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Magma can be found in the mantle or the crust.

Magma (from Ancient Greek μάγμα (mágma) 'thick unguent')[1] is the molten or semi-molten natural material from which all igneous rocks are formed.[2] Magma (sometimes colloquially but incorrectly referred to as lava) is found beneath the surface of the Earth, and evidence of magmatism has also been discovered on other terrestrial planets and some natural satellites.[3] Besides molten rock, magma may also contain suspended crystals and gas bubbles.[4]

Magma is produced by melting of the mantle or the crust in various tectonic settings, which on Earth include subduction zones, continental rift zones,[5] mid-ocean ridges and hotspots. Mantle and crustal melts migrate upwards through the crust where they are thought to be stored in magma chambers[6] or trans-crustal crystal-rich mush zones.[7] During magma's storage in the crust, its composition may be modified by fractional crystallization, contamination with crustal melts, magma mixing, and degassing. Following its ascent through the crust, magma may feed a volcano and be extruded as lava, or it may solidify underground to form an intrusion,[8] such as a dike, a sill, a laccolith, a pluton, or a batholith.[9]

While the study of magma has relied on observing magma after its transition into a lava flow, magma has been encountered in situ three times during geothermal drilling projects, twice in Iceland (see Use in energy production) and once in Hawaii.[10][11][12][13]

Physical and chemical properties

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Magma consists of liquid rock that usually contains suspended solid crystals.[14] As magma approaches the surface and the overburden pressure drops, dissolved gases bubble out of the liquid, so that magma near the surface consists of materials in solid, liquid, and gas phases.[15]

Composition

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See also: Igneous differentiation

Most magma is rich in silica.[8] Rare nonsilicate magma can form by local melting of nonsilicate mineral deposits[16] or by separation of a magma into separate immiscible silicate and nonsilicate liquid phases.[17]

Silicate magmas are molten mixtures dominated by oxygen and silicon, the most abundant chemical elements in the Earth's crust, with smaller quantities of aluminium, calcium, magnesium, iron, sodium, and potassium, and minor amounts of many other elements.[18] Petrologists routinely express the composition of a silicate magma in terms of the weight or molar mass fraction of the oxides of the major elements (other than oxygen) present in the magma.[19]

Because many of the properties of a magma (such as its viscosity and temperature) are observed to correlate with silica content, silicate magmas are divided into four chemical types based on silica content: felsic, intermediate, mafic, and ultramafic.[20]

Felsic magmas

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Felsic or silicic magmas have a silica content greater than 63%. They include rhyolite and dacite magmas. With such a high silica content, these magmas are extremely viscous, ranging from 108 cP (105 Pa⋅s) for hot rhyolite magma at 1,200 °C (2,190 °F) to 1011 cP (108 Pa⋅s) for cool rhyolite magma at 800 °C (1,470 °F).[21] For comparison, water has a viscosity of about 1 cP (0.001 Pa⋅s). Because of this very high viscosity, felsic lavas usually erupt explosively to produce pyroclastic (fragmental) deposits. However, rhyolite lavas occasionally erupt effusively to form lava spines, lava domes or "coulees" (which are thick, short lava flows).[22] The lavas typically fragment as they extrude, producing block lava flows. These often contain obsidian.[23]

Felsic lavas can erupt at temperatures as low as 800 °C (1,470 °F).[24] Unusually hot (>950 °C; >1,740 °F) rhyolite lavas, however, may flow for distances of many tens of kilometres, such as in the Snake River Plain of the northwestern United States.[25]

Intermediate magmas

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Intermediate or andesitic magmas contain 52% to 63% silica, and are lower in aluminium and usually somewhat richer in magnesium and iron than felsic magmas. Intermediate lavas form andesite domes and block lavas, and may occur on steep composite volcanoes, such as in the Andes.[26] They are also commonly hotter, in the range of 850 to 1,100 °C (1,560 to 2,010 °F)). Because of their lower silica content and higher eruptive temperatures, they tend to be much less viscous, with a typical viscosity of 3.5 million cP (3,500 Pa⋅s) at 1,200 °C (2,190 °F). This is slightly greater than the viscosity of smooth peanut butter.[27] Intermediate magmas show a greater tendency to form phenocrysts.[28] Higher iron and magnesium tends to manifest as a darker groundmass, including amphibole or pyroxene phenocrysts.[29]

Mafic magmas

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Mafic or basaltic magmas have a silica content of 52% to 45%. They are typified by their high ferromagnesian content, and generally erupt at temperatures of 1,100 to 1,200 °C (2,010 to 2,190 °F). Viscosities can be relatively low, around 104 to 105 cP (10 to 100 Pa⋅s), although this is still many orders of magnitude higher than water. This viscosity is similar to that of ketchup.[30] Basalt lavas tend to produce low-profile shield volcanoes or flood basalts, because the fluidal lava flows for long distances from the vent. The thickness of a basalt lava, particularly on a low slope, may be much greater than the thickness of the moving lava flow at any one time, because basalt lavas may "inflate" by supply of lava beneath a solidified crust.[31] Most basalt lavas are of ʻAʻā or pāhoehoe types, rather than block lavas. Underwater, they can form pillow lavas, which are rather similar to entrail-type pahoehoe lavas on land.[32]

Ultramafic magmas

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Ultramafic magmas, such as picritic basalt, komatiite, and highly magnesian magmas that form boninite, take the composition and temperatures to the extreme. All have a silica content under 45%. Komatiites contain over 18% magnesium oxide, and are thought to have erupted at temperatures of 1,600 °C (2,910 °F). At this temperature there is practically no polymerization of the mineral compounds, creating a highly mobile liquid.[33] Viscosities of komatiite magmas are thought to have been as low as 100 to 1000 cP (0.1 to 1 Pa⋅s), similar to that of light motor oil.[21] Most ultramafic lavas are no younger than the Proterozoic, with a few ultramafic magmas known from the Phanerozoic in Central America that are attributed to a hot mantle plume. No modern komatiite lavas are known, as the Earth's mantle has cooled too much to produce highly magnesian magmas.[34]

Alkaline magmas

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See also: Alkaline magma series

Some silicic magmas have an elevated content of alkali metal oxides (sodium and potassium), particularly in regions of continental rifting, areas overlying deeply subducted plates, or at intraplate hotspots.[35] Their silica content can range from ultramafic (nephelinites, basanites and tephrites) to felsic (trachytes). They are more likely to be generated at greater depths in the mantle than subalkaline magmas.[36] Olivine nephelinite magmas are both ultramafic and highly alkaline, and are thought to have come from much deeper in the mantle of the Earth than other magmas.[37]

Examples of magma compositions (wt%)[38]
Component Nephelinite Tholeiitic picrite Tholeiitic basalt Andesite Rhyolite
SiO2 39.7 46.4 53.8 60.0 73.2
TiO2 2.8 2.0 2.0 1.0 0.2
Al2O3 11.4 8.5 13.9 16.0 14.0
Fe2O3 5.3 2.5 2.6 1.9 0.6
FeO 8.2 9.8 9.3 6.2 1.7
MnO 0.2 0.2 0.2 0.2 0.0
MgO 12.1 20.8 4.1 3.9 0.4
CaO 12.8 7.4 7.9 5.9 1.3
Na2O 3.8 1.6 3.0 3.9 3.9
K2O 1.2 0.3 1.5 0.9 4.1
P2O5 0.9 0.2 0.4 0.2 0.0
Tholeiitic basalt magma
  1. SiO2 (53.8%)
  2. Al2O3 (13.9%)
  3. FeO (9.30%)
  4. CaO (7.90%)
  5. MgO (4.10%)
  6. Na2O (3.00%)
  7. Fe2O3 (2.60%)
  8. TiO2 (2.00%)
  9. K2O (1.50%)
  10. P2O5 (0.40%)
  11. MnO (0.20%)
Rhyolite magma
  1. SiO2 (73.2%)
  2. Al2O3 (14.0%)
  3. FeO (1.70%)
  4. CaO (1.30%)
  5. MgO (0.40%)
  6. Na2O (3.90%)
  7. Fe2O3 (0.60%)
  8. TiO2 (0.20%)
  9. K2O (4.10%)
  10. P2O5 (0.00%)
  11. MnO (0.00%)

Non-silicate magmas

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Some lavas of unusual composition have erupted onto the surface of the Earth. These include:

  • Carbonatite and natrocarbonatite lavas are known from Ol Doinyo Lengai volcano in Tanzania, which is the sole example of an active carbonatite volcano.[39] Carbonatites in the geologic record are typically 75% carbonate minerals, with lesser amounts of silica-undersaturated silicate minerals (such as micas and olivine), apatite, magnetite, and pyrochlore. This may not reflect the original composition of the lava, which may have included sodium carbonate that was subsequently removed by hydrothermal activity, though laboratory experiments show that a calcite-rich magma is possible. Carbonatite lavas show stable isotope ratios indicating they are derived from the highly alkaline silicic lavas with which they are always associated, probably by separation of an immiscible phase.[40] Natrocarbonatite lavas of Ol Doinyo Lengai are composed mostly of sodium carbonate, with about half as much calcium carbonate and half again as much potassium carbonate, and minor amounts of halides, fluorides, and sulphates. The lavas are extremely fluid, with viscosities only slightly greater than water, and are very cool, with measured temperatures of 491 to 544 °C (916 to 1,011 °F).[41]
  • Iron oxide magmas are thought to be the source of the iron ore at Kiruna, Sweden which formed during the Proterozoic.[17] Iron oxide lavas of Pliocene age occur at the El Laco volcanic complex on the Chile-Argentina border.[16] Iron oxide lavas are thought to be the result of immiscible separation of iron oxide magma from a parental magma of calc-alkaline or alkaline composition.[17] When erupted, the temperature of the molten iron oxide magma is about 700 to 800 °C (1,292 to 1,472 °F).[42]
  • Sulfur lava flows up to 250 metres (820 feet) long and 10 metres (33 feet) wide occur at Lastarria volcano, Chile. They were formed by the melting of sulfur deposits at temperatures as low as 113 °C (235 °F).[16]

Magmatic gases

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Main article: Volcanic gas

The concentrations of different gases can vary considerably. Water vapor is typically the most abundant magmatic gas, followed by carbon dioxide[43] and sulfur dioxide. Other principal magmatic gases include hydrogen sulfide, hydrogen chloride, and hydrogen fluoride.[44]

The solubility of magmatic gases in magma depends on pressure, magma composition, and temperature. Magma that is extruded as lava is extremely dry, but magma at depth and under great pressure can contain a dissolved water content in excess of 10%. Water is somewhat less soluble in low-silica magma than high-silica magma, so that at 1,100 °C and 0.5 GPa, a basaltic magma can dissolve 8% H2O while a granite pegmatite magma can dissolve 11% H2O.[45] However, magmas are not necessarily saturated under typical conditions.

Water concentrations in magmas (wt%)[46]
Magma composition H2O concentration
wt %
MORB (tholeiites) 0.1 – 0.2
Island tholeiite 0.3 – 0.6
Alkali basalts 0.8 – 1.5
Volcanic arc basalts 2–4
Basanites and nephelinites 1.5–2
Island arc andesites and dacites 1–3
Continental margin andesites and dacites 2–5
Rhyolites up to 7

Carbon dioxide is much less soluble in magmas than water, and frequently separates into a distinct fluid phase even at great depth. This explains the presence of carbon dioxide fluid inclusions in crystals formed in magmas at great depth.[46]

Rheology

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Graph showing logarithmic variation of magma viscosity (η) with silica content for three temperatures

Viscosity is a key melt property in understanding the behaviour of magmas. Whereas temperatures in common silicate lavas range from about 800 °C (1,470 °F) for felsic lavas to 1,200 °C (2,190 °F) for mafic lavas,[24] the viscosity of the same lavas ranges over seven orders of magnitude, from 104 cP (10 Pa⋅s) for mafic lava to 1011 cP (108 Pa⋅s) for felsic magmas.[24] The viscosity is mostly determined by composition but is also dependent on temperature.[21] The tendency of felsic lava to be cooler than mafic lava increases the viscosity difference.

The silicon ion is small and highly charged, and so it has a strong tendency to coordinate with four oxygen ions, which form a tetrahedral arrangement around the much smaller silicon ion. This is called a silica tetrahedron. In a magma that is low in silicon, these silica tetrahedra are isolated, but as the silicon content increases, silica tetrahedra begin to partially polymerize, forming chains, sheets, and clumps of silica tetrahedra linked by bridging oxygen ions. These greatly increase the viscosity of the magma.[47]

  • A single silica tetrahedron
    A single silica tetrahedron
  • Two silica tetrahedra joined by a bridging oxygen ion (tinted pink)
    Two silica tetrahedra joined by a bridging oxygen ion (tinted pink)

The tendency towards polymerization is expressed as NBO/T, where NBO is the number of non-bridging oxygen ions and T is the number of network-forming ions. Silicon is the main network-forming ion, but in magmas high in sodium, aluminium also acts as a network former, and ferric iron can act as a network former when other network formers are lacking. Most other metallic ions reduce the tendency to polymerize and are described as network modifiers. In a hypothetical magma formed entirely from melted silica, NBO/T would be 0, while in a hypothetical magma so low in network formers that no polymerization takes place, NBO/T would be 4. Neither extreme is common in nature, but basalt magmas typically have NBO/T between 0.6 and 0.9, andesitic magmas have NBO/T of 0.3 to 0.5, and rhyolitic magmas have NBO/T of 0.02 to 0.2. Water acts as a network modifier, and dissolved water drastically reduces melt viscosity. Carbon dioxide neutralizes network modifiers, so dissolved carbon dioxide increases the viscosity. Higher-temperature melts are less viscous, since more thermal energy is available to break bonds between oxygen and network formers.[15]

Most magmas contain solid crystals of various minerals, fragments of exotic rocks known as xenoliths and fragments of previously solidified magma. The crystal content of most magmas gives them thixotropic and shear thinning properties.[48] In other words, most magmas do not behave like Newtonian fluids, in which the rate of flow is proportional to the shear stress. Instead, a typical magma is a Bingham fluid, which shows considerable resistance to flow until a stress threshold, called the yield stress, is crossed.[49] This results in plug flow of partially crystalline magma. A familiar example of plug flow is toothpaste squeezed out of a toothpaste tube. The toothpaste comes out as a semisolid plug, because shear is concentrated in a thin layer in the toothpaste next to the tube, and only here does the toothpaste behave as a fluid. Thixotropic behavior also hinders crystals from settling out of the magma.[50] Once the crystal content reaches about 60%, the magma ceases to behave like a fluid and begins to behave like a solid. Such a mixture of crystals with melted rock is sometimes described as crystal mush.[51]

Magma is typically also viscoelastic, meaning it flows like a liquid under low stresses, but once the applied stress exceeds a critical value, the melt cannot dissipate the stress fast enough through relaxation alone, resulting in transient fracture propagation. Once stresses are reduced below the critical threshold, the melt viscously relaxes once more and heals the fracture.[52]

Temperature

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Temperatures of molten lava, which is magma extruded onto the surface, are almost all in the range 700 to 1,400 °C (1,300 to 2,600 °F), but very rare carbonatite magmas may be as cool as 490 °C (910 °F),[53] and komatiite magmas may have been as hot as 1,600 °C (2,900 °F).[54] Magma has occasionally been encountered during drilling in geothermal fields, including drilling in Hawaii that penetrated a dacitic magma body at a depth of 2,488 m (8,163 ft). The temperature of this magma was estimated at 1,050 °C (1,920 °F). Temperatures of deeper magmas must be inferred from theoretical computations and the geothermal gradient.[13]

Most magmas contain some solid crystals suspended in the liquid phase. This indicates that the temperature of the magma lies between the solidus, which is defined as the temperature at which the magma completely solidifies, and the liquidus, defined as the temperature at which the magma is completely liquid.[14] Calculations of solidus temperatures at likely depths suggests that magma generated beneath areas of rifting starts at a temperature of about 1,300 to 1,500 °C (2,400 to 2,700 °F). Magma generated from mantle plumes may be as hot as 1,600 °C (2,900 °F). The temperature of magma generated in subduction zones, where water vapor lowers the melting temperature, may be as low as 1,060 °C (1,940 °F).[55]

Density

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Magma densities depend mostly on composition, iron content being the most important parameter.[56]

Type Density (kg/m3)
Basaltic magma 2650–2800
Andesitic magma 2450–2500
Rhyolitic magma 2180–2250

Magma expands slightly at lower pressure or higher temperature.[56] When magma approaches the surface, its dissolved gases begin to bubble out of the liquid. These bubbles had significantly reduced the density of the magma at depth and helped drive it toward the surface in the first place.[57]

Origins

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The temperature within the interior of the earth is described by the geothermal gradient, which is the rate of temperature change with depth. The geothermal gradient is established by the balance between heating through radioactive decay in the Earth's interior and heat loss from the surface of the earth. The geothermal gradient averages about 25 °C/km in the Earth's upper crust, but this varies widely by region, from a low of 5–10 °C/km within oceanic trenches and subduction zones to 30–80 °C/km along mid-ocean ridges or near mantle plumes.[58] The gradient becomes less steep with depth, dropping to just 0.25 to 0.3 °C/km in the mantle, where slow convection efficiently transports heat. The average geothermal gradient is not normally steep enough to bring rocks to their melting point anywhere in the crust or upper mantle, so magma is produced only where the geothermal gradient is unusually steep or the melting point of the rock is unusually low. However, the ascent of magma towards the surface in such settings is the most important process for transporting heat through the crust of the Earth.[59]

Rocks may melt in response to a decrease in pressure,[60] to a change in composition (such as an addition of water),[61] to an increase in temperature,[62] or to a combination of these processes.[63] Other mechanisms, such as melting from a meteorite impact, are less important today, but impacts during the accretion of the Earth led to extensive melting, and the outer several hundred kilometers of the early Earth was probably a magma ocean.[64] Impacts of large meteorites in the last few hundred million years have been proposed as one mechanism responsible for the extensive basalt magmatism of several large igneous provinces.[65]

Decompression

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Decompression melting occurs because of a decrease in pressure.[66] It is the most important mechanism for producing magma from the upper mantle.[67]

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 the upward movement of solid mantle is critical in the evolution of the Earth.[63]

Decompression melting creates the ocean crust at mid-ocean ridges, making it by far the most important source of magma on Earth.[67] It also causes volcanism in intraplate regions, such as Europe, Africa and the Pacific sea floor. Intraplate volcanism is attributed to the rise of mantle plumes or to intraplate extension, with the importance of each mechanism being a topic of continuing research.[68]

Effects of water and carbon dioxide

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The change of rock composition most responsible for the 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 1,500 °C in the absence of water.[69] Water is driven out of the oceanic lithosphere in subduction zones, and it causes melting in the overlying mantle. Hydrous magmas with the composition of basalt or andesite 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.[70] These magmas form rocks of the calc-alkaline series, an important part of the continental crust.[71] With low density and viscosity, hydrous magmas are highly buoyant and will move upwards in Earth's mantle.[72]

The addition of carbon dioxide is relatively a much less important cause of magma formation than the 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.[73] 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.[74][75]

Temperature increase

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Increase in 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.[76] 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 1,000 kilometers within the middle crust along the southern margin of the Tibetan Plateau.[77] Granite and rhyolite are types of igneous rock commonly interpreted as products of the melting of continental crust because of increases in temperature. Temperature increases also may contribute to the melting of lithosphere dragged down in a subduction zone.[citation needed]

The melting process

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Phase diagram for the diopside-anorthite system

When rocks melt, they do so over a range of temperature, because most rocks are made of several minerals, which all have different melting points. The temperature at which the first melt appears (the solidus) is lower than the melting temperature of any one of the pure minerals. This is similar to the lowering of the melting point of ice when it is mixed with salt. The first melt is called the eutectic and has a composition that depends on the combination of minerals present.[78]

For example, a mixture of anorthite and diopside, which are two of the predominant minerals in basalt, begins to melt at about 1274 °C. This is well below the melting temperatures of 1392 °C for pure diopside and 1553 °C for pure anorthite. The resulting melt is composed of about 43 wt% anorthite.[79] As additional heat is added to the rock, the temperature remains at 1274 °C until either the anorthite or diopside is fully melted. The temperature then rises as the remaining mineral continues to melt, which shifts the melt composition away from the eutectic. For example, if the content of anorthite is greater than 43%, the entire supply of diopside will melt at 1274 °C., along with enough of the anorthite to keep the melt at the eutectic composition. Further heating causes the temperature to slowly rise as the remaining anorthite gradually melts and the melt becomes increasingly rich in anorthite liquid. If the mixture has only a slight excess of anorthite, this will melt before the temperature rises much above 1274 °C. If the mixture is almost all anorthite, the temperature will reach nearly the melting point of pure anorthite before all the anorthite is melted. If the anorthite content of the mixture is less than 43%, then all the anorthite will melt at the eutectic temperature, along with part of the diopside, and the remaining diopside will then gradually melt as the temperature continues to rise.[78]

Because of eutectic melting, the composition of the melt can be quite different from the source rock. For example, a mixture of 10% anorthite with diopside could experience about 23% partial melting before the melt deviated from the eutectic, which has the composition of about 43% anorthite. This effect of partial melting is reflected in the compositions of different magmas. A low degree of partial melting of the upper mantle (2% to 4%) can produce highly alkaline magmas such as melilitites, while a greater degree of partial melting (8% to 11%) can produce alkali olivine basalt.[80] Oceanic magmas likely result from partial melting of 3% to 15% of the source rock.[81] Some calk-alkaline granitoids may be produced by a high degree of partial melting, as much as 15% to 30%.[82] High-magnesium magmas, such as komatiite and picrite, may also be the products of a high degree of partial melting of mantle rock.[83]

Certain chemical elements, called incompatible elements, have a combination of ionic radius and ionic charge that is unlike that of the more abundant elements in the source rock. The ions of these elements fit rather poorly in the structure of the minerals making up the source rock, and readily leave the solid minerals to become highly concentrated in melts produced by a low degree of partial melting. Incompatible elements commonly include potassium, barium, caesium, and rubidium, which are large and weakly charged (the large-ion lithophile elements, or LILEs), as well as elements whose ions carry a high charge (the high-field-strength elements, or HSFEs), which include such elements as zirconium, niobium, hafnium, tantalum, the rare-earth elements, and the actinides. Potassium can become so enriched in melt produced by a very low degree of partial melting that, when the magma subsequently cools and solidifies, it forms unusual potassic rock such as lamprophyre, lamproite, or kimberlite.[84]

When enough rock is melted, the small globules of melt (generally occurring between mineral grains) link up and soften the rock. Under pressure within the earth, as little as a fraction of a percent of partial melting may be sufficient to cause melt to be squeezed from its source.[85] Melt rapidly separates from its source rock once the degree of partial melting exceeds 30%. However, usually much less than 30% of a magma source rock is melted before the heat supply is exhausted.[86]

Pegmatite may be produced by low degrees of partial melting of the crust.[87] Some granite-composition magmas are eutectic (or cotectic) melts, and they may be produced by low to high degrees of partial melting of the crust, as well as by fractional crystallization.[88]

Evolution of magmas

[edit]
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.
Main article: Igneous differentiation

Most magmas are fully melted only for small parts of their histories. More typically, they are mixes of melt and crystals, and sometimes also of gas bubbles.[15] Melt, crystals, and bubbles usually have different densities, and so they can separate as magmas evolve.[89]

As magma cools, minerals typically crystallize from the melt at different temperatures. This resembles the original melting process in reverse. However, because the melt has usually separated from its original source rock and moved to a shallower depth, the reverse process of crystallization is not precisely identical. For example, if a melt was 50% each of diopside and anorthite, then anorthite would begin crystallizing from the melt at a temperature somewhat higher than the eutectic temperature of 1274 °C. This shifts the remaining melt towards its eutectic composition of 43% diopside. The eutectic is reached at 1274 °C, the temperature at which diopside and anorthite begin crystallizing together. If the melt was 90% diopside, the diopside would begin crystallizing first until the eutectic was reached.[90]

If the crystals remained suspended in the melt, the crystallization process would not change the overall composition of the melt plus solid minerals. This situation is described as equillibrium crystallization. However, in a series of experiments culminating in his 1915 paper, Crystallization-differentiation in silicate liquids,[91] Norman L. Bowen demonstrated that crystals of olivine and diopside that crystallized out of a cooling melt of forsterite, diopside, and silica would sink through the melt on geologically relevant time scales. Geologists subsequently found considerable field evidence of such fractional crystallization.[89]

When crystals separate from a magma, then the residual magma 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.[92] Gabbro may have a liquidus temperature near 1,200 °C,[93] and the derivative granite-composition melt may have a liquidus temperature as low as about 700 °C.[94] 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.[89]

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. Assimilation near the roof of a magma chamber and fractional crystallization near its base can even take place simultaneously. Magmas of different compositions can mix with one another. In rare cases, melts can separate into two immiscible melts of contrasting compositions.[95]

Primary magmas

[edit]

When rock melts, the liquid is a primary magma. Primary magmas have not undergone any differentiation and represent the starting composition of a magma.[96] In practice, it is difficult to unambiguously identify primary magmas,[97] though it has been suggested that boninite is a variety of andesite crystallized from a primary magma.[98] The Great Dyke of Zimbabwe has also been interpreted as rock crystallized from a primary magma.[99] The interpretation of leucosomes of migmatites as primary magmas is contradicted by zircon data, which suggests leucosomes are a residue (a cumulate rock) left by extraction of a primary magma.[100]

Parental magma

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When it is impossible to find the primitive or primary magma composition, it is often useful to attempt to identify a parental magma.[97] A parental magma is a magma composition from which the observed range of magma chemistries has been derived by the processes of igneous differentiation. It need not be a primitive melt.[101]

For instance, a series of basalt flows are assumed to be related to one another. A composition from which they could reasonably be produced by fractional crystallization is termed a parental magma. Fractional crystallization models would be produced to test the hypothesis that they share a common parental magma.[102]

Migration and solidification

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Magma develops within the mantle or crust where the temperature and pressure conditions favor the molten state. After its formation, magma buoyantly rises toward the Earth's surface, due to its lower density than the source rock.[103] As it migrates through the crust, magma may collect and reside in magma chambers (though recent work suggests that magma may be stored in trans-crustal crystal-rich mush zones rather than dominantly liquid magma chambers [7]). Magma can remain in a chamber until it either cools and crystallizes to form intrusive rock, it erupts as a volcano, or it moves into another magma chamber.[citation needed]

Plutonism

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When magma cools it begins to form solid mineral phases. Some of these settle at the bottom of the magma chamber forming cumulates that might form mafic layered intrusions. Magma that cools slowly within a magma chamber usually ends up forming bodies of plutonic rocks such as gabbro, diorite and granite, depending upon the composition of the magma. Alternatively, if the magma is erupted it forms volcanic rocks such as basalt, andesite and rhyolite (the extrusive equivalents of gabbro, diorite and granite, respectively).[citation needed]

Volcanism

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Main article: Volcanism

Magma that is extruded onto the surface during a volcanic eruption is called lava. Lava cools and solidifies relatively quickly compared to underground bodies of magma. This fast cooling does not allow crystals to grow large, and a part of the melt does not crystallize at all, becoming glass. Rocks largely composed of volcanic glass include obsidian, scoria and pumice.

Before and during volcanic eruptions, volatiles such as CO2 and H2O partially leave the melt through a process known as exsolution. Magma with low water content becomes increasingly viscous. If massive exsolution occurs when magma heads upwards during a volcanic eruption, the resulting eruption is usually explosive.[104]

Use in energy production

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The Iceland Deep Drilling Project, while drilling several 5,000 m holes in an attempt to harness the heat in the volcanic bedrock below the surface of Iceland, struck a pocket of magma at 2,100 m in 2009. Because this was only the third time in recorded history that magma had been reached, IDDP decided to invest in the hole, naming it IDDP-1.[105]

A cemented steel case was constructed in the hole with a perforation at the bottom close to the magma. The high temperatures and pressure of the magma steam were used to generate 36 MW of power, making IDDP-1 the world's first magma-enhanced geothermal system.[105]

References

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Magma

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Magma is a hot, molten or semi-molten mixture of liquid rock, suspended crystals, and dissolved gases found beneath the Earth's surface, serving as the material for igneous rocks. When this material erupts through the surface, it is termed lava, distinguishing it from its subsurface state where it remains under . Magma forms through of existing rocks in the Earth's or lower crust, a process driven by mechanisms including elevated temperatures, decreased (decompression), and introduction of or other volatiles that lower points. These events commonly occur in diverse tectonic environments, such as zones where oceanic plates descend beneath continental plates, divergent boundaries like mid-ocean ridges, and intraplate hotspots. The chemical composition of magma, dominated by , varies significantly and is classified primarily by silica (SiO₂) content into (basaltic, 45–52% SiO₂, low and high ), intermediate (andesitic, 52–63% SiO₂), and (rhyolitic, 65–75% SiO₂, high and lower ) types. These compositional differences, along with volatile content (typically 1–6% water, CO₂, and SO₂), profoundly affect magma's flow behavior, eruption explosivity, and the resulting rock textures upon cooling. Magma's movement and solidification play a central role in the rock cycle, driving , continental growth, and .

Physical and Chemical Properties

Chemical Composition

Magma is predominantly composed of materials, with (SiO₂) constituting 45–75% by weight of the total composition, forming the backbone of its structure through . The primary elements include oxygen and as the most abundant, followed by aluminum, , calcium, sodium, , and potassium, which combine to form key s such as aluminum oxide (Al₂O₃, typically 12–18%), (FeO, 2–10%), (MgO, 1–10%), (CaO, 3–12%), (Na₂O, 2–5%), (K₂O, 0.5–5%), and (TiO₂, 0.5–2%). These oxides vary in proportion depending on the magma's source and evolution, but they collectively determine the mineral assemblage and overall behavior during ./07:_Igneous_Rocks/7.01:_Magma_and_How_It_Forms) Magmas are classified primarily by their silica content, which influences and associated rock types. Ultramafic magmas contain less than 45% SiO₂ and are rich in magnesium and iron, exemplified by komatiites derived from high-temperature melts. magmas have 45–52% SiO₂, featuring abundant ferromagnesian minerals like and , as seen in basaltic lavas. Intermediate magmas range from 52–63% SiO₂, balancing mafic and felsic components with minerals such as and , typical of andesites. Felsic magmas contain 63–75% SiO₂, dominated by and alkali feldspars, characteristic of rhyolitic compositions. Alkaline magmas deviate from the calc-alkaline series by being enriched in sodium and potassium oxides relative to silica, often with Na₂O + K₂O exceeding 5–12% depending on the total alkali-silica (TAS) diagram classification. These magmas produce rocks like phonolites or trachytes and are commonly associated with intraplate settings, such as rift zones. Carbonatites serve as examples, featuring high levels of carbonate minerals (>50 vol%) with low silica (<20%), marking them as distinctly alkaline. Non-silicate magmas are rare and include carbonatite melts, which are primarily carbonate-based rather than silica-dominated, originating from mantle sources with elevated CO₂ content. Sulfur-rich melts, such as immiscible sulfide liquids, also occur in specialized environments like carbonated alkaline systems, where they segregate from silicate phases and concentrate chalcophile elements. These non-silicate types highlight the diversity beyond typical silicate-dominated magmas. The stability and sequence of minerals in cooling magmas are described by Bowen's reaction series, which outlines the progressive crystallization from mafic to felsic phases. In the discontinuous branch, high-temperature minerals like olivine react to form pyroxene, then amphibole, biotite, and orthoclase; the continuous branch shows plagioclase evolving from calcium-rich to sodium-rich compositions. This series, based on experimental petrology, explains mineral associations in igneous rocks without exhaustive listings of all variants.

Physical Characteristics

Magma temperatures generally range from 700 to 1300°C, with variations primarily driven by chemical composition; felsic magmas typically exist at 700–900°C, whereas mafic magmas are hotter, spanning 1000–1300°C. These temperature baselines are influenced by the magma's silicate content, as higher silica concentrations in felsic types lower the melting point compared to iron- and magnesium-rich mafic compositions. Magma density varies between 2.2 and 3.0 g/cm³, generally decreasing as temperature rises due to thermal expansion and increasing with greater mafic mineral content, which adds denser components like olivine and pyroxene. This property can be approximated using the relation for thermal effects on density: ρ≈ρ0(1−αΔT)\rho \approx \rho_0 (1 - \alpha \Delta T) where ρ0\rho_0 is the reference density, α\alpha is the thermal expansion coefficient (typically 10^{-4} to 10^{-5} K^{-1} for silicate melts), and ΔT\Delta T is the temperature change from the reference state. The rheological behavior of magma is characterized by viscosities ranging from 10¹ to 10¹⁴ Pa·s, with low-viscosity values for hot, mafic types and high values for cooler, silica-rich varieties. Magmas with low crystal content flow as Newtonian fluids, where shear stress is linearly proportional to strain rate, but the presence of crystals introduces non-Newtonian properties, often modeled as a Bingham fluid with a yield strength given by: τ=τ0+ηγ˙\tau = \tau_0 + \eta \dot{\gamma} where τ\tau is shear stress, τ0\tau_0 is the yield stress (typically 10²–10⁴ Pa for crystal-laden magmas), η\eta is the plastic viscosity, and γ˙\dot{\gamma} is the strain rate. Suspensions of crystals, which can constitute up to 50% by volume near the solidus, significantly increase effective viscosity and impart shear-thinning or yield-stress behavior, hindering flow and promoting plug-like motion in conduits.

Dissolved Gases

Magma contains several key dissolved volatile components that significantly influence its behavior, with water (H₂O) being the most abundant, typically reaching concentrations up to 6 wt% in arc-related magmas, while carbon dioxide (CO₂) ranges from 0.1 to 1 wt%, and lesser amounts of sulfur dioxide (SO₂, derived from sulfur species at ~0.1 wt%), chlorine (Cl at 0.1–0.2 wt%), and fluorine (F at 0.05–0.1 wt%) are also present. The solubility of these gases in silicate melts generally follows Henry's law, expressed as P=KHXP = K_H X, where PP is the partial pressure of the gas, XX is its mole fraction in the melt, and KHK_H is the Henry's constant specific to the gas, melt composition, temperature, and pressure. The speciation of these volatiles within the melt affects their solubility and mobility. Dissolved H₂O exists primarily as hydroxyl groups (OH⁻) at low concentrations and as both hydroxyl and molecular H₂O at higher levels, with the equilibrium governed by the reaction \ceH2O+O2−<=>2OH−\ce{H2O + O^{2-} <=> 2OH^-}. Similarly, CO₂ speciates as molecular CO₂ and ions (CO₃²⁻), where the proportion of carbonate increases with higher oxygen activity and in the melt, while SO₂, Cl, and F dissolve mainly as oxidized or ionic species depending on conditions. As magma ascends and pressure decreases, these dissolved volatiles become supersaturated and exsolve from the melt, initiating vesiculation through the and growth of gas bubbles. This process expands the magma volume, enhances its , and can generate that triggers explosive eruptions when bubble connectivity leads to rapid gas escape. Isotopic compositions of these gases provide tracers for magma origins. The δ¹³C values of dissolved CO₂ typically range from -3 to -8‰ for -derived sources and become more positive (>-5‰) with crustal contamination, while δ¹⁸O in H₂O distinguishes mantle signatures (~5.5‰) from altered crustal fluids (often >6‰ or <5‰). Water's presence also lowers silicate melting temperatures, facilitating flux melting as explored in related sections.

Origins of Magma

Decompression Melting

Decompression melting is a primary mechanism for generating magma in the Earth's mantle, occurring when hot mantle rock ascends rapidly, reducing pressure without substantial conductive heat loss. This ascent follows an adiabatic path, where the temperature decreases along the adiabatic gradient of approximately 0.3–0.5 °C/km, which is comparable to the mantle geotherm but allows crossing of the solidus due to the steeper pressure dependence of the solidus. As pressure drops, the solidus temperature of the mantle peridotite decreases more rapidly than the adiabatic temperature path, causing the rock to intersect its melting curve and undergo partial melting. The process is governed by batch melting, in which the melt remains in equilibrium with the residual solid until extraction. The melt fraction FF is given by the equation F=C0−CsC0−ClF = \frac{C_0 - C_s}{C_0 - C_l} where C0C_0 is the bulk concentration of an element in the initial mantle, CsC_s is the concentration in the solid residue, and ClC_l is the concentration in the initial liquid. For incompatible trace elements, this simplifies to reflect the degree of melting, typically producing basaltic compositions from fertile peridotite sources. This mechanism predominates in tectonic settings involving mantle upwelling, such as mid-ocean ridges and continental rift zones, where divergence thins the lithosphere and facilitates ascent. Partial melting initiates at depths of 50–100 km and generates 5–20% melt volumes, depending on mantle potential temperature and upwelling rate. For instance, at the East Pacific Rise, rapid spreading drives decompression melting of asthenospheric mantle, yielding normal mid-ocean ridge basalt (MORB) magmas with tholeiitic compositions.

Flux Melting

Flux melting, also known as fluid-induced or volatile-induced melting, is a key process for generating magma in convergent tectonic settings, where the addition of water (H₂O) and carbon dioxide (CO₂) to mantle peridotite lowers the solidus temperature, enabling partial melting at conditions otherwise subsolidus. These volatiles act as fluxes by weakening inter-mineral bonds and altering phase equilibria, depressing the solidus of peridotite by 100–300 °C depending on their concentration (typically 0.1–2 wt% H₂O) and pressure; for instance, experiments on fertile lherzolite show a vapor-saturated solidus minimum of ~970 °C at 1.5 GPa, compared to ~1300 °C for the dry solidus. In phase diagrams of the system involving peridotite plus H₂O or CO₂, this manifests as a shift in the eutectic composition toward lower temperatures, where the first melts form at the intersection of the solidus with the volatile-saturated curve, producing low-degree, volatile-rich silicate liquids enriched in incompatible elements. The primary source of these fluxes is the subducting oceanic lithosphere, where hydrous minerals (e.g., amphibole, serpentine) and in the slab dehydrate and decarbonatize during prograde metamorphism at depths of 50–150 km, releasing aqueous fluids and supercritical CO₂-H₂O mixtures into the overlying mantle wedge. In warm-slab subduction zones, up to 90% of slab-bound H₂O may be liberated by ~100 km depth, with CO₂ release occurring via decarbonation reactions in altered oceanic crust and sediments, though CO₂ fluxes are generally lower (∼0.04–0.07 GtC/yr globally) and more depth-dependent. These fluids infiltrate the mantle wedge peridotite, hydrating olivine and pyroxenes to form nominally hydrous phases, which further destabilize the solidus and initiate ; CO₂'s fluxing effect is particularly pronounced at depths >100 km (~3 GPa), where it stabilizes carbonate melts that can hybridize with H₂O-bearing silicates to extend the melting interval. This process predominates in subduction zones, producing intermediate to arc magmas through ~5–15% of wedge, yielding andesitic compositions due to the volatile-rich, slab-influenced source. A representative example is the in the , where flux melting above the subducting Juan de Fuca plate generates hydrous basaltic to andesitic magmas that ascend to form stratovolcanoes like and , with melt volumes estimated at 10–15% based on modeling of erupted lavas. The resulting magmas exhibit elevated volatile contents that enhance their explosivity upon ascent.

Heating-Induced Melting

Heating-induced melting represents a mechanism of magma generation where thermal energy elevates the temperature of solid rock above its solidus, triggering primarily through rather than pressure reduction or volatile . This involves conductive or convective heating from adjacent hot bodies, such as mafic intrusions into the crust or ascending mantle plumes that warm surrounding . In the case of plumes, the rising hot material follows an adiabatic path during ascent, maintaining a nearly constant potential temperature while the actual temperature adjusts along the adiabat. Such melting is characteristic of intraplate environments, including hotspots and the initiation of large igneous provinces, where anomalous mantle heat flux penetrates otherwise stable lithosphere. For instance, the basaltic magmas forming the originate from induced by the hotspot plume, which delivers excess heat to the overlying and generates voluminous melts over millions of years. Effective requires an excess of more than 200°C above the solidus to provide sufficient energy for the of fusion, with plume excesses typically ranging from 200–300°C relative to ambient mantle conditions. The rate of is described by conductive heat flow following Fourier's law: q=−k∇Tq = -k \nabla T where qq is the (W/m²), kk is the thermal conductivity of the rock (typically 2–4 W/m·K for mantle ), and ∇T\nabla T is the . Compared to other melting mechanisms, heating-induced partial melting is less efficient, often producing only 1–5% melt fractions due to the gradual nature of thermal diffusion and limited volume of rock affected by the heat source. This results in smaller magma volumes unless sustained by large-scale plume activity. The resulting melts may exhibit compositional variations influenced by the source rock, such as enrichment in incompatible elements from low-degree melting.

Evolution of Magma

Primary Magmas

Primary magmas represent the initial liquids produced directly from the partial melting of mantle or crustal source rocks, formed as equilibrium partial melts that closely preserve the geochemical signatures of their origins without significant subsequent alteration. These melts are distinguished by their primitive compositions, which reflect the undepleted or minimally modified nature of the source material. In petrology, primary magmas are identified as those in chemical equilibrium with the residual mantle peridotite, avoiding contamination or fractionation processes that would alter their elemental ratios. For mantle-derived primary magmas, particularly basaltic types, characteristic features include high magnesium oxide (MgO) contents exceeding 8 wt%, often ranging from 8–20 wt% depending on the degree of and pressure conditions, alongside elevated levels of compatible trace elements such as (Ni > 200 ppm) and (Cr > 400 ppm). These compositions arise because the melts are saturated with and other phases from the source, retaining high concentrations of elements compatible with those minerals. Crustal-derived primary magmas, though less common for basalts, exhibit signatures tied to the specific , such as partial melts of or eclogite, but mantle remains the dominant source for primitive basaltic varieties. Experimental has confirmed these compositions through high-pressure simulations of melting, demonstrating that primary basaltic magmas form under anhydrous or hydrous conditions at depths of 20–100 km, yielding liquids with MgO contents consistent with observed natural samples. Seminal studies using piston-cylinder apparatus have reproduced the liquidus phase relations, showing that tholeiitic primary magmas equilibrate with , orthopyroxene, and in fertile sources. Such experiments underscore the role of mantle as the primary reservoir for generating basaltic melts worldwide. Identification of primary magmas in nature relies on proxies like melt inclusions trapped in phenocrysts and mantle xenoliths entrained in lavas, which encapsulate unmodified melt compositions shielded from crustal interaction. For instance, basalts (MORB) serve as a key example of primary oceanic magmas, with their primitive variants exhibiting high MgO and abundances that match experimentally derived melts from the . These inclusions and xenoliths provide direct windows into the initial melt stage, preserving volatile contents and isotopic ratios indicative of the source.

Magma Differentiation

Magma differentiation encompasses the physicochemical processes that transform primary magmas into a diverse array of derivative compositions, primarily through the separation of crystals from melt, incorporation of surrounding , and blending of compositionally distinct magmas. These mechanisms operate within magmatic reservoirs, leading to the formation of parental magmas that serve as precursors to the suites observed at Earth's surface. Unlike the initial generation of primary magmas from or lower crustal sources, differentiation focuses on post-generation evolution, often resulting in silica enrichment and that control the final erupted or intruded products. Fractional is a dominant in magma differentiation, where early-formed crystals settle or are removed from the melt, progressively altering the residual liquid's composition toward higher silica content. This separation prevents re-equilibration between crystals and melt, driving incompatible element enrichment in the liquid. The quantitative description of trace element behavior during this is given by the Rayleigh fractionation law: CLC0=FD−1\frac{C_L}{C_0} = F^{D-1} where CLC_L is the concentration of an element in the evolving liquid, C0C_0 is the initial concentration, FF is the fraction of melt remaining, and DD is the bulk (the ratio of the element's concentration in the solid to that in the liquid). This model, originally derived from principles and adapted to igneous systems, explains the observed depletions in compatible elements like magnesium and enrichments in incompatibles such as in evolved magmas. For instance, in basaltic systems, and can produce andesitic derivatives, as demonstrated in experimental and natural studies of tholeiitic suites. Assimilation involves the incorporation and dissolution of wallrock into the magma, leading to that modifies both major and compositions, as well as isotopic signatures. This process is particularly significant in continental settings, where magmas interact with pre-existing crust, adding radiogenic isotopes that distinguish contaminated magmas from pristine mantle-derived ones. A key indicator is the strontium isotope 87Sr/86Sr^{87}\mathrm{Sr}/^{86}\mathrm{Sr}, which typically exceeds 0.706 in upper due to long-term decay, contrasting with mantle values below 0.704; assimilation thus shifts magma ratios toward crustal signatures, often coupled with fractional in the assimilation-fractional (AFC) model. Quantitative modeling of AFC shows that even small amounts of assimilated material (e.g., 10-20% by mass) can significantly alter isotopic trends while conserving through simultaneous , as seen in arc plutons where granitic compositions reflect and entrainment of metasedimentary wallrock. Magma mixing occurs when compositionally disparate magmas converge in a chamber, producing hybrid melts through interlayering, , or injection, often evidenced by mingled textures and reversed in phenocrysts. This process generates intermediate compositions not achievable by simple or assimilation alone, contributing to the diversity within volcanic and plutonic complexes. In zoned plutons, such as those in the , injections into reservoirs create gradational boundaries and hybrid zones, with geochemical gradients reflecting diffusive re-equilibration and convective stirring over timescales of thousands to millions of years. Mixing is particularly prominent in subduction zones, where basaltic and andesitic end-members blend to form dacitic hybrids, as reconstructed from enclave populations and isotopic disequilibria. Parental magmas, which evolve from primary compositions through the above differentiation processes, define distinct evolutionary series such as tholeiitic and alkalic, each reflecting specific paths and source influences. Tholeiitic series, common in and settings, start from high-temperature, water-poor basaltic primaries and progress via olivine-plagioclase to silica-saturated derivatives like andesites, maintaining iron enrichment trends. In contrast, alkalic series, prevalent in intraplate like ocean islands, derive from lower-degree melts richer in alkalies and evolve toward peralkaline compositions such as trachytes through clinopyroxene and alkali feldspar , avoiding iron enrichment due to higher water contents that stabilize . These series highlight how initial primary magma variability, briefly referenced from source-derived compositions, interacts with differentiation to produce global igneous diversity.

Migration and Emplacement

Magma Ascent Mechanisms

Magma ascent from its source region to shallower crustal levels or the surface occurs through a combination of transport processes governed by the physical properties of the magma and surrounding rock. These mechanisms include buoyancy-driven rise of diapiric bodies, fracture propagation via hydraulic fracturing, and slower porous flow within crystal-rich mushes. The dominant mechanism depends on factors such as magma , which can range from 10^2 to 10^6 Pa·s for basaltic to rhyolitic compositions, and generated by contrasts or volatile content. Buoyancy-driven rise is a primary mechanism for the ascent of low-viscosity, partially molten through or lower crust, where the contrast between the magma (Δρ typically 50–300 kg/m³) and surrounding solid rock drives upward motion. For spherical diapirs under low conditions, the ascent velocity v is approximated by :
v≈Δρ g r29 ηv \approx \frac{\Delta \rho \, g \, r^2}{9 \, \eta}
where g is , r is the diapir radius (often 1–10 km), and η is the of the surrounding medium (10^{18}–10^{21} Pa·s for mantle rocks). This yields velocities of 0.1–10 cm/yr for kilometer-scale diapirs, enabling slow but persistent rise over geological timescales. Evolved magma compositions with higher can reduce , potentially stalling ascent. Fracture propagation facilitates rapid magma transport through the brittle upper crust via hydraulic fracturing, where overpressurized magma exploits tensile fractures to form dikes or sills. This process requires magma overpressure exceeding the host rock's tensile strength, typically 1–10 MPa for crustal materials, allowing fractures to propagate at speeds of ~0.1–10 m/s (up to tens m/s in some models). Dikes, which are vertical or near-vertical tabular intrusions, enable vertical ascent over tens of kilometers, while sills form horizontal sheets when propagation encounters lithological or stress barriers. Seismic swarms, characterized by migrating hypocenters at rates of 0.3–4.7 km/h (~7–113 km/day), often indicate active dyke propagation, as observed during the 2014 eruption in where swarms traced a ~48 km lateral dyke path, with an average rate of ~3.7 km/day. Porous flow dominates in crystal mushes, where melt (5–20% ) migrates through interconnected pores in a deformable framework at low velocities typically below 1 cm/yr. This compaction-driven process is limited by the mush's permeability (10^{-12}–10^{-15} m²) and viscosity contrasts, resulting in slow extraction rates that sustain long-term magma differentiation without bulk ascent. from ongoing melt production can enhance flow, but velocities remain subdued compared to fracturing. Viscosity strongly influences all mechanisms, as higher values (e.g., >10^5 Pa·s in crystal-laden magmas) impede flow and favor stalling, while low-viscosity melts promote faster ascent. , often 1–5 MPa from or gas accumulation, is critical for initiating fractures but diminishes with distance from the source. These factors interplay to determine whether magma reaches eruptive levels or emplaces as intrusions.

Plutonism

Plutonism refers to the process by which magma is emplaced intrusively into the , where it cools and solidifies slowly to form coarse-grained plutonic rocks without reaching the surface. This occurs after magma has ascended through various mechanisms, typically at depths of 2 to 30 kilometers, allowing for the development of large intrusive bodies that contribute significantly to continental crustal growth. Emplacement styles of plutons vary based on the rheological properties of the surrounding crust, magma viscosity, and tectonic setting. Laccoliths are mushroom-shaped intrusions that form by forceful injection of magma, causing doming of the overlying while remaining concordant at their base. Batholiths, in contrast, are extensive complexes of interconnected plutons, often spanning hundreds of kilometers, formed through multiple episodes of magma injection over millions of years. Two primary mechanisms drive emplacement: , where blocks of host rock are fractured and assimilated into the magma, creating passively, and forceful injection, where the magma's and actively deform and displace the surrounding crust. The slow cooling of plutons occurs over timescales ranging from 10^4 to 10^6 years, depending on the size of the intrusion and the thermal conductivity of the host rock, allowing for the growth of large mineral crystals. This process is primarily governed by conductive heat loss to the surrounding crust, described by the heat diffusion equation: ∂T∂t=κ∇2T\frac{\partial T}{\partial t} = \kappa \nabla^2 T where TT is , tt is time, κ\kappa is the of the rock (typically around 10^{-6} m²/s for crustal materials), and ∇2T\nabla^2 T represents the spatial variation in . The composition of the solidified plutonic rocks reflects the original magma type: magmas produce , characterized by high silica content (>65%) and minerals like and , while magmas yield , with lower silica (<52%) and dominant plagioclase and pyroxene. Surrounding the pluton, a contact metamorphism aureole develops due to the intense heat, altering the host rocks into hornfels or other high-grade assemblages within a few hundred meters to kilometers of the intrusion margin. Notable examples include the Sierra Nevada batholith in California, a vast Mesozoic assemblage of granitic plutons covering over 70,000 km², formed through repeated mafic to felsic intrusions associated with subduction. Plutonism also plays a key role in forming economic ore deposits, such as porphyry copper systems, where volatile-rich magmas exsolve fluids that precipitate copper sulfides around shallow intrusions, as seen in major deposits like those in the Andes.

Volcanism

Volcanism occurs when magma ascends to the Earth's surface, where reduced pressure causes dissolved volatiles to exsolve rapidly, transforming the magma into lava and driving eruptive activity. This process results in the extrusion of molten rock, accompanied by the release of gases and pyroclasts, shaping volcanic landforms such as cones, domes, and flows. The style and intensity of eruptions depend on magma composition, viscosity, and ascent dynamics, ranging from gentle effusive outflows to catastrophic explosions. Eruptions are triggered by mechanisms such as gas overpressure buildup within the magma, where exsolved volatiles create excess pressure exceeding the strength of the surrounding rock, leading to conduit rupture. Dome collapse, often due to gravitational instability or intrusion of fresh magma, can also initiate eruptions by suddenly decompressing the system and releasing trapped gases. The explosivity of these events is quantified using the Volcanic Explosivity Index (VEI), a scale from 0 to 8 that assesses eruption magnitude based on ejecta volume, plume height, and duration, with each unit representing an order-of-magnitude increase in explosivity. For instance, VEI 0-1 eruptions are non-explosive and localized, while VEI 7-8 events, like the 1815 Tambora eruption, eject over 100 km³ of material globally. Flow regimes during volcanism contrast effusive eruptions, characterized by low-viscosity basaltic magma flowing steadily from fissures or vents, with explosive eruptions involving high-viscosity silicic magma that fragments into pyroclasts. Effusive styles, such as Hawaiian eruptions, produce fluid lava fountains and flows with minimal fragmentation, while explosive Plinian eruptions generate towering ash columns through rapid decompression. Conduit models describe magma ascent where fragmentation occurs when gas overpressure reaches a threshold, typically at shallow depths of a few kilometers corresponding to pressures of 10-100 MPa, transitioning the flow from bubbly magma to a gas-pyroclast mixture. Lava rheology further influences surface morphology: pahoehoe forms smooth, ropy surfaces from low-shear-rate flows of less viscous lava, whereas 'a'ā develops rough, spiny textures due to higher shear stresses that break the crust, often at faster flow rates exceeding 5-10 m³/s. Representative examples illustrate these regimes: the ongoing effusive activity at Kīlauea Volcano in Hawai'i, since 1983, has produced extensive pāhoehoe and 'a'ā flows from low-viscosity basalt, covering approximately 144 km² (as of the end of the main 1983–2018 phase). In contrast, the 1991 Mount Pinatubo eruption in the Philippines was highly explosive (VEI 6), ejecting 10 km³ of dacitic magma and forming Plinian columns up to 35 km high, driven by viscous magma and rapid gas release. Volcanic hazards from these processes include pyroclastic flows—superheated avalanches of gas, ash, and rock fragments traveling at 100-700 km/h, incinerating everything in their path—and lahars, mudflows formed when pyroclastic debris mixes with water from crater lakes or melted snow, capable of burying communities kilometers away. Monitoring relies on sulfur dioxide (SO₂) emissions, measured via ground-based spectrometers or satellite instruments like NASA's Ozone Monitoring Instrument, as elevated SO₂ fluxes (e.g., >500 tons/day) signal magma ascent and potential unrest.

Applications and Study

Energy Production

Geothermal energy production primarily relies on hydrothermal systems, where hot reservoirs are heated by underlying magma bodies, enabling the extraction of for . These systems involve circulating through fractured rocks to capture conducted from magma chambers, which can reach temperatures exceeding 300°C. A prominent example is Iceland's Hellisheiði Power Plant, which utilizes a high-temperature hydrothermal heated by the Hengill volcanic system and produces 303 megawatts of electrical power, contributing significantly to the country's mix. Enhanced geothermal systems (EGS) extend this approach by drilling into hot dry rock formations lacking natural permeability, often influenced by nearby magma sources, and injecting to create artificial fractures for exchange. In EGS, is pumped into the reservoir, heated by conduction from the hot rock, and returned to the surface, where the drives turbines; the extracted follows Q=mcΔTQ = m c \Delta T, with QQ as , mm as , cc as , and ΔT\Delta T as temperature difference between injection and production fluids. This method allows access to deeper, magma-proximate resources, potentially increasing viable sites beyond conventional hydrothermal areas. Conceptual advancements in magma involve direct tapping of molten magma chambers for ultra-high-temperature heat, offering baseload power with immense . Projects like the Magma Testbed in aim to into a shallow magma body at approximately 2 kilometers depth to instrument and extract heat from supercritical fluids, building on prior accidental encounters such as the 2005 Puna Geothermal Venture in , which intersected magma at 2.5 kilometers and demonstrated potential but required immediate cessation due to extreme conditions. Key challenges include stability in temperatures over 900°C, where rock creep and can cause collapse, necessitating and cooling techniques during . As of late , global geothermal installed capacity stands at approximately 16.9 gigawatts, primarily from hydrothermal and early EGS plants, with technical potential estimated to exceed 200 gigawatts through expanded magma-influenced systems. These technologies yield environmental benefits, including lifecycle emissions of about 38 grams of CO₂ per —99% lower than plants—and minimal compounds, reducing and supporting climate mitigation without significant land disruption. Magma-heated geothermal systems also enable direct-use applications beyond electricity generation, such as , industrial drying, and emerging mineral extraction from brines (e.g., ), with global direct-use capacity exceeding 107 GWth as of 2023.

Scientific Investigation Methods

Scientific investigation of magma is challenged by its inaccessibility beneath the Earth's surface, necessitating indirect methods that integrate geophysical, geochemical, and experimental approaches to infer subsurface dynamics and compositions. These techniques address gaps in direct observation by detecting anomalies associated with molten rock, such as perturbations or volatile signatures preserved in erupted materials. Geophysical methods play a central role in imaging magma chambers and tracking surface manifestations of subsurface activity. , particularly through local earthquake , reveals low-velocity zones indicative of partial melts or magma reservoirs, as these regions attenuate seismic waves due to reduced rigidity. For instance, three-dimensional velocity models at Mount Etna have delineated Vp/Vs anomalies linked to active magmatic systems, enabling the mapping of eruptible magma storage. (InSAR) complements by measuring ground deformation caused by magma intrusion or pressure changes, with millimeter-scale precision over large areas. At , continuous InSAR monitoring using satellites like has captured episodic uplift and tied to magma recharge, providing real-time insights into caldera dynamics. Geochemical analyses focus on erupted products to reconstruct pre-eruptive conditions within magma systems. Melt inclusion studies, involving microscopic pockets of trapped melt within crystals, preserve volatile contents and compositions that reflect storage depths, temperatures, and degassing histories before eruption. Recent three-dimensional analyses of inclusions from volcanic samples have quantified gas-rich signatures, revealing rapid ascent rates and eruption triggers. Uranium-thorium (U-Th) dating of accessory minerals like zircon provides geochronological constraints on crystallization timescales, distinguishing between protracted differentiation and rapid events in young magmatic systems. This method has resolved storage durations of less than 10,000 years in rift-related volcanism, highlighting the efficiency of magma evolution. Experimental simulates magmatic processes under controlled conditions to validate field observations. High-pressure experiments, using piston-cylinder apparatuses, replicate crustal depths (up to 2-3 GPa) and temperatures (800-1200°C) to study phase equilibria, , and mixing in magmas. These simulations have demonstrated how hybrid compositions form during intrusions, informing interpretations of zoned plutons. Numerical modeling extends this by solving equations for patterns in cooling chambers, incorporating and compositional gradients. Such models predict double-diffusive leading to layering, as seen in simulations of basaltic systems where Rayleigh numbers exceed 10^6, establishing scales for chamber overturn. Post-2020 advances have integrated technology for enhanced forecasting and sampling, bridging observational gaps. , particularly Bayesian networks and on multi-sensor data, enables probabilistic eruption predictions by identifying precursors like seismic swarms or gas emissions with probabilities exceeding 80% accuracy in retrospective tests. Drone-based sampling has revolutionized volatile collection in hazardous zones, allowing unmanned flights into craters to gather plume gases and , as demonstrated in missions at active volcanoes like Turrialba in since 2021. Real-time satellite monitoring via tracks magma migration through deformation signals, with global datasets processed via to forecast unrest at over 1,000 volcanoes.

References

  1. https://education.nationalgeographic.org/resource/magma/
  2. https://www.usgs.gov/faqs/what-difference-between-magma-and-lava
  3. https://geo.libretexts.org/Bookshelves/Geology/Book%253A_An_Introduction_to_Geology_%28Johnson_Affolter_Inkenbrandt_and_Mosher%29/04%253A_Igneous_Processes_and_Volcanoes/4.03%253A_Magma_Generation
  4. https://opentextbc.ca/geology/chapter/3-2-magma-and-magma-formation/
  5. https://www2.tulane.edu/~sanelson/Natural_Disasters/volcan&magma.htm
  6. https://geo.libretexts.org/Bookshelves/Geology/Physical_Geology_%28Earle%29/04%253A_Volcanism/4.02%253A_Magma_Composition_and_Eruption_Style
  7. https://education.nationalgeographic.org/resource/magma-role-rock-cycle/
  8. https://opentextbc.ca/physicalgeology2ed/chapter/3-2-magma-and-magma-formation/
  9. https://opengeology.org/petrology/03-magma/
  10. https://courses.lumenlearning.com/suny-geophysical/chapter/magma-composition/
  11. https://www.alexstrekeisen.it/english/vulc/index.php
  12. https://comptes-rendus.academie-sciences.fr/geoscience/articles/10.5802/crgeos.101/
  13. https://pmc.ncbi.nlm.nih.gov/articles/PMC9130134/
  14. https://www.science.org/doi/10.1126/sciadv.adl3127
  15. https://www.mantleplumes.org/WebDocuments/Bowen1922.pdf
  16. https://opentextbc.ca/geology/chapter/3-3-crystallization-of-magma/
  17. https://www2.tulane.edu/~sanelson/eens1110/igneous.htm
  18. https://volcanoes.sdsu.edu/Controls.html
  19. https://www.usgs.gov/observatories/calvo/news/how-hot-hot-when-it-comes-volcanoes
  20. https://magma.geol.ucsb.edu/papers/Lesher%2CSpera_EoV.pdf
  21. https://www.eoas.ubc.ca/~krussell/VISCOSITY/grdViscosity_files/grdViscosityPaper.pdf
  22. https://www.jstage.jst.go.jp/article/jmps/105/6/105_100611/_pdf/-char/en
  23. https://agupubs.onlinelibrary.wiley.com/doi/full/10.1029/2008GC002138
  24. https://www.science.org/doi/10.1126/science.abm5174
  25. https://pubs.usgs.gov/pp/1750/chapters/pp2008-1750_chapter27.pdf
  26. https://pubs.geoscienceworld.org/msa/rimg/article/69/1/333/289113/Interpreting-H2O-and-CO2-Contents-in-Melt
  27. https://agupubs.onlinelibrary.wiley.com/doi/full/10.1029/2006RG000216
  28. https://www.jstage.jst.go.jp/article/jmps/118/1/118_221224b/_html/-char/en
  29. https://www.sciencedirect.com/science/article/abs/pii/S0016703701008729
  30. https://agupubs.onlinelibrary.wiley.com/doi/full/10.1029/2004JB003113
  31. https://www.sciencedirect.com/science/article/pii/S0377027320302055
  32. https://www.mdpi.com/2076-3263/15/7/266
  33. https://agupubs.onlinelibrary.wiley.com/doi/10.1029/2023GL107474
  34. https://www.e-education.psu.edu/rocco/node/1988
  35. https://agupubs.onlinelibrary.wiley.com/doi/full/10.1029/2021JB023562
  36. https://www2.tulane.edu/~sanelson/eens212/earths_interior.htm
  37. https://www.sciencedirect.com/topics/earth-and-planetary-sciences/mid-ocean-ridge-basalt
  38. https://www.nps.gov/para/learn/nature/magma-melts-and-eruption-types.htm
  39. https://www2.tulane.edu/~sanelson/eens1110/igneous.pdf
  40. https://www.lyellcollection.org/doi/full/10.1144/sp410.13
  41. https://academic.oup.com/petrology/article/55/10/2067/1455622
  42. https://www.sciencedirect.com/science/article/abs/pii/S0012821X06004869
  43. https://agupubs.onlinelibrary.wiley.com/doi/full/10.1029/2002JB001839
  44. https://www.pnas.org/doi/10.1073/pnas.1507889112
  45. https://pubs.usgs.gov/publication/70159878
  46. https://www.researchgate.net/publication/223075677_Mantle_plumes
  47. https://www.sciencedirect.com/science/article/pii/0012821X9500122S
  48. https://agupubs.onlinelibrary.wiley.com/doi/10.1029/2005GC000927
  49. https://www.nature.com/articles/nature06376
  50. https://essopenarchive.org/users/539634/articles/1182545/master/file/data/999966_0_merged_1719326996/999966_0_merged_1719326996.pdf
  51. http://jupiter.ethz.ch/~gfdteaching/dymali/2022/downloads/ThermalStructure-Notes.pdf
  52. https://pubs.geoscienceworld.org/gsa/gsabulletin/article-abstract/100/7/1040/182164/Possible-petrogenetic-relations-between-low-and
  53. https://www.sciencedirect.com/science/article/abs/pii/002449377990063X
  54. https://www.frontiersin.org/journals/earth-science/articles/10.3389/feart.2018.00114/full
  55. https://academic.oup.com/petrology/article-abstract/4/1/75/1500186
  56. https://agupubs.onlinelibrary.wiley.com/doi/abs/10.1029/91jb02840
  57. https://pubs.geoscienceworld.org/gsa/geology/article/28/11/991/191020/Using-melt-inclusions-to-determine-parent-magma
  58. https://agupubs.onlinelibrary.wiley.com/doi/10.1029/91JB02841
  59. https://www.sciencedirect.com/science/article/abs/pii/S0024493712000461
  60. https://wwwrcamnl.wr.usgs.gov/isoig/isopubs/itchch2.html
  61. https://agupubs.onlinelibrary.wiley.com/doi/abs/10.1029/91JB02150
  62. https://pubs.usgs.gov/pp/1320/report.pdf
  63. https://link.springer.com/article/10.1007/s00410-019-1576-4
  64. https://www.geochemicalperspectivesletters.org/article2125/
  65. https://www.nature.com/articles/s41467-025-61620-5
  66. https://www.science.org/doi/10.1126/science.adn2838
  67. https://se.copernicus.org/articles/12/1549/2021/
  68. https://agupubs.onlinelibrary.wiley.com/doi/full/10.1029/2023GC011115
  69. https://users.monash.edu.au/~weinberg/PDF_Papers/Weinberg_Podladchikov_1994_JGR.pdf
  70. https://gonnermann.rice.edu/publications/gonnermann-2015-magma-trans.pdf
  71. https://academic.oup.com/gji/article/236/3/1863/7577610
  72. https://www.sciencedirect.com/science/article/pii/S0012821X18301420
  73. https://pmc.ncbi.nlm.nih.gov/articles/PMC6335481/
  74. https://royalsocietypublishing.org/doi/10.1098/rsta.2018.0019
  75. https://seismo.berkeley.edu/wiki_cider/images/3/32/Gonnerman_cider_2017.pdf
  76. https://pubs.usgs.gov/publication/pp1483
  77. http://profharwood.x10host.com/GEOL101/Study/Plutons.htm
  78. https://agupubs.onlinelibrary.wiley.com/doi/10.1029/2012JB009325
  79. https://www.sciencedirect.com/science/article/abs/pii/S0191814116301390
  80. https://timslab.princeton.edu/document/144
  81. https://www.researchgate.net/publication/258662643_Insights_into_the_mechanisms_and_timescales_of_pluton_assembly_from_deformation_patterns_of_mafic_enclaves
  82. https://pmc.ncbi.nlm.nih.gov/articles/PMC9506720/
  83. https://www.nps.gov/subjects/geology/igneous.htm
  84. https://geo.libretexts.org/Bookshelves/Geology/Book%3A_An_Introduction_to_Geology_%28Johnson_Affolter_Inkenbrandt_and_Mosher%29/04%3A_Igneous_Processes_and_Volcanoes/4.01%3A_Classification_of_Igneous_Rocks
  85. https://opengeology.org/textbook/6-metamorphic-rocks/
  86. https://pubs.usgs.gov/sir/2010/5070/b/pdf/SIR10-5070B.pdf
  87. https://www.nature.com/articles/srep40566
  88. https://www.usgs.gov/programs/VHP/volcanic-gases-can-be-harmful-health-vegetation-and-infrastructure
  89. https://www.sciencedirect.com/science/article/abs/pii/S0012821X98002775
  90. https://www.sciencedirect.com/science/article/pii/S0377027325000629
  91. https://www.usgs.gov/observatories/yvo/news/volcanic-explosivity-index-a-tool-comparing-sizes-explosive-volcanic
  92. https://pubs.usgs.gov/pp/1801/downloads/pp1801_Chap8_Mangan.pdf
  93. https://www.nature.com/articles/s41467-018-05293-3
  94. https://www.usgs.gov/volcanoes/kilauea/science/puuoo-eruption-lasted-35-years
  95. https://pubs.usgs.gov/pp/1801/downloads/pp1801_Chap10_Tilling.pdf
  96. https://www.usgs.gov/news/volcano-watch-how-big-eruption
  97. https://www.usgs.gov/programs/VHP/pyroclastic-flows-move-fast-and-destroy-everything-their-path
  98. https://www.usgs.gov/programs/VHP/lahars-move-rapidly-down-valleys-rivers-concrete
  99. https://www.usgs.gov/faqs/how-are-volcanic-gases-measured
  100. https://www.mdpi.com/2076-3263/10/6/212
  101. https://www.spe-aberdeen.org/uploads/20250227-Geothermal-and-CCUS-in-Iceland.pdf
  102. https://www.energy.gov/eere/geothermal/enhanced-geothermal-systems
  103. https://www.sciencedirect.com/science/article/pii/S0375650516300712
  104. https://kmt.is/
  105. https://www.usgs.gov/observatories/hvo/news/volcano-watch-will-drilling-magma-start-a-volcanic-eruption
  106. https://link.springer.com/article/10.1007/s44421-025-00004-3
  107. https://www.thinkgeoenergy.com/thinkgeoenergys-top-10-geothermal-countries-2024-power/
  108. https://www.iea.org/reports/the-future-of-geothermal-energy/executive-summary
  109. https://www.nrel.gov/docs/fy15osti/57187.pdf
  110. https://www.irena.org/publications/2023/Jan/Geothermal-Power-Technology-Brief
  111. https://www.nature.com/articles/s41598-024-74349-w
  112. https://www.usgs.gov/volcanoes/yellowstone/science/monitoring-deformation-yellowstone-national-park
  113. https://agupubs.onlinelibrary.wiley.com/doi/10.1029/2021JB022341
  114. https://academic.oup.com/petrology/article/66/3/egaf012/8003508
  115. https://www.annualreviews.org/content/journals/10.1146/annurev-earth-040610-133330
  116. https://www.usgs.gov/observatories/yvo/news/under-pressure-and-temperature-volcano-lessons-learned-laboratory
  117. https://academic.oup.com/petrology/article/51/3/731/1535211
  118. https://agupubs.onlinelibrary.wiley.com/doi/full/10.1029/2024GL112029
  119. https://www.gi.alaska.edu/news/uaf-drones-sample-active-costa-rica-volcano
  120. https://sentinels.copernicus.eu/web/success-stories/-/copernicus-sentinel-1-to-help-create-global-volcano-monitoring-system
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