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Schematic representation of a metamorphic reaction. Abbreviations of minerals: act = actinolite; chl = chlorite; ep = epidote; gt = garnet; hbl = hornblende; plag = plagioclase. Two minerals represented in the figure do not participate in the reaction, they can be quartz and K-feldspar. This reaction takes place in nature when a mafic rock goes from amphibolite facies to greenschist facies.
A cross-polarized thin section image of a garnet-mica-schist from Salangen Municipality, Norway showing the strong strain fabric of schists. The black crystal is garnet, the pink-orange-yellow colored strands are muscovite mica, and the brown crystals are biotite mica. The grey and white crystals are quartz and (limited) feldspar.

Metamorphism is the transformation of existing rock (the protolith) to rock with a different mineral composition or texture. Metamorphism takes place at temperatures in excess of 150 °C (300 °F), and often also at elevated pressure or in the presence of chemically active fluids, but the rock remains mostly solid during the transformation.[1] Metamorphism is distinct from weathering or diagenesis, which are changes that take place at or just beneath Earth's surface.[2]

Various forms of metamorphism exist, including regional, contact, hydrothermal, shock, and dynamic metamorphism. These differ in the characteristic temperatures, pressures, and rate at which they take place and in the extent to which reactive fluids are involved. Metamorphism occurring at increasing pressure and temperature conditions is known as prograde metamorphism, while decreasing temperature and pressure characterize retrograde metamorphism.

Metamorphic petrology is the study of metamorphism. Metamorphic petrologists rely heavily on statistical mechanics and experimental petrology to understand metamorphic processes.

Metamorphic processes

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(Left) Randomly-orientated grains in a rock before metamorphism. (Right) Grains align orthogonal to the applied stress if a rock is subjected to stress during metamorphism

Metamorphism is the set of processes by which existing rock is transformed physically or chemically at elevated temperature, without actually melting to any great degree. The importance of heating in the formation of metamorphic rock was first recognized by the pioneering Scottish naturalist, James Hutton, who is often described as the father of modern geology. Hutton wrote in 1795 that some rock beds of the Scottish Highlands had originally been sedimentary rock, but had been transformed by great heat.[3]

Hutton also speculated that pressure was important in metamorphism. This hypothesis was tested by his friend, James Hall, who sealed chalk into a makeshift pressure vessel constructed from a cannon barrel and heated it in an iron foundry furnace. Hall found that this produced a material strongly resembling marble, rather than the usual quicklime produced by heating of chalk in the open air. French geologists subsequently added metasomatism, the circulation of fluids through buried rock, to the list of processes that help bring about metamorphism. However, metamorphism can take place without metasomatism (isochemical metamorphism) or at depths of just a few hundred meters where pressures are relatively low (for example, in contact metamorphism).[3]

Rock can be transformed without melting because heat causes atomic bonds to break, freeing the atoms to move and form new bonds with other atoms. Pore fluid present between mineral grains is an important medium through which atoms are exchanged.[4] This permits recrystallization of existing minerals or crystallization of new minerals with different crystalline structures or chemical compositions (neocrystallization).[1] The transformation converts the minerals in the protolith into forms that are more stable (closer to chemical equilibrium) under the conditions of pressure and temperature at which metamorphism takes place.[5][6]

Metamorphism is generally regarded to begin at temperatures of 100 to 200 °C (212 to 392 °F). This excludes diagenetic changes due to compaction and lithification, which result in the formation of sedimentary rocks.[7] The upper boundary of metamorphic conditions lies at the solidus of the rock, which is the temperature at which the rock begins to melt. At this point, the process becomes an igneous process.[8] The solidus temperature depends on the composition of the rock, the pressure, and whether the rock is saturated with water. Typical solidus temperatures range from 650 °C (1,202 °F) for wet granite at a few hundred megapascals (MPa) of pressure[9] to about 1,080 °C (1,980 °F) for wet basalt at atmospheric pressure.[10] Migmatites are rocks formed at this upper limit, which contains pods and veins of material that has started to melt but has not fully segregated from the refractory residue.[11]

The metamorphic process can occur at almost any pressure, from near surface pressure (for contact metamorphism) to pressures in excess of 16 kbar (1600 MPa).[12]

Recrystallization

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Basalt hand sample showing fine texture
Amphibolite formed by metamorphism of basalt showing coarse texture

The change in the grain size and orientation in the rock during the process of metamorphism is called recrystallization. For instance, the small calcite crystals in the sedimentary rocks limestone and chalk change into larger crystals in the metamorphic rock marble.[13] In metamorphosed sandstone, recrystallization of the original quartz sand grains results in very compact quartzite, also known as metaquartzite, in which the often larger quartz crystals are interlocked.[14] 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 points of contact (pressure solution) and redeposition in pore space.[15]

During recrystallization, the identity of the mineral does not change, only its texture. Recrystallization generally begins when temperatures reach above half the melting point of the mineral on the Kelvin scale.[16]

Pressure solution begins during diagenesis (the process of lithification of sediments into sedimentary rock) but is completed during early stages of metamorphism. For a sandstone protolith, the dividing line between diagenesis and metamorphism can be placed at the point where strained quartz grains begin to be replaced by new, unstrained, small quartz grains, producing a mortar texture that can be identified in thin sections under a polarizing microscope. With increasing grade of metamorphism, further recrystallization produces foam texture, characterized by polygonal grains meeting at triple junctions, and then porphyroblastic texture, characterized by coarse, irregular grains, including some larger grains (porphyroblasts.)[17]

A mylonite (through a petrographic microscope)

Metamorphic rocks are typically more coarsely crystalline than the protolith from which they formed. Atoms in the interior of a crystal are surrounded by a stable arrangement of neighboring atoms. This is partially missing at the surface of the crystal, producing a surface energy that makes the surface thermodynamically unstable. Recrystallization to coarser crystals reduces the surface area and so minimizes the surface energy.[18]

Although grain coarsening is a common result of metamorphism, rock that is intensely deformed may eliminate strain energy by recrystallizing as a fine-grained rock called mylonite. Certain kinds of rock, such as those rich in quartz, carbonate minerals, or olivine, are particularly prone to form mylonites, while feldspar and garnet are resistant to mylonitization.[19]

Phase change

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Phase diagram of Al2SiO5
(nesosilicates)

Phase change metamorphism is the creating of a new mineral with the same chemical formula as a mineral of the protolith. This involves a rearrangement of the atoms in the crystals. An example is provided by the aluminium silicate minerals, kyanite, andalusite, and sillimanite. All three have the identical composition, Al2SiO5. Kyanite is stable at surface conditions. However, at atmospheric pressure, kyanite transforms to andalusite at a temperature of about 190 °C (374 °F). Andalusite, in turn, transforms to sillimanite when the temperature reaches about 800 °C (1,470 °F). At pressures above about 4 kbar (400 MPa), kyanite transforms directly to sillimanite as the temperature increases.[20] A similar phase change is sometimes seen between calcite and aragonite, with calcite transforming to aragonite at elevated pressure and relatively low temperature.[21]

Neocrystallization

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Neocrystallization involves the creation of new mineral crystals different from the protolith. Chemical reactions digest the minerals of the protolith which yields new minerals. This is a very slow process as it can also involve the diffusion of atoms through solid crystals.[22]

An example of a neocrystallization reaction is the reaction of fayalite with plagioclase at elevated pressure and temperature to form garnet. The reaction is:[23]

Many complex high-temperature reactions may take place between minerals without them melting, and each mineral assemblage produced provides us with a clue as to the temperatures and pressures at the time of metamorphism. These reactions are possible because of rapid diffusion of atoms at elevated temperature. Pore fluid between mineral grains can be an important medium through which atoms are exchanged.[4]

A particularly important group of neocrystallization reactions are those that release volatiles such as water and carbon dioxide. During metamorphism of basalt to eclogite in subduction zones, hydrous minerals break down, producing copious quantities of water.[24] The water rises into the overlying mantle, where it lowers the melting temperature of the mantle rock, generating magma via flux melting.[25] The mantle-derived magmas can ultimately reach the Earth's surface, resulting in volcanic eruptions. The resulting arc volcanoes tend to produce dangerous eruptions, because their high water content makes them extremely explosive.[26]

Examples of dehydration reactions that release water include:[27]

An example of a decarbonation reaction is:[28]

Plastic deformation

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In plastic deformation pressure is applied to the protolith, which causes it to shear or bend, but not break. In order for this to happen temperatures must be high enough that brittle fractures do not occur, but not so high that diffusion of crystals takes place.[22] As with pressure solution, the early stages of plastic deformation begin during diagenesis.[29]

Types

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Regional

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Regional metamorphism is a general term for metamorphism that affects entire regions of the Earth's crust.[30] It most often refers to dynamothermal metamorphism, which takes place in orogenic belts (regions where mountain building is taking place),[31] but also includes burial metamorphism, which results simply from rock being buried to great depths below the Earth's surface in a subsiding basin.[32][33]

Dynamothermal

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A metamorphic rock, deformed during the Variscan orogeny, at Vall de Cardós, Lérida, Spain

To many geologists, regional metamorphism is practically synonymous with dynamothermal metamorphism.[30] This form of metamorphism takes place at convergent plate boundaries, where two continental plates or a continental plate and an island arc collide. The collision zone becomes a belt of mountain formation called an orogeny. The orogenic belt is characterized by thickening of the Earth's crust, during which the deeply buried crustal rock is subjected to high temperatures and pressures and is intensely deformed.[33][34] Subsequent erosion of the mountains exposes the roots of the orogenic belt as extensive outcrops of metamorphic rock,[35] characteristic of mountain chains.[33]

Metamorphic rock formed in these settings tends to shown well-developed foliation.[33] Foliation develops when a rock is being shortened along one axis during metamorphism. This causes crystals of platy minerals, such as mica and chlorite, to become rotated such that their short axes are parallel to the direction of shortening. This results in a banded, or foliated, rock, with the bands showing the colors of the minerals that formed them. Foliated rock often develops planes of cleavage. Slate is an example of a foliated metamorphic rock, originating from shale, and it typically shows well-developed cleavage that allows slate to be split into thin plates.[36]

The type of foliation that develops depends on the metamorphic grade. For instance, starting with a mudstone, the following sequence develops with increasing temperature: The mudstone is first converted to slate, which is a very fine-grained, foliated metamorphic rock, characteristic of very low grade metamorphism. Slate in turn is converted to phyllite, which is fine-grained and found in areas of low grade metamorphism. Schist is medium to coarse-grained and found in areas of medium grade metamorphism. High-grade metamorphism transforms the rock to gneiss, which is coarse to very coarse-grained.[37]

Rocks that were subjected to uniform pressure from all sides, or those that lack minerals with distinctive growth habits, will not be foliated. Marble lacks platy minerals and is generally not foliated, which allows its use as a material for sculpture and architecture.

Collisional orogenies are preceded by subduction of oceanic crust.[38] The conditions within the subducting slab as it plunges toward the mantle in a subduction zone produce their own distinctive regional metamorphic effects, characterized by paired metamorphic belts.[39]

The pioneering work of George Barrow on regional metamorphism in the Scottish Highlands showed that some regional metamorphism produces well-defined, mappable zones of increasing metamorphic grade. This Barrovian metamorphism is the most recognized metamorphic series in the world. However, Barrovian metamorphism is specific to pelitic rock, formed from mudstone or siltstone, and it is not unique even in pelitic rock. A different sequence in the northeast of Scotland defines Buchan metamorphism, which took place at lower pressure than the Barrovian.[40]

Burial

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Sioux Quartzite, a product of burial metamorphism

Burial metamorphism takes place simply through rock being buried to great depths below the Earth's surface in a subsiding basin.[33] Here the rock is subjected to high temperatures and the great pressure caused by the immense weight of the rock layers above. Burial metamorphism tends to produce low-grade metamorphic rock. This shows none of the effects of deformation and folding so characteristic of dynamothermal metamorphism.[41]

Examples of metamorphic rocks formed by burial metamorphism include some of the rocks of the Midcontinent Rift System of North America, such as the Sioux Quartzite,[42] and in the Hamersley Basin of Australia.[43]

Contact

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A metamorphic aureole in the Henry Mountains, Utah. The greyish rock on top is the igneous intrusion, consisting of porphyritic granodiorite from the Henry Mountains laccolith, and the pinkish rock on the bottom is the sedimentary country rock, a siltstone. In between, the metamorphosed siltstone is visible as both the dark layer (~5  cm thick) and the pale layer below it.

Contact metamorphism occurs typically around intrusive igneous rocks as a result of the temperature increase caused by the intrusion of magma into cooler country rock. The area surrounding the intrusion where the contact metamorphism effects are present is called the metamorphic aureole,[44] the contact aureole, or simply the aureole.[45] Contact metamorphic rocks are usually known as hornfels. Rocks formed by contact metamorphism may not present signs of strong deformation and are often fine-grained[46][47] and extremely tough.[48] The Yule Marble used on the Lincoln Memorial exterior and the Tomb of the Unknown Soldier in Arlington National Cemetery was formed by contact metamorphism.[49]

Contact metamorphism is greater adjacent to the intrusion and dissipates with distance from the contact.[50] The size of the aureole depends on the heat of the intrusion, its size, and the temperature difference with the wall rocks. Dikes generally have small aureoles with minimal metamorphism, extending not more than one or two dike thicknesses into the surrounding rock,[51] whereas the aureoles around batholiths can be up to several kilometers wide.[52][53]

The metamorphic grade of an aureole is measured by the peak metamorphic mineral which forms in the aureole. This is usually related to the metamorphic temperatures of pelitic or aluminosilicate rocks and the minerals they form. The metamorphic grades of aureoles at shallow depth are albite-epidote hornfels, hornblende hornfels, pyroxene hornfels, and sillimanite hornfels, in increasing order of temperature of formation. However, the albite-epidote hornfels is often not formed, even though it is the lowest temperature grade.[54]

Magmatic fluids coming from the intrusive rock may also take part in the metamorphic reactions. An extensive addition of magmatic fluids can significantly modify the chemistry of the affected rocks. In this case the metamorphism grades into metasomatism. If the intruded rock is rich in carbonate the result is a skarn.[55] Fluorine-rich magmatic waters which leave a cooling granite may often form greisens within and adjacent to the contact of the granite.[56] Metasomatic altered aureoles can localize the deposition of metallic ore minerals and thus are of economic interest.[57][58]

Fenitization, or Na-metasomatism, is a distinctive form of contact metamorphism accompanied by metasomatism. It takes place around intrusions of a rare type of magma called a carbonatite that is highly enriched in carbonates and low in silica. Cooling bodies of carbonatite magma give off highly alkaline fluids rich in sodium as they solidify, and the hot, reactive fluid replaces much of the mineral content in the aureole with sodium-rich minerals.[59]

A special type of contact metamorphism, associated with fossil fuel fires, is known as pyrometamorphism.[60][61]

Hydrothermal

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Hydrothermal metamorphism is the result of the interaction of a rock with a high-temperature fluid of variable composition. The difference in composition between an existing rock and the invading fluid triggers a set of metamorphic and metasomatic reactions. The hydrothermal fluid may be magmatic (originate in an intruding magma), circulating groundwater, or ocean water.[33] Convective circulation of hydrothermal fluids in the ocean floor basalts produces extensive hydrothermal metamorphism adjacent to spreading centers and other submarine volcanic areas. The fluids eventually escape through vents on the ocean floor known as black smokers.[62] The patterns of this hydrothermal alteration are used as a guide in the search for deposits of valuable metal ores.[63]

Shock

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Main article: Shock metamorphism

Shock metamorphism occurs when an extraterrestrial object (a meteorite for instance) collides with the Earth's surface. Impact metamorphism is, therefore, characterized by ultrahigh pressure conditions and low temperature. The resulting minerals (such as SiO2 polymorphs coesite and stishovite) and textures are characteristic of these conditions.[64]

Dynamic

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Dynamic metamorphism is associated with zones of high strain such as fault zones.[33] In these environments, mechanical deformation is more important than chemical reactions in transforming the rock. The minerals present in the rock often do not reflect conditions of chemical equilibrium, and the textures produced by dynamic metamorphism are more significant than the mineral makeup.[65]

There are three deformation mechanisms by which rock is mechanically deformed. These are cataclasis, the deformation of rock via the fracture and rotation of mineral grains;[66] plastic deformation of individual mineral crystals; and movement of individual atoms by diffusive processes.[67] The textures of dynamic metamorphic zones are dependent on the depth at which they were formed, as the temperature and confining pressure determine the deformation mechanisms which predominate.[68]

At the shallowest depths, a fault zone will be filled with various kinds of unconsolidated cataclastic rock, such as fault gouge or fault breccia. At greater depths, these are replaced by consolidated cataclastic rock, such as crush breccia, in which the larger rock fragments are cemented together by calcite or quartz. At depths greater than about 5 kilometers (3.1 mi), cataclasites appear; these are quite hard rocks consist of crushed rock fragments in a flinty matrix, which forms only at elevated temperature. At still greater depths, where temperatures exceed 300 °C (572 °F), plastic deformation takes over, and the fault zone is composed of mylonite. Mylonite is distinguished by its strong foliation, which is absent in most cataclastic rock.[69] It is distinguished from the surrounding rock by its finer grain size.[70]

There is considerable evidence that cataclasites form as much through plastic deformation and recrystallization as brittle fracture of grains, and that the rock may never fully lose cohesion during the process. Different minerals become ductile at different temperatures, with quartz being among the first to become ductile, and sheared rock composed of different minerals may simultaneously show both plastic deformation and brittle fracture.[71]

The strain rate also affects the way in which rocks deform. Ductile deformation is more likely at low strain rates (less than 10−14 sec−1) in the middle and lower crust, but high strain rates can cause brittle deformation. At the highest strain rates, the rock may be so strongly heated that it briefly melts, forming a glassy rock called pseudotachylite.[72][73] Pseudotachylites seem to be restricted to dry rock, such as granulite.[74]

Classification of metamorphic rocks

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Main article: Metamorphic rock § Classification

Metamorphic rocks are classified by their protolith, if this can be determined from the properties of the rock itself. For example, if examination of a metamorphic rock shows that its protolith was basalt, it will be described as a metabasalt. When the protolith cannot be determined, the rock is classified by its mineral composition or its degree of foliation.[75][76][77]

Metamorphic grades

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Metamorphic grade is an informal indication of the amount or degree of metamorphism.[78]

In the Barrovian sequence (described by George Barrow in zones of progressive metamorphism in Scotland), metamorphic grades are also classified by mineral assemblage based on the appearance of key minerals in rocks of pelitic (shaly, aluminous) origin:

Low grade ------------------- Intermediate --------------------- High grade

Greenschist ------------- Amphibolite ----------------------- Granulite
Slate --- Phyllite ---------- Schist ---------------------- Gneiss --- Migmatite
Chlorite zone
Biotite zone
Garnet zone
Staurolite zone
Kyanite zone
Sillimanite zone

A more complete indication of this intensity or degree is provided by the concept of metamorphic facies.[78]

Metamorphic facies

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Main article: Metamorphic facies

Metamorphic facies are recognizable terranes or zones with an assemblage of key minerals that were in equilibrium under specific range of temperature and pressure during a metamorphic event. The facies are named after the metamorphic rock formed under those facies conditions from basalt.[79]

The particular mineral assemblage is somewhat dependent on the composition of that protolith, so that (for example) the amphibolite facies of a marble will not be identical with the amphibolite facies of a pellite. However, the facies are defined such that metamorphic rock with as broad a range of compositions as is practical can be assigned to a particular facies. The present definition of metamorphic facies is largely based on the work of the Finnish geologist, Pentti Eskola in 1921, with refinements based on subsequent experimental work. Eskola drew upon the zonal schemes, based on index minerals, that were pioneered by the British geologist, George Barrow.[12]

The metamorphic facies is not usually considered when classifying metamorphic rock based on protolith, mineral mode, or texture. However, a few metamorphic facies produce rock of such distinctive character that the facies name is used for the rock when more precise classification is not possible. The chief examples are amphibolite and eclogite. The British Geological Survey strongly discourages use of granulite as a classification for rock metamorphosed to the granulite facies. Instead, such rock will often be classified as a granofels.[76] However, this is not universally accepted.[77]

Temperatures and pressures of metamorphic facies
Temperature Pressure Facies
Low Low Zeolite
Lower Moderate Lower Moderate Prehnite-Pumpellyite
Moderate to High Low Hornfels
Low to Moderate Moderate to High Blueschist
Moderate → High Moderate Greenschist→Amphibolite→Granulite
Moderate to High High Eclogite

See diagram for more detail.

Prograde and retrograde

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Metamorphism is further divided into prograde and retrograde metamorphism. Prograde metamorphism involves the change of mineral assemblages (paragenesis) with increasing temperature and (usually) pressure conditions. These are solid state dehydration reactions, and involve the loss of volatiles such as water or carbon dioxide. Prograde metamorphism results in rock characteristic of the maximum pressure and temperature experienced. Metamorphic rocks usually do not undergo further change when they are brought back to the surface.[80]

Retrograde metamorphism involves the reconstitution of a rock via revolatisation under decreasing temperatures (and usually pressures), allowing the mineral assemblages formed in prograde metamorphism to revert to those more stable at less extreme conditions. This is a relatively uncommon process, because volatiles produced during prograde metamorphism usually migrate out of the rock and are not available to recombine with the rock during cooling. Localized retrograde metamorphism can take place when fractures in the rock provide a pathway for groundwater to enter the cooling rock.[80]

Equilibrium mineral assemblages

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Petrogenetic grid showing aluminium silicate-muscovite-quartz-K feldspar phase boundaries
ACF compatibility diagrams (aluminium-calcium-iron) showing phase equilibria in metamorphic mafic rocks at different P-T circumstances (metamorphic facies). Dots represent mineral phases, thin grey lines are equilibria between two phases. Mineral abbreviations: act = actinolite; cc = calcite; chl = chlorite; di = diopside; ep = epidote; glau = glaucophane; gt = garnet; hbl = hornblende; ky = kyanite; law = lawsonite; plag = plagioclase; om = omphacite; opx = orthopyroxene; zo = zoisite

Metamorphic processes act to bring the protolith closer to thermodynamic equilibrium, which is its state of maximum stability. For example, shear stress (nonhydrodynamic stress) is incompatible with thermodynamic equilibrium, so sheared rock will tend to deform in ways that relieve the shear stress.[81] The most stable assemblage of minerals for a rock of a given composition is that which minimizes the Gibbs free energy[82]

where:

In other words, a metamorphic reaction will take place only if it lowers the total Gibbs free energy of the protolith. Recrystallization to coarser crystals lowers the Gibbs free energy by reducing surface energy,[18] while phase changes and neocrystallization reduce the bulk Gibbs free energy. A reaction will begin at the temperature and pressure where the Gibbs free energy of the reagents becomes greater than that of the products.[83]

A mineral phase will generally be more stable if it has a lower internal energy, reflecting tighter binding between its atoms. Phases with a higher density (expressed as a lower molar volume V) are more stable at higher pressure, while minerals with a less ordered structure (expressed as a higher entropy S) are favored at high temperature. Thus andalusite is stable only at low pressure, since it has the lowest density of any aluminium silicate polymorph, while sillimanite is the stable form at higher temperatures, since it has the least ordered structure.[84]

The Gibbs free energy of a particular mineral at a specified temperature and pressure can be expressed by various analytic formulas. These are calibrated against experimentally measured properties and phase boundaries of mineral assemblages. The equilibrium mineral assemblage for a given bulk composition of rock at a specified temperature and pressure can then be calculated on a computer.[85][86]

However, it is often very useful to represent equilibrium mineral assemblages using various kinds of diagrams.[87] These include petrogenetic grids[88][89] and compatibility diagrams (compositional phase diagrams.)[90][91]

Petrogenetic grids

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Main article: Petrogenetic grid

A petrogenetic grid is a geologic phase diagram that plots experimentally derived metamorphic reactions at their pressure and temperature conditions for a given rock composition. This allows metamorphic petrologists to determine the pressure and temperature conditions under which rocks metamorphose.[88][89] The Al2SiO5 nesosilicate phase diagram shown is a very simple petrogenetic grid for rocks that only have a composition consisting of aluminum (Al), silicon (Si), and oxygen (O). As the rock undergoes different temperatures and pressure, it could be any of the three given polymorphic minerals.[84] For a rock that contains multiple phases, the boundaries between many phase transformations may be plotted, though the petrogenetic grid quickly becomes complicated. For example, a petrogenetic grid might show both the aluminium silicate phase transitions and the transition from aluminum silicate plus potassium feldspar to muscovite plus quartz.[92]

Compatibility diagrams

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Main article: Compatibility diagram

Whereas a petrogenetic grid shows phases for a single composition over a range of temperature and pressure, a compatibility diagram shows how the mineral assemblage varies with composition at a fixed temperature and pressure. Compatibility diagrams provide an excellent way to analyze how variations in the rock's composition affect the mineral paragenesis that develops in a rock at particular pressure and temperature conditions.[90][91] Because of the difficulty of depicting more than three components (as a ternary diagram), usually only the three most important components are plotted, though occasionally a compatibility diagram for four components is plotted as a projected tetrahedron.[93]

See also

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Footnotes

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References

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Further reading

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[edit]
Revisions and contributorsEdit on WikipediaRead on Wikipedia

Metamorphism

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Metamorphism is the by which rocks are transformed in their solid state due to changes in , , and chemically active fluids, resulting in alterations to their composition, texture, and sometimes , without . This geological occurs primarily within the and is driven by tectonic forces, such as plate collisions or , which impose the necessary conditions on pre-existing rocks known as protoliths. The resulting metamorphic rocks provide key insights into the thermal and tectonic history of the planet, often exhibiting distinct features like —layered alignments of —or non-foliated textures depending on the conditions. The agents of metamorphism—heat, pressure, and fluids—work together or independently to recrystallize minerals and reorganize the rock's fabric. Heat, sourced from igneous intrusions or deep burial, promotes mineral growth and reactions, while pressure, including directed stress from tectonic deformation, can align minerals into parallel orientations. Chemically active fluids, often derived from dehydration of rocks or magmatic sources, facilitate ion exchange and metasomatism, which may alter the rock's bulk composition. These changes are typically prograde, progressing with increasing temperature and pressure, though retrograde metamorphism can occur during uplift and cooling. Metamorphism is classified into several types based on the dominant conditions and scale. Contact metamorphism occurs locally around igneous intrusions where heat from alters surrounding rocks, producing non-foliated minerals like those in or . Regional metamorphism, the most widespread type, affects large areas during orogenic events, leading to foliated rocks such as , , and through combined heat and pressure. Other types include dynamic (cataclastic) metamorphism from intense shearing along faults, hydrothermal metamorphism driven by hot fluids in oceanic or volcanic settings, burial metamorphism in sedimentary basins, and shock metamorphism from impacts. Common metamorphic rocks illustrate the diversity of outcomes: foliated varieties like phyllite (fine-grained with silky sheen) and gneiss (banded with separated and ) form under regional conditions, while non-foliated rocks such as quartzite (from ) and marble (from ) result from thermal or contact processes without strong directional stress. These rocks are integral to the rock cycle, often serving as protoliths for further metamorphism or sources of economic minerals like , , and . Understanding metamorphism is crucial for interpreting mountain-building processes, resource exploration, and the evolution of .

Fundamentals

Definition and Characteristics

Metamorphism is the process by which pre-existing rocks, known as protoliths, undergo solid-state changes in and texture due to variations in , , and fluid composition, without reaching the point of . This transformation occurs within the and results in the recrystallization of minerals, leading to a new rock type that retains much of the original chemical composition. The term "metamorphism" derives from the Greek words "meta," meaning change or after, and "morphe," meaning form, reflecting the alteration in rock structure and appearance. The concept was first systematically explored by Scottish geologist in his 1795 publication Theory of the Earth, where he examined metamorphic rocks in , such as schists intruded by veins, to support his ideas on geological cycles driven by internal . Key characteristics of metamorphism include textural modifications, such as the development of —planar alignment of minerals like due to directed pressure—and mineralogical reconstitution, where new minerals form through recrystallization without significant volume change. Typically, metamorphism is isochemical, preserving the bulk chemical composition of the as atoms rearrange into stable minerals under new conditions, though can occur in fluid-rich environments, introducing or removing chemical components via or influx. These changes distinguish metamorphism from , which involves low-temperature, low-pressure alterations in sedimentary rocks during burial, and from igneous processes, which require partial or complete melting to form new rocks. Representative examples of metamorphic rocks illustrate these traits: forms from through the recrystallization of , resulting in a coarse-grained, non-foliated texture suitable for , while develops from under higher-grade conditions, exhibiting pronounced from aligned platy minerals like .

Agents of Metamorphism

Metamorphism is driven by three primary agents: , , and chemically active fluids, which act individually or in combination to alter the and texture of pre-existing rocks without them. These agents provide the energy and conditions necessary for solid-state transformations, with their effects depending on the geological setting. Typical metamorphic conditions involve temperatures ranging from 200°C to 800°C and pressures from 0.1 to 10 kbar (0.01 to 1 GPa), though extremes can exceed these values in specific environments. Heat, or thermal energy, is a fundamental agent that accelerates atomic and reaction kinetics by increasing molecular vibrations, thereby enabling bond breaking and in structures. Sources of include geothermal gradients, where rises with depth at approximately 25–30°C per kilometer, and localized inputs from intrusions that can rapidly elevate temperatures in surrounding rocks. This thermal influence promotes processes like annealing, which enlarges grain sizes and reduces , particularly at temperatures above 200°C where reaction rates become geologically significant. Pressure encompasses both lithostatic (confining) and differential (directed) components, each exerting distinct influences on rock behavior. Lithostatic pressure arises from the of overlying material and is isotropic, increasing linearly with depth at about 0.3 kbar per kilometer; it primarily affects rock volume by compressing pore spaces and facilitating denser assemblages, typically measured in kilobars (kbar) or gigapascals (GPa). In contrast, differential pressure involves unequal stresses from tectonic forces, such as plate convergence, which deform rocks and align minerals, often at pressures exceeding 1 kbar in active orogenic belts. Chemically active fluids, predominantly water-rich with dissolved ions like Na⁺, K⁺, and CO₂, play a crucial role by lowering activation energies for reactions and enhancing diffusion through rock matrices. These fluids, often derived from dehydration of subducting slabs or magmatic exsolution, facilitate —a process where elements are added or removed, altering bulk rock composition and forming new minerals such as micas or garnets. In water-saturated systems, fluid pressure can reduce , promoting brittle-ductile transitions, and their presence is essential for many metamorphic reactions that would otherwise proceed too slowly. The interactions among these agents determine the dominant metamorphic style: prevails in contact metamorphism near igneous bodies, where temperatures can reach 800°C with minimal ; , especially differential stress, dominates regional metamorphism in deeply buried terrains under 2–10 kbar; and are key in hydrothermal settings, often combining with to drive metasomatic changes. Strain from differential can further amplify effects by localizing fluid infiltration along shear zones, illustrating the synergistic nature of these agents in natural systems.

Metamorphic Processes

Recrystallization

Recrystallization is a fundamental metamorphic process involving the reorganization of atoms within existing mineral grains to form new, strain-free crystals, thereby refining the rock's texture without the formation of entirely new mineral phases. This atomic-scale rearrangement occurs through diffusion and dislocation movement, primarily driven by the minimization of stored strain energy accumulated from prior deformation. Key mechanisms include recovery, where dislocations reorganize into lower-energy configurations such as subgrain boundaries; subgrain rotation, in which progressive misorientation of subgrains transforms low-angle boundaries into high-angle grain boundaries; and grain boundary migration, where boundaries advance to consume deformed regions, often resulting in lobate shapes. These processes enable the rock to achieve a more stable microstructure under sustained metamorphic conditions. Recrystallization manifests in two primary types: static and dynamic. Static recrystallization takes place after deformation ceases, during annealing under elevated temperatures without ongoing stress, allowing grains to grow and equilibrate. In contrast, dynamic recrystallization occurs simultaneously with deformation, where new grains nucleate and grow amid active , often leading to finer-grained textures that influence ongoing . For instance, in phyllosilicates under greenschist-facies conditions, static recrystallization enhances preferred orientations through dissolution and regrowth of grains aligned with cleavage planes. The effects of recrystallization include the development of larger, more equidimensional grains that replace irregular, strained ones, promoting textural maturity and reducing internal defects. This grain coarsening also diminishes by sealing intergranular voids through boundary migration and diffusive , enhancing rock cohesion. In quartzites, derived from quartz-rich protoliths, recrystallization produces , polygonal grains that contribute to the rock's and low permeability, as observed in regionally metamorphosed terrains where original sedimentary fabrics are obliterated. Recrystallization is favored at temperatures exceeding 300°C, where rates become sufficient to enable mobility and boundary movement, though it remains a time-dependent requiring prolonged exposure to metamorphic conditions. Below this threshold, such as in low-grade settings, recrystallization is sluggish due to limited thermal activation. Microstructural evidence for recrystallization is commonly revealed through electron backscatter diffraction (EBSD) and (TEM), showing equiangular triple junctions at approximately 120° angles between grains, indicative of energy minimization at equilibrium boundaries. These features are particularly evident in dynamically recrystallized , where subgrain networks evolve into polygonal mosaics.

Phase Transformations

Phase transformations during metamorphism involve solid-state changes in structure or composition driven by thermodynamic instability under altered pressure and temperature conditions. These transformations primarily manifest as polymorphic transitions, where a mineral adopts a new while retaining its , or as adjustments in solid solutions, where ions redistribute within the lattice to achieve stability. For instance, (α-SiO₂) transforms to the denser under high pressures exceeding 2-3 GPa, a change characteristic of shock metamorphism from impacts. Similarly, solid-solution adjustments occur in minerals like , where compositional zoning evolves to minimize free energy in response to changing conditions. Thermodynamically, these transformations proceed toward the minimization of (G = H - TS), where the stable phase at given P-T conditions has the lowest G. The boundaries between phases are defined by equilibrium curves on P-T diagrams, with slopes determined by the Clapeyron equation: dP/dT = ΔS/ΔV, where ΔS is the entropy change and ΔV is the volume change across the transition. For reconstructive polymorphs involving bond breaking, ΔV is often negative (denser high-P phase), yielding positive slopes, as seen in the quartz-coesite boundary. In the Al₂SiO₅ system, the at approximately 0.5 GPa and 500°C separates fields for (low-pressure, low-temperature), (high-pressure), and (high-temperature), allowing pelitic rocks to record specific metamorphic paths through these polymorphs. Another key example is the calcite-aragonite transition in carbonates, where , the orthorhombic high-pressure polymorph, stabilizes above ~0.3 GPa at surface temperatures, though it rarely persists due to kinetic factors. Kinetically, phase transformations are governed by and growth processes, which often impose barriers leading to metastable persistence of parent phases. requires overcoming an related to interfacial free energy and strain, with rates increasing exponentially with overstepping of equilibrium conditions (ΔG driving force). Growth follows via interface advance, but -limited mechanisms can slow the process, as in the aragonite-to-calcite reversion, where experiments show transformation times exceeding millions of years at 300-500°C due to high barriers. Fluids may briefly accelerate these kinetics by lowering activation energies through enhanced , though their role is secondary to P-T drivers. Metastable assemblages, such as low-pressure polymorphs in high-grade terrains, thus commonly survive beyond equilibrium boundaries. Diagnostic features of phase transformations include pseudomorphs, where the new phase replaces the original while preserving its external shape and sometimes internal fabric due to epitaxial growth or topotactic relations. For example, pseudomorphs after in deformed pelites retain prismatic outlines, indicating the transformation path. Such textures provide evidence of the sequence and conditions of metamorphic evolution without complete textural reset.

Neocrystallization

Neocrystallization refers to the formation of entirely new mineral species during metamorphism through chemical reactions that involve the breakdown of pre-existing minerals and the recombination of their constituent atoms into novel structures. This process contrasts with mere textural adjustments, as it requires of ions across grain boundaries or through fluids to enable the synthesis of minerals with compositions not present in the . Such reactions are driven by changes in and , often facilitated by the presence of aqueous fluids that lower energies for atomic mobility. A key aspect of neocrystallization involves devolatilization reactions, where volatile components such as H₂O or CO₂ are released, promoting the stability of or less hydrous phases at higher metamorphic grades. These reactions typically proceed in a sequence that reflects progressive metamorphic conditions, with the expelled volatiles potentially influencing in the surrounding rock. Neocrystallization can occur in two primary modes: isochemical, within a where the bulk composition remains unchanged except for volatile loss, and metasomatic, in an open system where external fluids introduce or remove elements, leading to significant chemical alterations. The distinction hinges on fluid-rock interactions, with often linked to advective transport of solutes. Representative examples illustrate neocrystallization across metamorphic facies. In the facies, commonly forms from the reaction of clay minerals like or with iron-magnesium-bearing phases, yielding a green phyllosilicate that defines the facies assemblage. At higher grades in the amphibolite facies, nucleates and grows through the breakdown of hydrous minerals such as or , incorporating calcium, aluminum, and iron to form almandine-rich porphyroblasts. A classic reaction exemplifying this process is the dehydration of in the presence of : Muscovite+Quartz→Sillimanite+K-feldspar+H2O\text{Muscovite} + \text{Quartz} \rightarrow \text{Sillimanite} + \text{K-feldspar} + \text{H}_2\text{O} This net-transfer reaction occurs around 600–700°C and 3–5 kbar, marking the transition to higher-grade pelitic assemblages. Evidence for neocrystallization is preserved in microstructural features, such as zoned crystals that record progressive compositional changes during growth, with cores reflecting earlier, lower-grade conditions and rims indicating later, higher-grade overgrowths. Reaction rims—narrow zones of new minerals forming at interfaces between reactants—further attest to localized diffusion and incomplete equilibrium, often surrounding relict grains of the original assemblage. These textures provide direct petrologic indicators of the reaction pathways and fluid involvement in neocrystallization.

Deformation Mechanisms

Deformation mechanisms during metamorphism describe the ways in which differential stress alters the internal and fabric of rocks, primarily through mechanical processes that accommodate strain without significant volume change. These mechanisms transition from brittle to ductile behaviors as , , and vary, influencing the development of aligned orientations and shear-related textures in metamorphic rocks. Brittle deformation predominates at low temperatures (typically below 350°C) and shallow crustal depths, where rocks under stress, while ductile deformation becomes feasible at higher temperatures (above 300–450°C, depending on ), allowing continuous flow. Key brittle mechanisms include cataclasis, involving the grinding and fragmentation of grains along fault planes, which reduces grain size but produces angular fragments and fault gouge. Pressure solution, a semi-brittle active at low to moderate temperatures, entails dissolution of minerals at points of high (e.g., grain contacts) and reprecipitation in low-stress regions, leading to and the formation of or sutured boundaries. In ductile regimes, dislocation creep dominates, where crystal defects () move via glide and climb, enabling plastic deformation; this mechanism requires elevated temperatures to activate , with deforming ductily around 300°C and around 450°C. These mechanisms generate distinctive fabrics that record history. , a planar fabric, arises from the preferred alignment of platy or elongate minerals (e.g., micas, amphiboles) perpendicular to the maximum principal stress, often through mechanical or pressure solution. Lineation, a linear fabric, develops from the elongation of minerals or alignment of fold axes during non-coaxial shear. In low-grade settings, slaty cleavage forms as a pervasive in fine-grained rocks like , resulting from the and alignment of phyllosilicates under compressional stress. Rock during deformation follows constitutive flow laws that quantify dependence on stress and . For dislocation creep, the power-law creep equation is commonly applied: ϵ˙=Aσnexp⁡(−[Q](/page/Q)RT)\dot{\epsilon} = A \sigma^n \exp\left(-\frac{[Q](/page/Q)}{RT}\right) where ϵ˙\dot{\epsilon} is the , σ\sigma is the differential stress, AA is a material-specific constant, nn is the stress exponent (typically 3–5 for non-linear viscous flow), QQ is the for deformation, RR is the , and TT is absolute ; this relation underscores how increasing exponentially accelerates ductile flow. Prominent examples include mylonites, which form in ductile shear zones through progressive grain-size reduction via cataclasis transitioning to dislocation creep and dynamic recrystallization, resulting in fine-grained matrices (50–90% in mylonites) with strong foliation and shear bands (S-C fabrics) that indicate sense of shear. Boudinage structures exemplify rheological layering, where stiffer (competent) layers in a softer matrix fracture and separate into sausage-like segments during extension, with symmetric types (e.g., drawn boudins) forming under pure shear and asymmetric types (e.g., domino or gash boudins) under simple shear. These processes span scales, from microscale features like kink bands in individual mica grains to macroscale folds and regional shear zones. Syn-deformational recrystallization often accompanies ductile mechanisms, reducing stored strain energy without net volume loss.

Types of Metamorphism

Regional Metamorphism

Regional metamorphism occurs over vast areas, typically spanning hundreds to thousands of kilometers, and is primarily associated with tectonic processes at convergent plate boundaries where is thickened through collision or . This type of metamorphism affects large volumes of rock in orogenic belts, resulting from deep burial, elevated temperatures, and directed pressures that drive mineralogical and textural changes in the . The process unfolds over durations of 10 to 100 million years, allowing for the gradual equilibration of mineral assemblages under evolving conditions. A key subtype is dynamothermal regional metamorphism, which combines thermal effects from radiogenic and tectonic heat sources with high pressures and intense deformation, leading to the development of and other fabrics. Characteristic of this subtype is progressive zoning, where metamorphic grade increases systematically away from the structural core of the orogen, delineated by isograds—lines of constant mineral appearance or disappearance. These Barrovian zones, named after George Barrow's observations in the , reflect a continuum from low-grade conditions near the margins to high-grade or in deeper levels, with index s such as marking mid-grade transitions. Prominent examples include the Appalachian orogenic belt in eastern , where collision produced widespread Barrovian metamorphism with staurolite-bearing schists in zones of intermediate grade, and the Himalayan belt, resulting from ongoing India-Asia convergence since approximately 50 million years ago. In these settings, pressure- (P-T) paths typically form clockwise loops, beginning with rapid that increases both and , reaching peak conditions at depth, followed by tectonic uplift and cooling that decrease more abruptly than . These paths underscore the role of in initiating metamorphism, with deformation fabrics like schistosity developing concurrently to accommodate strain during orogenesis.

Contact Metamorphism

Contact metamorphism is a localized thermal process that affects rocks adjacent to igneous intrusions, such as plutons, dikes, or sills, primarily through from the cooling without involving significant tectonic deformation. This metamorphism develops in the shallow crust, forming a contact aureole—a zone of altered —that typically ranges from 10 meters to several kilometers in width, depending on the size of the intrusion and the presence of volatiles. Key characteristics include steep thermal gradients, with temperatures decreasing rapidly away from the intrusion, resulting in fine-grained, non-foliated rocks that exhibit granoblastic textures due to static recrystallization. , a common product, forms from various protoliths like or and displays equidimensional grains without preferred orientation. The pressure-temperature conditions are marked by high temperatures ranging from 500°C to 800°C and low pressures below 2 kbar, reflecting the shallow depths and dominance of effects over pressure. Static recrystallization predominates, allowing minerals to grow and equilibrate under these conditions without shearing. Aureoles display concentric zonation, with inner zones experiencing high-grade metamorphism closest to the intrusion and progressively lower-grade zones outward, reflecting the diminishing heat influence. Representative examples include formation at the contacts between rocks and intrusions, where calc-silicate develop, and cordierite-bearing assemblages in pelitic rocks near granitic plutons. Magmatic fluids can enhance alteration in some settings by facilitating mineral reactions.

Hydrothermal Metamorphism

Hydrothermal metamorphism involves the interaction of hot, chemically reactive fluids with rocks, leading to mineralogical and chemical changes primarily through , where elements are added or removed from the rock system. This process typically occurs in localized settings such as near faults, volcanic systems, or ocean floors, where fluids like or magmatic waters circulate through fractured rocks at temperatures often exceeding 150°C. In oceanic environments, particularly along mid-ocean ridges, heated percolates through and , driving alteration under relatively low pressures. On continents, magmatic fluids associated with intrusions or fault zones facilitate similar reactions in volcanic or plutonic rocks. The primary mechanisms include dissolution-precipitation, where minerals dissolve in the fluid and new phases precipitate, often generating that enhances further fluid flow, and , which replaces ions in the crystal lattice without significant volume change. A key example is serpentinization of , an exothermic where and react with water to form serpentine minerals like or , releasing and fixing magnesium into the rock. This process commonly occurs in the at mid-ocean ridges, altering ultramafic rocks to hydrous assemblages such as , , and . Hydrothermal metamorphism manifests in distinct types based on tectonic setting. In ocean-ridge systems, it produces greenschist-facies assemblages, including , , and , through interaction of with at temperatures of 250–350°C. In contrast, continental settings often result in propylitic alteration, characterized by the formation of , , and in igneous rocks at lower temperatures below 250°C, typically peripheral to more intense alteration zones. Chemical changes are pronounced, with fluids introducing and magnesium while facilitating sodium and . For instance, magnesium during greenschist-facies alteration of oceanic basalts can fix significant Mg from , altering the bulk composition of a 1-km-thick crustal section. In ultramafic rocks, serpentinization adds Si and Mg while mobilizing Fe. Representative examples include epidosites in complexes, which form through intense hydrothermal alteration of metabasalts in upflow zones of seafloor systems, resulting in epidote-quartz assemblages that release metals like and into fluids. Listvenites, another hallmark, arise from silicification and of serpentinized via CO₂-rich hydrothermal fluids, producing quartz-fuchsite-carbonate rocks that indicate high time-integrated fluid flux under conditions around 290–340°C.

Shock and Dynamic Metamorphism

Shock metamorphism results from the hypervelocity impacts of meteorites or other extraterrestrial bodies, which generate extreme pressures exceeding 10 GPa and temperatures up to several thousand degrees over durations of microseconds to seconds. These conditions induce irreversible changes in rocks without significant chemical alteration, primarily through mechanical deformation and localized melting. A hallmark feature is shatter cones, which are striated, conical fracture surfaces that form at shock pressures starting from about 2 GPa, though they are most prominent at higher intensities where they may contain embedded microdeformation structures. Another diagnostic indicator is planar deformation features (PDFs) in grains, consisting of closely spaced, parallel lamellae of amorphous silica or transformed phases, which develop at pressures between approximately 8 and 35 GPa. These features arise from propagation that causes selective gliding along crystallographic planes, producing a planar fabric orientation. The in exemplifies shock metamorphism, with abundant shatter cones and PDFs preserved in granitic rocks, confirming pressures well above 10 GPa during the ~2 billion-year-old event. Dynamic metamorphism, in contrast, occurs within active fault zones during rapid seismic slip, where intense shearing at ultra-high strain rates of 10310^3 to 10610^6 s−1^{-1} generates frictional heating sufficient to cause localized melting. This process, often termed cataclasis under dynamic conditions, produces pseudotachylite—fine-grained, glassy veins or injections that represent quenched frictional melts, typically millimeters to centimeters thick and parallel to the fault plane. The melt forms adiabatically due to the high slip velocities (up to several meters per second) and confined shear zones, where heat dissipation is minimal during the brief rupture duration. Planar fabrics, including foliated cataclasites and vein injections, reflect the extreme directional strain, distinguishing these rocks from broader deformation features. Along the San Andreas Fault in California, pseudotachylite veins have been observed in drill cores, linked to frictional heating during large earthquakes, though some occurrences may involve comminution rather than full melting. Both shock and dynamic metamorphism are characterized by their rapidity, with strain rates orders of magnitude higher than in tectonic settings, leading to non-equilibrium products like amorphous phases and shock twins rather than recrystallized equilibria. The heating mechanism is predominantly adiabatic in these environments, as the short timescales prevent significant conductive loss, unlike the slower, diffusion-dominated transfer in other metamorphic types. This results in preserved shock indicators that provide direct evidence of high-energy events, essential for identifying ancient impacts or seismic histories in the geological record.

Burial Metamorphism

Burial metamorphism encompasses the progressive alteration of sedimentary rocks under conditions of increasing and due to in thick sedimentary sequences, without significant tectonic deformation or igneous influence. It typically develops in stable tectonic settings such as passive continental margins or foreland basins, where sediments accumulate to depths of 5–15 km, leading to low-grade metamorphic changes that extend diagenetic processes. The process occurs within the anchizone to low epizone, characterized by limited textural reorganization and rare development of slaty cleavage in pelitic rocks. Mineral assemblages generally fall within the to prehnite-pumpellyite range, featuring minerals like laumontite, , and mixed-layer / clays, reflecting subtle recrystallization and phase adjustments. Pressure-temperature conditions are moderate, with temperatures of 100–300°C and pressures around 1–5 kbar, driven primarily by geothermal gradients; fluids involved are largely connate waters expelled during compaction, maintaining low in zeolitic assemblages. A prominent example is found in the Tertiary sediments of the Gulf Coast of the , where burial to depths exceeding 3 km results in the transformation of to -rich mixed-layer clays, with the degree of illite ordering and layer percentage serving as a crystallinity index to gauge metamorphic grade. This index increases systematically with depth and temperature, marking the transition from to low-grade metamorphism. If subsequent tectonic activity imparts deformation, metamorphism can evolve into regional metamorphism, though the boundary remains gradational. Burial metamorphism thus represents the initial stages of progressive grade increase in sedimentary sequences.

Classification of Metamorphic Rocks

Metamorphic Grade

Metamorphic grade serves as a proxy for the maximum and conditions experienced by rocks during metamorphism, providing a measure of the intensity of the process. It is typically assessed on a scale ranging from low grade, characterized by facies conditions at temperatures below 200–300°C and low pressures, to high grade in facies at 700–900°C and moderate to high pressures, and even ultra-high grade in eclogite or ultra-high-temperature (UHT) assemblages exceeding 900°C under high pressures. This progression reflects increasing thermal and structural reorganization of the rock, often correlating with metamorphic but distinguished by its focus on overall intensity rather than specific parageneses. Qualitative evaluation of metamorphic grade relies on index minerals, which appear sequentially in pelitic rocks as grade increases, marking the progression from low to high conditions. Common index minerals include in low-grade (around 200–400°C), followed by in the lower (350–500°C), and in medium-grade settings (450–650°C), with higher grades featuring , , and above 600°C. These minerals define reaction isograds, lines on geological maps delineating the first appearance of a specific index mineral due to prograde reactions, such as the isograd from the + → + + H₂O, or the isograd from + + → + K-feldspar + H₂O. Such isograds allow mapping of grade variations across terrains, with the sequence providing a relative scale of metamorphic intensity. For quantitative assessment, geothermobarometry employs compositions to estimate peak and , offering precise values beyond qualitative indices. A widely used method is the - , based on Fe-Mg exchange between coexisting and , which records equilibration temperatures in medium- to high-grade pelitic rocks (typically 450–700°C). Calibrated experimentally, this uses the distribution coefficient KD=(Mg/Fe)[garnet](/page/Garnet)(Mg/Fe)[biotite](/page/Biotite)K_D = \frac{(Mg/Fe)_{[garnet](/page/Garnet)}}{(Mg/Fe)_{[biotite](/page/Biotite)}}, where higher temperatures shift the exchange toward more Mg-rich cores, allowing calculations via empirical equations that account for minor effects (about 40°C per GPa). Complementary barometers, such as those using -plagioclase, refine estimates when combined, enabling reconstruction of the maximum conditions but requiring careful selection of unreset pairs. Variations in metamorphic grade progression are exemplified by Barrovian and Buchan sequences, which reflect different pressure-temperature paths. The Barrovian sequence, typical of regional metamorphism in collisional settings, follows a moderate (around 20–30°C/km) with higher pressures, producing index minerals in the order → → → → → , often under 4–8 kbar and up to 700°C. In contrast, the Buchan sequence occurs in low-pressure environments, such as those influenced by igneous intrusions, with a steeper (>40°C/km) yielding andalusite and alongside and , but lacking , at pressures below 3 kbar and similar temperatures. These sequences highlight how tectonic context influences grade zoning, with Barrovian types dominant in orogenic belts and Buchan in more localized heating regimes. A key limitation in determining metamorphic grade arises from retrograde metamorphism, which can reset or overprint prograde assemblages during cooling and uplift, obscuring peak conditions. Without abundant fluids—unlike during prograde stages—retrograde reactions proceed sluggishly at low temperatures (<400°C), often failing to equilibrate fully and leaving relict high-grade minerals intact while forming partial low-grade overgrowths. This incomplete resetting complicates index mineral identification and geothermobarometric calculations, as altered rims on minerals like garnet may yield erroneously low temperatures, necessitating petrographic and isotopic techniques to distinguish peak from retrograde features.

Metamorphic Facies

Metamorphic facies represent sets of mineral assemblages that are stable together under specific ranges of pressure (P), temperature (T), and fluid conditions, providing a framework for classifying metamorphic rocks based on the physical conditions of their formation. The concept was introduced by Pentti Eskola in 1915, who defined a facies as a group of rocks sharing similar mineral compositions despite variations in protolith chemistry, reflecting equilibrium under coherent P-T fields. Facies are typically named after characteristic rock types, such as or , derived from observations in metabasic rocks. Unlike metamorphic grade, which denotes a general increase in intensity of metamorphism, facies emphasize distinct mineralogical signatures tied to particular P-T regimes. Key metamorphic facies include the blueschist facies, which forms under high-pressure and low-temperature conditions (typically >5 kbar and <500°C), characterized by minerals like glaucophane, lawsonite, and epidote in metabasic rocks. The eclogite facies occurs at ultra-high pressures (>10-15 kbar) and moderate to high temperatures (500-800°C), featuring diagnostic minerals such as omphacite (a sodic clinopyroxene) and pyrope-rich in compositions. In contrast, the granulite facies develops at high temperatures (>700°C) and relatively low pressures (<10 kbar), with assemblages including orthopyroxene, , and , often in dry conditions that inhibit hydrous minerals. Other common facies, like (300-500°C, <5 kbar), feature , , and as index minerals indicating moderate conditions. Facies series describe sequences of facies transitions observed in metamorphic terranes with increasing grade, reflecting tectonic settings and thermal gradients. Akiho Miyashiro (1961) classified these into categories such as the high-pressure Franciscan-type series ( to eclogite) and the low-pressure Buchan-type series ( to ), based on the relative timing of and increases during prograde metamorphism. These series illustrate how evolve coherently across regional belts, with diagnostic minerals like marking entry into greenschist conditions and omphacite signaling eclogite-facies overprinting. Modern refinements to the facies concept incorporate pseudosections, which are calculated phase diagrams tailored to specific bulk rock compositions, allowing for more precise mapping of mineral stability fields under variable chemical conditions. Unlike traditional facies boundaries derived from average compositions, pseudosections account for variability, revealing how minor elements influence assemblage stability and refining P-T estimates for individual samples. This approach, enabled by thermodynamic modeling software, has enhanced the resolution of transitions in complex terranes.

Prograde and Retrograde Sequences

Prograde metamorphism encompasses the mineralogical and textural transformations in rocks driven by increasing and during and heating, leading to a progression from low-grade to high-grade assemblages. This directional evolution typically features dehydration reactions, where hydrous phases like or decompose to produce anhydrous minerals such as or , thereby increasing the variance and complexity of mineral parageneses as the metamorphic grade rises. These changes reflect the rock's response to tectonic , with prograde sequences often transitioning through such as from to conditions. In contrast, retrograde metamorphism involves the adjustments to decreasing and during exhumation and cooling, commonly resulting in hydration reactions that introduce or reform hydrous minerals. A classic example is the overprinting of sericite (fine-grained ) on grains, where reacts with to form sericite + + sodium ions in solution, altering the high-grade prograde fabric without fully reversing it due to kinetic barriers at lower temperatures. Retrograde effects are generally less pervasive than prograde ones, as reduced thermal energy slows reaction rates and limits fluid infiltration, often preserving relict prograde minerals amid partial overprints. The trajectories of these sequences are depicted by pressure-temperature-time (P-T-t) paths, which illustrate the rock's thermal and baric history; regional metamorphism commonly traces loops, with elevating both and before uplift reverses the path, whereas contact metamorphism produces tight, hairpin-shaped loops due to localized, rapid heating followed by swift cooling. Diagnostic evidence includes compositional zoning in porphyroblasts like , where core-to-rim chemical gradients (e.g., increasing Mg/Fe ratios) preserve prograde growth histories, and reaction textures such as or symplectites that signal retrograde hydration or decomposition. Prograde durations typically range from 1 to 10 million years, enabling near-equilibrium recrystallization, while retrograde phases proceed more rapidly—often on the order of hundreds of thousands to a few million years—facilitated by accelerated uplift and fluid access.

Mineral Equilibria and Modeling

Equilibrium Mineral Assemblages

In metamorphic , equilibrium mineral assemblages represent the stable parageneses of that minimize the for a specific bulk rock composition under given (P) and (T) conditions. These assemblages form through prograde reactions where adjust to achieve stability, with the overall system seeking the lowest-energy configuration as dictated by the second law of . The concept underpins the prediction of stability in rocks, allowing petrologists to infer P-T conditions from observed parageneses. The variance of these assemblages is governed by the Gibbs phase rule, which quantifies the (F) available to the system:
F=C−P+2F = C - P + 2
where C is the number of independent chemical components and P is the number of phases (s plus fluid, if present). In simplified applications to solid-dominated metamorphic systems, this rule predicts that most assemblages are divariant (F = 2), stable over broad regions of P-T space defined by the bulk composition. Boundaries between divariant fields mark univariant reactions (F = 1), such as discontinuous net-transfer reactions (e.g., + = + + H₂O), where a specific appears or disappears along a . A practical example is the use of AKF diagrams, which project mineral compositions in the Al₂O₃-K₂O-(FeO + MgO) for pelitic rocks, ignoring SiO₂ if in excess as . In these diagrams, coexisting minerals are connected by tie-lines, delineating divariant compatibility fields; for instance, in low-grade pelites, the assemblage chloritoid + + occupies a field bounded by univariant reactions involving or formation at higher grades. The mineralogical further constrains assemblages, stating that the number of minerals at equilibrium cannot exceed the number of components, ensuring parsimony in observed parageneses (e.g., a three-component yields at most three phases). Real-world deviations from ideal behavior arise due to non-ideal mixing in solid solutions, such as in (NaAlSi₃O₈-KAlSi₃O₈) or (e.g., pyrope-almandine). Activity models correct for these effects by quantifying deviations from , often using Margules or Darken quadratic formalisms to parameterize excess free energy terms that influence phase boundaries and reaction equilibria. These models are essential for accurate thermodynamic calculations, as non-ideality increases at lower temperatures, altering predicted assemblage stabilities compared to ideal approximations.

Petrogenetic Grids

Petrogenetic grids are qualitative pressure-temperature (P-T) diagrams that map the boundaries of metamorphic reactions in simplified chemical systems, illustrating the stability fields of assemblages during metamorphism. These grids typically focus on model compositions such as the KFMASH system (K₂O-FeO-MgO-Al₂O₃-SiO₂-H₂O), which approximates common pelitic rocks by excluding minor components to highlight key phase relations. They consist of univariant reaction lines separating divariant stability fields, with invariant points marking locations where multiple reactions converge, often forming triple points that define topological changes in assemblage stability. The concept of petrogenetic grids emerged in the early 20th century, building on field observations of mineral zonation in metamorphic terranes. Alfred Harker, in his 1932 monograph on metamorphism, provided foundational descriptions of mineral assemblages and their textural relations, laying groundwork for diagrammatic representations of phase equilibria, though without explicit P-T grids. The term "petrogenetic grid" was formalized by Norman L. Bowen in 1940, who proposed networks of equilibrium curves to characterize metamorphic events, drawing from igneous analogies. Significant advancements occurred in the and , with qualitative grids for basic rocks appearing in the , followed by Arlington L. Albee's 1965 schematic grid for metapelites in the KFASH, KMASH, and KFMASH systems, which integrated limited experimental data and Schreinemakers' rules for topological consistency. Construction of petrogenetic grids involves plotting experimentally or thermodynamically derived reaction boundaries on a P-T plane, assuming excess and a phase, to delineate assemblage variance—typically divariant (variance = 2) fields bounded by univariant lines. Invariant points occur where three univariant reactions intersect, such as the convergence of dehydration reactions involving , , and in pelitic systems, allowing prediction of phase-in or phase-out sequences. These grids are constructed for specific bulk compositions to simplify the natural variability, emphasizing net-transfer reactions that control prograde growth. Petrogenetic grids are primarily used to predict the stability of mineral assemblages across P-T space and to interpret the topological shifts at triple points, where small changes in conditions can lead to abrupt assemblage changes, aiding in the reconstruction of metamorphic conditions from rock samples. For instance, they help identify petrogenetic indicator reactions, like the breakdown of to form or , constraining the pressure regime of regional metamorphism. A representative example is the petrogenetic grid for metapelites in the KFMASH system, which delineates boundaries between low-grade assemblages (e.g., + ) and higher-grade ones (e.g., + + ), including reactions such as + + quartz = + H₂O. This grid also incorporates melting boundaries, contrasting subsolidus paths with vapor-absent curves, such as melting to form K-feldspar + + melt, which become prominent above approximately 700–800°C at mid-crustal pressures. Such diagrams, as refined by Thompson (1976), illustrate how invariant points like the [Ms + Chl + Bt + Gt] point control the transition from hydrous to assemblages in pelitic sequences. Modern extensions of petrogenetic grids include pseudosections, which are calculated P-T phase diagrams for a fixed bulk rock composition using thermodynamic datasets and software such as THERMOCALC or Perple_X. Unlike qualitative grids for model systems, pseudosections account for specific effective compositions and solid solutions, enabling detailed P-T path reconstruction for individual rocks and incorporating more components for accuracy. These tools have advanced significantly since the , with updates to thermodynamic databases improving predictions as of 2020. Despite their utility, petrogenetic grids assume attainment of chemical equilibrium among phases, which may not reflect natural kinetic barriers that inhibit reaction completion during rapid heating or deformation. They also neglect the effects of fluid composition variations and minor elements, potentially oversimplifying in complex natural systems.

Compatibility Diagrams

Compatibility diagrams are projective representations used in metamorphic petrology to illustrate the mineral assemblages compatible with specific bulk rock compositions at fixed and conditions, facilitating the analysis of phase relations in multicomponent systems. These diagrams simplify complex chemical systems by projecting from excess components, typically (SiO₂) and (H₂O), which are assumed to be ubiquitous and saturated, thereby reducing the dimensionality to three key components. Developed initially by Pentti Eskola in the early , they provide a graphical framework for understanding mineral compatibilities without requiring full thermodynamic calculations. The two primary types are the ACF and AKF diagrams. The ACF diagram, standing for Al₂O₃-CaO-(FeO+MgO+MnO), is suited for Ca- and Al-rich rocks such as metabasites and marbles, where it projects from excess K₂O (assigned to alkali feldspar) and Na₂O (to ). In contrast, the AKF diagram, for Al₂O₃-K₂O-(FeO+MgO+MnO), targets K-rich, Ca-poor pelitic rocks like metasediments, projecting from excess CaO (to or other Ca-minerals). Both assume excess and H₂O, allowing focus on the variable components that control mineral stability. A related variant, the AFM diagram (Al₂O₃-FeO-MgO), projects further from K₂O (assigned to or ) and is commonly applied to Fe-Mg-rich assemblages in both pelitic and mafic rocks. Construction of these diagrams involves normalizing the bulk rock composition to the three components after subtracting fixed phases. For instance, in an ACF diagram, the molecular proportions of Al₂O₃ (minus Na₂O and K₂O), CaO (minus contributions from phosphates), and FeO+MgO+MnO are summed and renormalized to 100%, then plotted as coordinates in a ternary triangle. compositions are similarly projected and connected: divariant assemblages (invariant in the projection) form tie-triangles enclosing compatible three- fields, while univariant boundaries are tie-lines linking two . rules govern field topology, ensuring that fields do not cross tie-lines and that higher-variance assemblages (e.g., two- fields) adjacent to lower-variance ones (e.g., tie-triangles), maintaining thermodynamic consistency. This structure adheres to the , where the number of phases in an assemblage indicates the variance () in the projected system. These diagrams enable applications such as identifying possible metamorphic reactions through tie-line flips, where a change in conditions shifts a stable tie-line to cross another, indicating a univariant reaction (e.g., + → + in AKF). They also assess variance by counting phases: a three-phase tie-triangle implies divariance (F=2, fixed P-T), while excess phases suggest incompatibility or . Modern computational tools like THERMOCALC, developed by and Powell, generate calculated compatibility diagrams incorporating thermodynamic datasets and solid solutions for more precise, isobaric-isothermal projections. An illustrative example is the AFM diagram for mafic rocks in the amphibolite facies, where and (projected out) define stable fields with or ; bulk compositions plotting within the amphibole- tie-line indicate their coexistence, while those in amphibole- fields reflect higher-grade conditions.

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