Magma: Difference between revisions
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==Origins of magma by partial melting== |
==Origins of magma by partial melting== |
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It is usually very difficult to change the bulk composition of a large mass of rock, so composition is the basic control on whether a rock will melt at any given temperature and pressure. The composition of a rock may also be considered to include ''volatile'' phases such as [[water]] and [[carbon dioxide]]. |
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===Partial melting=== |
===Partial melting=== |
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Melting of solid rocks to form magma is controlled by three physical parameters: temperature, pressure, and composition. The most common mechanisms of magma generation in the mantle are [[Igneous rock#Decompression|decompression melting]],<ref>Geological Society of America, ''Plates, Plumes, And Paradigms'', pp. 590 ff., 2005, {{ISBN|0-8137-2388-4}}</ref> heating ( |
Melting of solid rocks to form magma is controlled by three physical parameters: temperature, pressure, and composition. The most common mechanisms of magma generation in the mantle are [[Igneous rock#Decompression|decompression melting]],<ref>Geological Society of America, ''Plates, Plumes, And Paradigms'', pp. 590 ff., 2005, {{ISBN|0-8137-2388-4}}</ref> heating (such as by interaction with a hot [[mantle plume]]<ref>{{Cite journal|last=Campbell|first=I. H.|date=2005-12-01|title=Large Igneous Provinces and the Mantle Plume Hypothesis|journal=Elements|volume=1|issue=5|pages=265–269|doi=10.2113/gselements.1.5.265|issn=1811-5209}}</ref>), and lowering of the [[Solidus (chemistry)|solidus]] (such as by the addition of water<ref>{{Cite journal|last1=Asimow|first1=P. D.|last2=Langmuir|first2=C. H.|date=2003|title=The importance of water to oceanic mantle melting regimes|journal=Nature|volume=421|issue=6925|pages=815–820|doi=10.1038/nature01429|pmid=12594505|bibcode=2003Natur.421..815A|s2cid=4342843|issn=0028-0836}}</ref>). Mechanisms are discussed further in the entry for [[igneous rock]]. |
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When rocks melt, they do so slowly and gradually because most rocks are made of several [[minerals]], which all have different melting points; moreover, the physical and chemical relationships controlling the melting are complex. As a rock melts, for example, its volume changes. 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.<ref>{{Cite journal|last=Faul|first=Ulrich H.|date=2001|title=Melt retention and segregation beneath mid-ocean ridges|journal=Nature|volume=410|issue=6831|pages=920–923|doi=10.1038/35073556|pmid=11309614|bibcode=2001Natur.410..920F|s2cid=4403804|issn=0028-0836}}</ref> Melts can stay in place long enough to melt to 20% or even 35%, but rocks are rarely melted in excess of 50%, because eventually the melted rock mass becomes a crystal-and-melt mush that can then ascend ''en masse'' as a [[diapir]], which may then cause further decompression melting. |
When rocks melt, they do so slowly and gradually because most rocks are made of several [[minerals]], which all have different melting points; moreover, the physical and chemical relationships controlling the melting are complex. As a rock melts, for example, its volume changes. 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.<ref>{{Cite journal|last=Faul|first=Ulrich H.|date=2001|title=Melt retention and segregation beneath mid-ocean ridges|journal=Nature|volume=410|issue=6831|pages=920–923|doi=10.1038/35073556|pmid=11309614|bibcode=2001Natur.410..920F|s2cid=4403804|issn=0028-0836}}</ref> Melts can stay in place long enough to melt to 20% or even 35%, but rocks are rarely melted in excess of 50%, because eventually the melted rock mass becomes a crystal-and-melt mush that can then ascend ''en masse'' as a [[diapir]], which may then cause further decompression melting. |
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| ⚫ | Within the solid earth, the temperature of a rock is controlled by the [[geothermal gradient]] and the [[radioactive decay]] within the rock. The geothermal gradient averages about 25 °C/km with a wide range from a low of 5–10 °C/km within oceanic trenches and subduction zones to 30–80 °C/km ' |
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===Geochemical implications of partial melting=== |
===Geochemical implications of partial melting=== |
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Revision as of 01:26, 9 February 2021
Magma (from Ancient Greek μάγμα (mágma) meaning "thick unguent"[1]) is the molten or semi-molten natural material from which all igneous rocks are formed.[2] Magma 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 at various tectonic settings, including 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 their storage in the crust, magma compositions may be modified by fractional crystallization, contamination with crustal melts, magma mixing, and degassing. Following their ascent through the crust, magmas may feed a volcano to be extruded as lava, or solidify underground to form an intrusion,[8] such as a igneous dike or a sill.
While the study of magma has historically relied on observing magma in the form of lava flows, magma has been encountered in situ three times during geothermal drilling projects—twice in Iceland (see Magma usage for energy production), and once in Hawaii.[9][10][11][12]
Physical and chemical properties of magma
Magma consists of liquid in which there are usually suspended solid crystals.[13] As magma approaches the surface, and the overburden pressure drops, dissolved gases begin to separate from the liquid as bubbles, so that a magma near the surface consists of both solid, liquid, and gas phases.[14]
Composition
Most magmatic liquids are rich in silica.[8] Rare nonsilicate magmas can form by local melting of nonsilicate mineral deposits[15] or by separation of a magma into separate immiscible silicate and nonsilicate liquid phases.[16]
Silicate magmas are molten mixtures dominated by oxygen and silicon, the Earth's most abundant chemical elements, with smaller quantities of aluminium, calcium, magnesium, iron, sodium, and potassium, and minor amounts of many other elements.[17] 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.[18]
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.[19]
Felsic magma
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 for hot rhyolite magma at 1,200 °C (2,190 °F) to 1011 cP for cool rhyolite magma at 800 °C (1,470 °F).[20] For comparison, water has a viscosity of about 1 cP. 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).[21] The lavas typically fragment as they extrude, producing block lava flows. These often contain obsidian.[22]
Felsic lavas can erupt at temperatures as low as 800 °C (1,470 °F).[23] 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.[24]
Intermediate magma
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.[25] 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 × 106 cP at 1,200 °C (2,190 °F). This is slightly greater than the viscosity of smooth peanut butter.[26] Intermediate magmas show a greater tendency to form phenocrysts,[27] Higher iron and magnesium tends to manifest as a darker groundmass, including amphibole or pyroxene phenocrysts.[28]
Mafic magmas
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, although this is still many orders of magnitude higher than water. This viscosity is similar to that of ketchup.[29] 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.[30] 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.[31]
Ultramafic magmas
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.[32] Viscosities of komatiite magmas are thought to have been as low as 100 to 1000 cP, similar to that of light motor oil.[20] 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.[33]
Akaline magmas
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.[34] 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.[35] 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.[36]
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Tholeiitic basalt magma SiO2 (53.8%) Al2O3 (13.9%) FeO (9.3%) CaO (7.9%) MgO (4.1%) Na2O (3.0%) Fe2O3 (2.6%) TiO2 (2.0%) K2O (1.5%) P2O5 (0.4%) MnO (0.2%)
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Rhyolite magma SiO2 (73.2%) Al2O3 (14%) FeO (1.7%) CaO (1.3%) MgO (0.4%) Na2O (3.9%) Fe2O3 (0.6%) TiO2 (0.2%) K2O (4.1%) P2O5 (0.%) MnO (0.%)
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Nonsilicic magmas
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.[38] 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.[39] 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).[40]
- Iron oxide magmas are thought to be the source of the iron ore at Kiruna, Sweden which formed during the Proterozoic.[16] Iron oxide lavas of Pliocene age occur at the El Laco volcanic complex on the Chile-Argentina border.[15] 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.[16]
- 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).[15]
Magmatic gases
The concentrations of different gases can vary considerably. Water vapor is typically the most abundant magmatic gas, followed by carbon dioxide[41] and sulfur dioxide. Other principal magmatic gases include hydrogen sulfide, hydrogen chloride, and hydrogen fluoride.[42]
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.[43] However, magmas are not necessarily saturated under typical conditions.
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Carbon dioxide is much less soluble in magmas than water, and frequently separates into a separate fluid phase even at great depth. This explains the presence of carbon dioxide fluid inclusions in crystals formed in magmas at great depth.[44]
Rheology
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,[23] the viscosity of the same lavas ranges over seven orders of magnitude, from 104 cP for mafic lava to 1011 cP for felsic magmas.[23] The viscosity is mostly determined by composition, but is also dependent on temperature.[20] The tendency for 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.[45]
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A single silica tetrahedron -
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.[14]
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.[46] 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.[47] 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.[48] 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.[49]
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.[50]
Temperature
Temperatures of lavas are in the range 700 °C to 1300 °C (or 1300 °F to 2400 °F), but very rare carbonatite magmas may be as cool as 490 °C,[51] and komatiite magmas may have been as hot as 1600 °C.[52] These are temperatures of magma that has been extruded at the surface. Magmas have occasionally been encountered during drilling in geothermal field, 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 1050 °C (1922 °F). Temperatures of deeper magmas must be inferred from theoretical computations and the geothermal gradient.[12] under mid-ocean ridges and volcanic arc environments.
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.[13] Calculations of solidus temperatures at likely depths suggests that magma generated beneath areas of rifting starts at a temperature of about 1300 °C to 1500 °C. Magma generated from mantle plumes may be as hot as 1600 °C. The temperature of magma generated in subduction zones, where water vapor lowers the melting temperature, may be as low as 1060 °C.[53]
Density
Magma densities depend mostly on composition, with the iron content being the most important parameter. Magmas also expand slightly at lower pressure and higher temperature.
| Type | Density (kg/m3) |
|---|---|
| Basalt magma | 2650–2800[54] |
| Andesite magma | 2450–2500[54] |
| Rhyolite magma | 2180–2250[54] |
As magmas approach the surface, the dissolved gases in the magma begin to bubble out of the liquid. These bubbles significantly reduce the density of the magma and help drive it further towards the surface.[55]
Origins of magma by partial melting
Partial melting
Melting of solid rocks to form magma is controlled by three physical parameters: temperature, pressure, and composition. The most common mechanisms of magma generation in the mantle are decompression melting,[56] heating (such as by interaction with a hot mantle plume[57]), and lowering of the solidus (such as by the addition of water[58]). Mechanisms are discussed further in the entry for igneous rock.
When rocks melt, they do so slowly and gradually because most rocks are made of several minerals, which all have different melting points; moreover, the physical and chemical relationships controlling the melting are complex. As a rock melts, for example, its volume changes. 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.[59] Melts can stay in place long enough to melt to 20% or even 35%, but rocks are rarely melted in excess of 50%, because eventually the melted rock mass becomes a crystal-and-melt mush that can then ascend en masse as a diapir, which may then cause further decompression melting.
Within the solid earth, the temperature of a rock is controlled by the geothermal gradient and the radioactive decay within the rock. The geothermal gradient averages about 25 °C/km with a wide range from a low of 5–10 °C/km within oceanic trenches and subduction zones to 30–80 °C/km '
The presence of volatile phases in a rock under pressure can stabilize a melt fraction. The presence of even 0.8% water may reduce the temperature of melting by as much as 100 °C. Conversely, the loss of water and volatiles from a magma may cause it to essentially freeze or solidify.
Geochemical implications of partial melting
The degree of partial melting is critical to determination of the characteristics of the magma it produces, and the likelihood that a melt forms reflects the degrees to which incompatible and compatible elements are involved. Incompatible elements commonly include potassium, barium, caesium, and rubidium.
Rock types produced by small degrees of partial melting in the Earth's mantle are typically alkaline (Ca, Na), potassic (K) or peralkaline (in which the aluminium to silica ratio is high). Typically, primitive melts of this composition form lamprophyre, lamproite, kimberlite and sometimes nepheline-bearing mafic rocks such as alkali basalts and essexite gabbros or even carbonatite.
Pegmatite may be produced by low degrees of partial melting of the crust. 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. At high degrees of partial melting of the crust, granitoids such as tonalite, granodiorite and monzonite can be produced, but other mechanisms are typically important in producing them.
Evolution of magmas
Primary melts
When rock melts, the liquid is a primary melt. Primary melts have not undergone any differentiation and represent the starting composition of a magma. In nature it is rare to find primary melts. The leucosomes of migmatites are examples of primary melts. Primary melts derived from the mantle are especially important, and are known as primitive melts or primitive magmas. By finding the primitive magma composition of a magma series it is possible to model the composition of the mantle from which a melt was formed, which is important in understanding evolution of the mantle.[clarification needed]
Parental melts
When it is impossible to find the primitive or primary magma composition, it is often useful[according to whom?] to attempt to identify a parental melt. A parental melt 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.
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 melt. Fractional crystallization models would be produced to test the hypothesis that they share a common parental melt.
At high degrees of partial melting of the mantle, komatiite and picrite are produced.
Migration and solidification of magmas
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. 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 cools and crystallizes forming igneous rock, it erupts as a volcano, or moves into another magma chamber.There are two known processes by which magma changes: by crystallization within the crust or mantle to form a pluton, or by volcanic eruption to become lava or tephra.
Plutonism
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).
Volcanism
During a volcanic eruption the magma that leaves the underground 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.
Magma usage for energy production
The Iceland Deep Drilling Project, while drilling several 5,000m holes in an attempt to harness the heat in the volcanic bedrock below the surface of Iceland, struck a pocket of magma at 2,100m 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.
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 36MW of power, making IDDP-1 the world's first magma-enhanced geothermal system.[60]
References
- ^ "Definition of Magma". Merriam-Webster Dictionary. Merriam-Webster. Retrieved 28 October 2018.
- ^ BOWEN, NORMAN L. (1947). "MAGMAS". Geological Society of America Bulletin. 58 (4): 263. doi:10.1130/0016-7606(1947)58[263:M]2.0.CO;2. ISSN 0016-7606.
- ^ Greeley, Ronald; Schneid, Byron D. (1991-11-15). "Magma Generation on Mars: Amounts, Rates, and Comparisons with Earth, Moon, and Venus". Science. 254 (5034): 996–998. Bibcode:1991Sci...254..996G. doi:10.1126/science.254.5034.996. ISSN 0036-8075. PMID 17731523. S2CID 206574665.
- ^ Spera, Frank J. (2000), "Physical Properties of Magma", in Sigurdsson, Haraldur (editor-in-chief) (ed.), Encyclopedia of Volcanoes, Academic Press, pp. 171–190, ISBN 978-0126431407
{{citation}}:|editor-first=has generic name (help) - ^ Foulger, G.R. (2010). Plates vs. Plumes: A Geological Controversy. Wiley–Blackwell. ISBN 978-1-4051-6148-0.
- ^ Detrick, R. S.; Buhl, P.; Vera, E.; Mutter, J.; Orcutt, J.; Madsen, J.; Brocher, T. (1987). "Multi-channel seismic imaging of a crustal magma chamber along the East Pacific Rise". Nature. 326 (6108): 35–41. Bibcode:1987Natur.326...35D. doi:10.1038/326035a0. ISSN 0028-0836. S2CID 4311642.
- ^ a b Sparks, R. Stephen J.; Cashman, Katharine V. (2017). "Dynamic Magma Systems: Implications for Forecasting Volcanic Activity". Elements. 13 (1): 35–40. doi:10.2113/gselements.13.1.35. ISSN 1811-5209.
- ^ a b MCBIRNEY, A. R.; NOYES, R. M. (1979-08-01). "Crystallization and Layering of the Skaergaard Intrusion". Journal of Petrology. 20 (3): 487–554. Bibcode:1979JPet...20..487M. doi:10.1093/petrology/20.3.487. ISSN 0022-3530.
- ^ Scientists' Drill Hits Magma: Only Third Time on Record, UC Davis News and Information, June 26, 2009.
- ^ Magma Discovered in Situ for First Time. Physorg (December 16, 2008)
- ^ Puna Dacite Magma at Kilauea: Unexpected Drilling Into an Active Magma Posters, 2008 Eos Trans. AGU, 89(53), Fall Meeting.
- ^ a b Teplow, William; Marsh, Bruce; Hulen, Jeff; Spielman, Paul; Kaleikini, Mike; Fitch, David; Rickard, William (2009). "Dacite Melt at the Puna Geothermal Venture Wellfield, Big Island of Hawaii" (PDF). GRC Transactions. 33: 989–994. Retrieved 8 February 2021.
- ^ a b Philpotts, Anthony R.; Ague, Jay J. (2009). Principles of igneous and metamorphic petrology (2nd ed.). Cambridge, UK: Cambridge University Press. pp. 19–20. ISBN 9780521880060.
- ^ a b Schmincke, Hans-Ulrich (2003). Volcanism. Berlin: Springer. p. 49-50. ISBN 9783540436508.
- ^ a b c Guijón, R.; Henríquez, F.; Naranjo, J.A. (2011). "Geological, Geographical and Legal Considerations for the Conservation of Unique Iron Oxide and Sulphur Flows at El Laco and Lastarria Volcanic Complexes, Central Andes, Northern Chile". Geoheritage. 3 (4): 99–315. doi:10.1007/s12371-011-0045-x. S2CID 129179725.
- ^ a b c Harlov, D.E.; et al. (2002). "Apatite–monazite relations in the Kiirunavaara magnetite–apatite ore, northern Sweden". Chemical Geology. 191 (1–3): 47–72. Bibcode:2002ChGeo.191...47H. doi:10.1016/s0009-2541(02)00148-1.
- ^ Philpotts & Ague 2009, p. 19.
- ^ Philpotts & Ague 2009, pp. 132–133.
- ^ Casq, R.A.F.; Wright, J.V. (1987). Volcanic Successions. Unwin Hyman Inc. p. 528. ISBN 978-0-04-552022-0.
- ^ a b c Philpotts & Ague 2009, p. 23.
- ^ Philpotts & Ague 2009, pp. 70–77.
- ^ Schmincke 2003, p. 132.
- ^ a b c Philpotts & Ague 2009, p. 20.
- ^ Bonnichsen, B.; Kauffman, D.F. (1987). "Physical features of rhyolite lava flows in the Snake River Plain volcanic province, southwestern Idaho". Geological Society of America Special Paper. Geological Society of America Special Papers. 212: 119–145. doi:10.1130/SPE212-p119. ISBN 0-8137-2212-8.
- ^ Schmincke 2003, pp. 21–24, 132, 143.
- ^ Philpotts & Ague 2009, pp. 23–611.
- ^ Takeuchi, Shingo (5 October 2011). "Preeruptive magma viscosity: An important measure of magma eruptibility". Journal of Geophysical Research. 116 (B10): B10201. Bibcode:2011JGRB..11610201T. doi:10.1029/2011JB008243.
- ^ Philpotts & Ague 2009, pp. 1376–377.
- ^ Philpotts & Ague 2009, pp. 23–25.
- ^ Philpotts & Ague 2009, p. 53-55, 59-64.
- ^ Schmincke 2003, pp. 128–132.
- ^ Arndt, N.T. (1994). "Archean komatiites". In Condie, K.C. (ed.). Archean Crustal Evolution. Amsterdam: Elsevier. p. 19. ISBN 978-0-444-81621-4.
- ^ Philpotts & Ague 2009, pp. 399–400.
- ^ Philpotts & Ague 2009, pp. 139–148.
- ^ Philpotts & Ague 2009, pp. 606–607.
- ^ "Stikine Volcanic Belt: Volcano Mountain". Catalogue of Canadian volcanoes. Archived from the original on 2009-03-07. Retrieved 23 November 2007.
- ^ Philpotts & Ague 2009, p. 145.
- ^ Vic Camp, How volcanoes work, Unusual Lava Types, San Diego State University, Geology
- ^ Philpotts & Ague 2009, pp. 396–397.
- ^ Keller, Jörg; Krafft, Maurice (November 1990). "Effusive natrocarbonatite activity of Oldoinyo Lengai, June 1988". Bulletin of Volcanology. 52 (8): 629–645. Bibcode:1990BVol...52..629K. doi:10.1007/BF00301213. S2CID 129106033.
- ^ Pedone, M.; Aiuppa, A.; Giudice, G.; Grassa, F.; Francofonte, V.; Bergsson, B.; Ilyinskaya, E. (2014). "Tunable diode laser measurements of hydrothermal/volcanic CO2 and implications for the global CO2 budget". Solid Earth. 5 (2): 1209–1221. Bibcode:2014SolE....5.1209P. doi:10.5194/se-5-1209-2014.
- ^ Schmincke 2003, p. 42.
- ^ Philpotts & Ague 2009, pp. 244–250.
- ^ a b Schmincke 2003, p. 44.
- ^ Schmincke 2003, pp. 38–41.
- ^ Pinkerton, H.; Bagdassarov, N. (2004). "Transient phenomena in vesicular lava flows based on laboratory experiments with analogue materials". Journal of Volcanology and Geothermal Research. 132 (2–3): 115–136. Bibcode:2004JVGR..132..115B. doi:10.1016/s0377-0273(03)00341-x.
- ^ Schmincke 2003, pp. 39–40.
- ^ Philpotts & Ague 2009, p. 40.
- ^ Philpotts & Ague 2009, p. 16.
- ^ Wadsworth, Fabian B.; Witcher, Taylor; Vossen, Caron E. J.; Hess, Kai-Uwe; Unwin, Holly E.; Scheu, Bettina; Castro, Jonathan M.; Dingwell, Donald B. (December 2018). "Combined effusive-explosive silicic volcanism straddles the multiphase viscous-to-brittle transition". Nature Communications. 9 (1): 4696. Bibcode:2018NatCo...9.4696W. doi:10.1038/s41467-018-07187-w. ISSN 2041-1723. PMC 6224499. PMID 30409969.
- ^ Weidendorfer, D.; Schmidt, M.W.; Mattsson, H.B. (2017). "A common origin of carbonatite magmas". Geology. 45 (6): 507–510. Bibcode:2017Geo....45..507W. doi:10.1130/G38801.1.
- ^ Herzberg, C.; Asimow, P. D.; Arndt, N.; Niu, Y.; Lesher, C. M.; Fitton, J. G.; Cheadle, M. J.; Saunders, A. D. (2007). "Temperatures in ambient mantle and plumes: Constraints from basalts, picrites, and komatiites". Geochemistry, Geophysics, Geosystems. 8 (2): n/a. Bibcode:2007GGG.....8.2006H. doi:10.1029/2006gc001390. hdl:20.500.11919/1080. ISSN 1525-2027.
- ^ Philpotts & Ague 2009, pp. 593–597.
- ^ a b c usu.edu - Geology 326, "Properties of Magmas", 2005-02-11
- ^ Schmincke 2003, p. 50.
- ^ Geological Society of America, Plates, Plumes, And Paradigms, pp. 590 ff., 2005, ISBN 0-8137-2388-4
- ^ Campbell, I. H. (2005-12-01). "Large Igneous Provinces and the Mantle Plume Hypothesis". Elements. 1 (5): 265–269. doi:10.2113/gselements.1.5.265. ISSN 1811-5209.
- ^ Asimow, P. D.; Langmuir, C. H. (2003). "The importance of water to oceanic mantle melting regimes". Nature. 421 (6925): 815–820. Bibcode:2003Natur.421..815A. doi:10.1038/nature01429. ISSN 0028-0836. PMID 12594505. S2CID 4342843.
- ^ Faul, Ulrich H. (2001). "Melt retention and segregation beneath mid-ocean ridges". Nature. 410 (6831): 920–923. Bibcode:2001Natur.410..920F. doi:10.1038/35073556. ISSN 0028-0836. PMID 11309614. S2CID 4403804.
- ^ Wilfred Allan Elders, Guðmundur Ómar Friðleifsson and Bjarni Pálsson (2014). Geothermics Magazine, Vol. 49 (January 2014). Elsevier Ltd.
