to yield serpentine and Mg(OH)2

(brucite). Serpentine is stable at high pH in the presence of brucite like calcium silicate hydrate,

(C-S-H) phases formed along with portlandite (Ca(OH)2) in hardened Portland cement paste after

the hydration of belite (Ca2SiO4), the artificial calcium equivalent of forsterite.

Analogy of reaction 1c with belite hydration in ordinary Portland cement:

Belite + water → C-S-H phase + portlandite

2 Ca2SiO4 + 4 H2O → 3 CaO · 2 SiO2 · 3 H2O + Ca(OH)2

After reaction, the poorly soluble reaction products (aqueous silica or dissolved magnesium ions)

can be transported in solution out of the serpentinized zone by diffusion or advection.

A similar suite of reactions involves pyroxene-group minerals, though less readily and with

complication of the additional end-products due to the wider compositions of pyroxene and

pyroxene-olivine mixes. Talc and magnesian chlorite are possible products, together with the

serpentine minerals antigorite, lizardite, and chrysotile. The final mineralogy depends both on

rock and fluid compositions, temperature, and pressure. Antigorite forms in reactions at

temperatures that can exceed 600°C during metamorphism, and it is the serpentine group mineral

stable at the highest temperatures. Lizardite and chrysotile can form at low temperatures very

near the Earth's surface. Fluids involved in serpentinite formation commonly are highly reactive

and may transport calcium and other elements into surrounding rocks; fluid reaction with these

rocks may create metasomatic reaction zones enriched in calcium and called rodingites.

In the presence of carbon dioxide, however, serpentinitization may form either magnesite

(MgCO3) or generate methane (CH4). It is thought that some hydrocarbon gases may be

produced by serpentinite reactions within the oceanic crust.

Reaction 2a:

Olivine + water + carbonic acid → serpentine + magnetite + methane

(Fe,Mg)2SiO4 + nH2O + CO2 → Mg3Si2O5(OH)4 + Fe3O4 + CH4

or, in balanced form:

18Mg2SiO4 + 6Fe2SiO4 + 26H2O + CO2 → 12Mg3Si2O5(OH)4 + 4Fe3O4 + CH4

Reaction 2b:

Olivine + water + carbonic acid → serpentine + magnetite + magnesite + silica

(Fe,Mg)2SiO4 + nH2O + CO2 → Mg3Si2O5(OH)4 + Fe3O4 + MgCO3 + SiO2

Reaction 2a is favored if the serpentinite is Mg-poor or if there isn't enough carbon dioxide to

promote talc formation. Reaction 2b is favored in highly magnesian compositions and low partial

pressure of carbon dioxide.

The degree to which a mass of ultramafic rock undergoes serpentinisation depends on the

starting rock composition and on whether or not fluids transport calcium, magnesium and other

elements away during the process. If an olivine composition contains sufficient fayalite, then

olivine plus water can completely metamorphose to serpentine and magnetite in a closed system.

In most ultramafic rocks formed in the Earth's mantle, however, the olivine is about 90%

forsterite endmember, and for that olivine to react completely to serpentine, magnesium must be

transported out of the reacting volume.

Serpentinitization of a mass of peridotite usually destroys all previous textural evidence because

the serpentine minerals are weak and behave in a very ductile fashion. However, some masses of

serpentinite are less severely deformed, as evidenced by the apparent preservation of textures

inherited from the peridotite, and the serpentinites may have behaved in a rigid fashion.

Hydrogen production by anaerobic oxidation of fayalite ferrous ions

In the absence of atmospheric oxygen (O2), in deep geological conditions prevailing far away

from Earth atmosphere, hydrogen (H2) is produced by the anaerobic oxidation of ferrous ions

(Fe2+

) present in the crystal lattice of the iron-endmember fayalite by the protons (H+) of water.

Considering three formula units of fayalite (Fe2(SiO4)) for the purpose of stoechiometry and

reaction mass balance, four ferrous ions will undergo oxidation by water protons while the two

remaining will stay unoxidised. Neglecting the orthosilicate anions not involved in the redox

process, it is then possible to schematically write the two half-redox reactions as follows:

4 (Fe2+ → Fe3+ + e–) (oxidation of ferrous ions)

2 (H2O + 2 e– → O2– + H2) (reduction of protons into hydrogen)

This leads to the global redox reaction involving ferrous ions oxidation by water:

4 Fe2+ + 2 H2O → 4 Fe3+ + 2 O2– + 2 H2

The two unoxidised ferrous (Fe2+

) ions still available in the three formula units of fayalite finally

combine with the four ferric (Fe3+

) cations and oxide anions (O2–

) to form two formula units of

magnetite (Fe3O4).

Finally, considering the required rearrangement of the orthosilicate anions into free silica (SiO2)

and free oxide anions (O2–

), it is possible to write the complete reaction of anaerobic oxidation

and hydrolysis of fayalite according to the following mass balance:

3 Fe2SiO4 + 2 H2O → 2 Fe3O4 + 3 SiO2 + 3 H2

fayalite + water → magnetite + quartz + hydrogen

This reaction closely resembles the Schikorr reaction observed in the anaerobic oxidation of the

ferrous hydroxide in contact with water:

3 Fe(OH)2 → Fe3O4 + 2 H2O + H2

ferrous hydroxide → magnetite + water + hydrogen

Carbon sequestration

Serpentinite, along with olivine, its precursor, has been proposed as an efficient reagent for

carbon sequestration using the magnesite reaction, mentioned hereabove, or a variation where

serpentine is reacted with carbon dioxide and hydrogen to form magnesite, magnetite, and silica.

The ideal composition of olivine or serpentinite for this process is thus highly magnesian, to

favor the production of magnesite and the fixation of carbon.

Swiss ovenstone

A variety of chlorite talc schist associated with Alpine serpentinite is found in Val d’Anniviers,

Switzerland and was used as ovenstone in stove construction.[2]

Serpentinization on Mars

The presence of traces of methane in the atmosphere of Mars has been hypothesized to be

possible evidence for life on Mars. Serpentinization has been proposed as an alternative non-

biological source for the observed methane traces.[3][4]

Hornfels

hornfels

Hornfels (German, meaning "hornstone," after its frequent association with glacial "horn peaks"

in the Alps, being a very hard rock and thus more likely to resist glacial action and form horn-

shaped peaks such as Matterhorn) is the group designation for a series of contact metamorphic

rocks that have been baked and indurated by the heat of intrusive igneous masses and have been

rendered massive, hard, splintery, and in some cases exceedingly tough and durable.

Most hornfels are fine-grained, and while the original rocks (such as sandstone, shale, slate,

limestone and diabase) may have been more or less fissile owing to the presence of bedding or

cleavage planes, this structure is effaced or rendered inoperative in the hornfels. Though they

may show banding, due to bedding, etc., they break across this as readily as along it; in fact, they

tend to separate into cubical fragments rather than into thin plates.

The most common hornfels (the biotite hornfelses ) are dark-brown to black with a somewhat

velvety luster owing to the abundance of small crystals of shining black mica. The lime hornfels

are often white, yellow, pale-green, brown and other colors. Green and darkgreen are the

prevalent tints of the hornfels produced by the alteration of igneous rocks. Although for the most

part the constituent grains are too small to be determined by the unaided eye, there are often

larger crystals of cordierite, garnet or andalusite scattered through the fine matrix, and these may

become very prominent on the weathered faces of the rock.

Structure

The structure of the hornfels is very characteristic. Very rarely do any of the minerals show

crystalline form, but the small grains fit closely together like the fragments of a mosaic; they are

usually of nearly equal dimensions. This has been called pfiaster or pavement structure from the

resemblance to rough pavement work. Each mineral may also enclose particles of the others; in

the quartz, for example, small crystals of graphite, biotite, iron oxides, sillimanite or feldspar

may appear in great numbers. Often the whole of the grains are rendered semi-opaque in this

way. The minutest crystals may show traces of crystalline outlines; undoubtedly they are of new

formation and have originated in situ. This leads us to believe that the whole rock has been

recrystallized at a high temperature and in the solid state so that there was little freedom for the

mineral molecules to build up well-individualized crystals. The regeneration of the rock has been

sufficient to efface most of the original structures and to replace the former minerals more-or-

less completely by new ones. But crystallization has been hampered by the solid condition of the

mass and the new minerals are formless and have been unable to reject impurities, but have

grown around them.

Compositions of Hornfels

Slates, shales and clays yield biotite hornfels in which the most conspicuous mineral is black

mica, the small scales of which are transparent under the microscope and have a dark reddish

brown color and strong dichroism. There is also quartz, and often a considerable amount of

feldspar, while graphite, tourmaline and iron oxides frequently occur in lesser quantity. In these

biotite hornfels the minerals, which consist of aluminiun silicates, are commonly found; they are

usually andalusite and sillimanite, but kyanite appears also in hornfels, especially in those that

have a schistose character. The andalusite may be pink and is then often pleochroic in thin

sections, or it may be white with the cross-shaped dark enclosures of the matrix that are

characteristic of chiastolite. Sillimanite usually forms exceedingly minute needles embedded in

quartz.

In the rocks of this group cordierite also occurs, not rarely, and may have the outlines of

imperfect hexagonal prisms that are divided up into six sectors when seen in polarized light. In

biotite hornfels, a faint striping may indicate the original bedding of the unaltered rock and

corresponds to small changes in the nature of the sediment deposited. More commonly there is a

distinct spotting, visible on the surfaces of the hand specimens. The spots are round or elliptical,

and may be paler or darker than the rest of the rock. In some cases they are rich in graphite or

carbonaceous matter; in others they are full of brown mica; some spots consist of rather coarser

grains of quartz than occur in the matrix. The frequency with which this feature reappears in the

less altered slates and hornfels is rather remarkable, especially as it seems certain that the spots

are not always of the same nature or origin. Tourmaline hornfels are found sometimes near the

margins of tourmaline granites; they are black with small needles of schorl that under the

microscope are dark brown and richly pleochroic. As the tourmaline contains boron, there must

have been some permeation of vapors from the granite into the sediments. Rocks of this group

are often seen in the Cornish tin-mining districts, especially near the ludes.

A second great group of hornfels are the calcite-silicate hornfels that arise from the thermal

alteration of impure limestone. The purer beds recrystallize as marbles, but where there has been

originally an admixture of sand or clay lime-bearing silicates are formed, such as diopside,

epidote, garnet, sphene, vesuvianite and scapolite; with these phlogopite, various feldspars,

pyrites, quartz and actinolite often occur. These rocks are fine-grained, and though often banded,

are tough and much harder than the original limestones. They are excessively variable in their

mineralogical composition, and very often alternate in thin seams with biotite hornfels and

indurated quartzites. When perfused with boric and fluoric vapors from the granite they may

contain much axinite, fluorite and datolite, but the altiminous silicates (andalusite, &c.) are

absent from these rocks.

From diabases, basalts, andesites and other igneous rocks a third type of hornfels is produced.

They consist essentially of feldspar with hornblende (generally of brown color) and pale

pyroxene. Sphene, biotite and iron oxides are the other common constituents, but these rocks

show much variety of composition and structure. Where the original mass was decomposed and

contained calcite, zeolites, chlorite and other secondary minerals either in veins or in cavities,

there are usually rounded a reas or irregular streaks containing a suite of new minerals, which

may resemble those of the calcium-silicate hornfelses above described. The original porphyritic,

fluidal, vesicular or fragmental structures of the igneous rock are clearly visible in the less

advanced stages of hornfelsing, but become less evident as the alteration progresses.

In some districts hornfelsed rocks occur that have acquired a schistose structure through

shearing, and these form transitions to schists and gneisses that contain the same minerals as the

hornfels, but have a schistose instead of a hornfels structure. Among these may be mentioned

cordierite and sillimanite gneisses, andalusite and kyanite mica-schists, and those schistose

calcite-silicate rocks that are known as cipolins. That these are sediments that have undergone

thermal alteration is generally admitted, but the exact conditions under which they were formed

are not always clear. The essential features of hornfelsing are ascribed to the action of heat,

pressure and permeating vapors, regenerating a rock mass without the production of fusion (at

least on a large scale). It has been argued, however, that often there is extensive chemical change

owing to the introduction of matter from the granite into the rocks surrounding it. The formation

of new feldspar in the hornfelses is pointed out as evidence of this. While this felspathization

may have occurred in a few localities, it seems conspicuously absent from others. Most

authorities at the present time regard the changes as being purely of a physical and not of a

chemical nature.