Waldemar Lindgren
Waldemar Lindgren was a Swedish-American geologist. Lindgren was one of the founders of modern economic geology. Waldemar Lindgren was born in Vassmolösa, Kalmar Municipality, in the historical province of Småland in southern Sweden, the son of Johan and Emma Lindgren. Lindgren's father was a member of parliament, his mother the daughter of a clergyman. Lindgren attended the Freiberg Mining Academy, graduating as a mining engineer in 1882. In 1884, he began a 31-year career with the U. S. Geological Survey, working on ore deposits in the Rocky Mountains. In 1905, he helped found the journal Economic Geology. In 1912, he was appointed head of the Department of Geology at the Massachusetts Institute of Technology. Lindgren was elected a foreign member of the Royal Swedish Academy of Sciences in 1931. Lindgren was a fellow of the Mineralogical Society of America, he has served as President of the Geological Society of America in 1924 and of the Society of Economic Geologists. Lindgren's published writings run to nearly 200 titles, not counting discussions and more than 1,000 abstracts.
Most of his publications are on the great ore deposits. Editions of Mineral Deposits, his used textbook were published in 1913, 1919, 1928 and 1933, he died in 1939 in Mass.. The Gold Belt of the Blue Mountains of Oregon Extract from the 22nd Annual Report Part 2: Ore Deposits The water resources of Molokai, Hawaiian Islands The copper deposits of the Clifton-Morenci district, Arizona Geology and gold deposits of the Cripple Creek District, Colorado; the ore deposits of New Mexico The Tertiary Gravels of the Sierra Nevada of California Mineral Deposits Works by or about Waldemar Lindgren at Internet Archive Memorial from American Mineralogist Brief bio, National Mining Hall of Fame Biographical sketch from Dictionary of American Biography Lindgren at Freiberg Mining Academy at the Wayback Machine, from SEG Newsletter, accessed 1/17/08
Geothermal gradient
Geothermal gradient is the rate of increasing temperature with respect to increasing depth in the Earth's interior. Away from tectonic plate boundaries, it is about 25–30 °C/km of depth near the surface in most of the world. Speaking, geo-thermal refers to the Earth but the concept may be applied to other planets; the Earth's internal heat comes from a combination of residual heat from planetary accretion, heat produced through radioactive decay, latent heat from core crystallization, heat from other sources. The major heat-producing isotopes in the Earth are potassium-40, uranium-238, uranium-235, thorium-232. At the center of the planet, the temperature may be up to 7,000 K and the pressure could reach 360 GPa; because much of the heat is provided by radioactive decay, scientists believe that early in Earth history, before isotopes with short half-lives had been depleted, Earth's heat production would have been much higher. Heat production was twice that of present-day at 3 billion years ago, resulting in larger temperature gradients within the Earth, larger rates of mantle convection and plate tectonics, allowing the production of igneous rocks such as komatiites that are no longer formed.
Temperature within the Earth increases with depth. Viscous or molten rock at temperatures between 650 to 1,200 °C are found at the margins of tectonic plates, increasing the geothermal gradient in the vicinity, but only the outer core is postulated to exist in a molten or fluid state, the temperature at the Earth's inner core/outer core boundary, around 3,500 kilometres deep, is estimated to be 5650 ± 600 Kelvin; the heat content of the Earth is 1031 joules. Much of the heat is created by decay of radioactive elements. An estimated 45 to 90 percent of the heat escaping from the Earth originates from radioactive decay of elements located in the mantle. Gravitational potential energy released during the accretion of the Earth. Heat released during differentiation. Latent heat released as the liquid outer core crystallizes at the inner core boundary. Heat may be generated by tidal forces on the Earth; the resulting earth tides dissipate energy in Earth's interior as heat. There is no reputable science to suggest that any significant heat may be created by the Earth's magnetic field, as suggested by some contemporary folk theories.
In Earth's continental crust, the decay of natural radioactive isotopes makes a significant contribution to geothermal heat production. The continental crust is abundant in lower density minerals but contains significant concentrations of heavier lithophilic minerals such as uranium; because of this, it holds the most concentrated global reservoir of radioactive elements found in the Earth. In layers closer to Earth's surface occurring isotopes are enriched in the granite and basaltic rocks; these high levels of radioactive elements are excluded from the Earth's mantle due to their inability to substitute in mantle minerals and consequent enrichment in melts during mantle melting processes. The mantle is made up of high density minerals with higher concentrations of elements that have small atomic radii such as magnesium and calcium; the geothermal gradient is steeper in the lithosphere than in the mantle because the mantle transports heat by convection, leading to a geothermal gradient, determined by the mantle adiabat, rather than by the conductive heat transfer processes that predominate in the lithosphere, which acts as a thermal boundary layer of the convecting mantle.
Heat flows from its sources within the Earth to the surface. Total heat loss from the Earth is estimated at 44.2 TW. Mean heat flow is 101 mW/m2 over oceanic crust; this is 0.087 watt/square meter on average, but is much more concentrated in areas where the lithosphere is thin, such as along mid-ocean ridges and near mantle plumes. The Earth's crust acts as a thick insulating blanket which must be pierced by fluid conduits in order to release the heat underneath. More of the heat in the Earth is lost through plate tectonics, by mantle upwelling associated with mid-ocean ridges; the final major mode of heat loss is by conduction through the lithosphere, the majority of which occurs in the oceans due to the crust there being much thinner and younger than under the continents. The heat of the Earth is replenished by radioactive decay at a rate of 30 TW; the global geothermal flow rates are more than twice the rate of human energy consumption from all primary sources. Heat from Earth's interior can be used as an energy source, known as geothermal energy.
The geothermal gradient has been used for space heating and bathing since ancient Roman times, more for generating electricity. As the human population continues to grow, so does energy use and the correlating environmental impacts that are consistent with global primary sources of energy; this has caused a growing interest in finding sources of energy that are renewable and have reduced greenhouse gas emissions. In areas of high geothermal energy density, current technology allows for the generation of electrical power because of the corresponding high temperatures. Generating electrical power from geothermal resources requires no fuel while providing true baseload energy at a reliability rate that exceeds 90%. In order to extract geothermal energy, it is necessary to efficiently transfer heat from a
Boussinesq approximation (water waves)
In fluid dynamics, the Boussinesq approximation for water waves is an approximation valid for weakly non-linear and long waves. The approximation is named after Joseph Boussinesq, who first derived them in response to the observation by John Scott Russell of the wave of translation; the 1872 paper of Boussinesq introduces the equations now known as the Boussinesq equations. The Boussinesq approximation for water waves takes into account the vertical structure of the horizontal and vertical flow velocity; this results in non-linear partial differential equations, called Boussinesq-type equations, which incorporate frequency dispersion. In coastal engineering, Boussinesq-type equations are used in computer models for the simulation of water waves in shallow seas and harbours. While the Boussinesq approximation is applicable to long waves – that is, when the wavelength is large compared to the water depth – the Stokes expansion is more appropriate for short waves; the essential idea in the Boussinesq approximation is the elimination of the vertical coordinate from the flow equations, while retaining some of the influences of the vertical structure of the flow under water waves.
This is useful because the waves propagate in the horizontal plane and have a different behaviour in the vertical direction. As in Boussinesq's case, the interest is in the wave propagation; this elimination of the vertical coordinate was first done by Joseph Boussinesq in 1871, to construct an approximate solution for the solitary wave. Subsequently, in 1872, Boussinesq derived the equations known nowadays as the Boussinesq equations; the steps in the Boussinesq approximation are: a Taylor expansion is made of the horizontal and vertical flow velocity around a certain elevation, this Taylor expansion is truncated to a finite number of terms, the conservation of mass for an incompressible flow and the zero-curl condition for an irrotational flow are used, to replace vertical partial derivatives of quantities in the Taylor expansion with horizontal partial derivatives. Thereafter, the Boussinesq approximation is applied to the remaining flow equations, in order to eliminate the dependence on the vertical coordinate.
As a result, the resulting partial differential equations are in terms of functions of the horizontal coordinates. As an example, consider potential flow over a horizontal bed in the plane, with x the horizontal and z the vertical coordinate; the bed is located at z = − h. A Taylor expansion is made of the velocity potential φ around the bed level z = −h: φ = φ b + z = − h + 1 2 2 z = − h + 1 6 3 z = − h + 1 24 4 z = − h + ⋯, where φb is the velocity potential at the bed. Invoking Laplace's equation for φ, as valid for incompressible flow, gives: φ = { φ b − 1 2 2 ∂ 2 φ b ∂
Volcanogenic massive sulfide ore deposit
Volcanogenic massive sulfide ore deposits known as VMS ore deposits, are a type of metal sulfide ore deposit copper-zinc which are associated with and created by volcanic-associated hydrothermal events in submarine environments. These deposits are sometimes called volcanic-hosted massive sulfide deposits; the density is 4500 kg/m3. They are predominantly stratiform accumulations of sulfide minerals that precipitate from hydrothermal fluids on or below the seafloor in a wide range of ancient and modern geological settings. In modern oceans they are synonymous with sulfurous plumes called black smokers, they occur within environments dominated by volcanic or volcanic derived rocks, the deposits are coeval and coincident with the formation of said volcanic rocks. As a class, they represent a significant source of the world's copper, lead and silver ores, with cobalt, barium, selenium, cadmium, bismuth, tellurium and germanium as co- or by-products. Volcanogenic massive sulfide deposits are forming today on the seafloor around undersea volcanoes along many mid ocean ridges, within back-arc basins and forearc rifts.
Mineral exploration companies are exploring for seafloor massive sulfide deposits. The close association with volcanic rocks and eruptive centers sets VMS deposits apart from similar ore deposit types which share similar source and trap processes. Volcanogenic massive sulfide deposits are distinctive in that ore deposits are formed in close temporal association with submarine volcanism and are formed by hydrothermal circulation and exhalation of sulfides which are independent of sedimentary processes, which sets VMS deposits apart from sedimentary exhalative deposits. There is a subclass of VMS deposits, the volcanic- and sediment-hosted massive sulfide deposits, that do share characteristics that are hybrid between the VMS and SEDEX deposits. Notable examples of this class include the deposits of New Brunswick, Canada; the source of metal and sulfur in VMS deposits is a combination of incompatible elements which are leached from the volcanic pile in the sub-seafloor hydrothermal alteration zone by hydrothermal circulation.
Hydrothermal circulation is considered to be driven via heat in the crust related to deep-seated gabbro intrusions. Transport of metals occurs via convection of hydrothermal fluids, the heat for this supplied by the magma chamber which sits below the volcanic edifice. Cool ocean water is drawn into the hydrothermal zone and is heated by the volcanic rock and is expelled into the ocean, the process enriching the hydrothermal fluid in sulfur and metal ions; the ore materials are trapped within a fumarole field or a black smoker field when they are expelled into the ocean and precipitate sulfide minerals as stratiform sulfide ore. Some deposits show evidence of formation via deposition of sulfide via replacement of altered volcanosedimentary rocks and may form by invasion of sulfur-rich brines into unconsolidated sediments; the typical location for VMS deposits is at the top of the felsic volcanic sequence, within a sequence of volcaniclastic tuffaceous epiclastics, sediments or fine tuffs which are related to the underlying volcanics.
The hangingwall to the deposit is broadly related to a more mafic sequence of volcanic rocks, either andesite, or basalt or absent or sediments only. VMS deposits are associated spatially and temporally with felsic volcanic rocks present in the stratigraphy below the deposit, as the direct footwall to the deposit. Sediments are contiguous with VMS deposits in some form or another and are present as cherts and chemical sediments deposited within a submarine environment; the hanging wall to the deposit can be volcanic units contiguous and coeval with the footwall rocks, indicating mineralisation was developed in an inter-eruptive pause. Hybrid VMS-SEDEX deposits of the siliciclastic associations may be developed within interflow sediments or within units of sedimentary rocks which are present discontinuously throughout a larger and contiguous volcanic package. Altogether, these geological features have been interpreted to show an association of VMS deposits with hydrothermal systems developed above or around submarine volcanic centres.
VMS deposits have a wide variety of morphologies, with mound shaped and bowl shaped deposits most typical. The bowl-shaped formations formed due to venting of hydrothermal solutions into submarine depressions - in many cases, this type of deposit can be confused with sedimentary exhalative deposits; the mound-shaped deposits formed in a way similar to that of modern massive sulfide deposits - via production of a hydrothermal mound formed by successive black smoker chimneys. Deposits that have formed in environments dominated by sedimentary rocks or permeable volcanic rocks can show a tabular morphology that mimics the geometry of the surrounding rocks. VMS deposits have an ideal form of a conical area of altered volcanic or volcanogenic sedimentary rock within the feeder zone, called the stringer sulfide or stockwork zone, o
Metamorphism
Metamorphism is the change of minerals or geologic texture in pre-existing rocks, without the protolith melting into liquid magma. The change occurs due to heat and the introduction of chemically active fluids; the chemical components and crystal structures of the minerals making up the rock may change though the rock remains a solid. Changes at or just beneath Earth's surface due to weathering or diagenesis are not classified as metamorphism. Metamorphism occurs between diagenesis, melting; the geologists who study metamorphism are known as "metamorphic petrologists." To determine the processes underlying metamorphism, they rely on statistical mechanics and experimental petrology. Three types of metamorphism exist: contact and regional. Metamorphism produced with increasing pressure and temperature conditions is known as prograde metamorphism. Conversely, decreasing temperatures and pressure characterize retrograde metamorphism. Metamorphic rocks can change without melting. Heat causes atomic bonds to break, the atoms move and form new bonds with other atoms, creating new minerals with different chemical components or crystalline structures, or enabling recrystallization.
When pressure is applied, somewhat flattened grains that orient in the same direction have a more stable configuration. The temperature lower limit on what is considered to be a metamorphic process is considered to be 100 – 200 °C; the upper boundary of metamorphic conditions is related to the onset of melting processes in the rock. The maximum temperature for metamorphism is 700 – 900 °C, depending on the pressure and on the composition of the rock. Migmatites are rocks formed at this upper limit, which contains pods and veins of material that has started to melt but has not segregated from the refractory residue. Since the 1980s it has been recognized that rocks are dry enough and of a refractory enough composition to record without melting "ultra-high" metamorphic temperatures of 900 – 1100 °C; the metamorphic process has to be over pressure of at least 100 mega pascals but below 300 mega pascals, the depth of 100 mega pascals varies depending on what type of rock is applying pressure. Regional or Barrovian metamorphism covers large areas of continental crust associated with mountain ranges those associated with convergent tectonic plates or the roots of eroded mountains.
Conditions producing widespread regionally metamorphosed rocks occur during an orogenic event. The collision of two continental plates or island arcs with continental plates produce the extreme compressional forces required for the metamorphic changes typical of regional metamorphism; these orogenic mountains are eroded, exposing the intensely deformed rocks typical of their cores. The conditions within the subducting slab as it plunges toward the mantle in a subduction zone produce regional metamorphic effects, characterized by paired metamorphic belts; the techniques of structural geology are used to unravel the collisional history and determine the forces involved. Regional metamorphism can be described and classified into metamorphic facies or metamorphic zones of temperature/pressure conditions throughout the orogenic terrane. Contact metamorphism occurs around intrusive igneous rocks as a result of the temperature increase caused by the intrusion of magma into cooler country rock; the area surrounding the intrusion where the contact metamorphism effects are present is called the metamorphic aureole.
Contact metamorphic rocks are known as hornfels. Rocks formed by contact metamorphism may not present signs of strong deformation and are fine-grained. Contact metamorphism is greater adjacent to the intrusion and dissipates with distance from the contact; the size of the aureole depends on the heat of the intrusion, its size, the temperature difference with the wall rocks. Dikes have small aureoles with minimal metamorphism whereas large ultramafic intrusions can have thick and well-developed contact metamorphism; the metamorphic grade of an aureole is measured by the peak metamorphic mineral which forms in the aureole. This is related to the metamorphic temperatures of pelitic or aluminosilicate rocks and the minerals they form; the metamorphic grades of aureoles are sillimanite hornfels, pyroxene hornfels. Magmatic fluids coming from the intrusive rock may take part in the metamorphic reactions. An extensive addition of magmatic fluids can modify the chemistry of the affected rocks. In this case the metamorphism grades into metasomatism.
If the intruded rock is rich in carbonate the result is a skarn. Fluorine-rich magmatic waters which leave a cooling granite may form greisens within and adjacent to the contact of the granite. Metasomatic altered aureoles can localize the deposition of metallic ore minerals and thus are of economic interest. A special type of contact metamorphism, associated with fossil fuel fires, is known as pyrometamorphism. Hydrothermal metamorphism is the result of the interaction of a rock with a high-temperature fluid of variable composition; the difference in composition between an existing rock and the invading fluid triggers a set of metamorphic and metasomatic reactions. The hydrothermal fluid may be magmatic, circulating ocean water. Convective circulation of hydrothermal fluids in the ocean floor basalts produces extensive hydrothermal metamorphism adjacent to spreading centers and other submarine volcanic areas
Magma
Magma is the molten or semi-molten natural material from which all igneous rocks are formed. Magma is found beneath the surface of the Earth, evidence of magmatism has been discovered on other terrestrial planets and some natural satellites. Besides molten rock, magma may contain suspended crystals and gas bubbles. Magma is produced by melting of the mantle and/or the crust at various tectonic settings, including subduction zones, continental rift zones, mid-ocean ridges and hotspots. Mantle and crustal melts migrate upwards through the crust where they are thought to be stored in magma chambers or trans-crustal crystal-rich mush zones. During their storage in the crust, magma compositions may be modified by fractional crystallization, contamination with crustal melts, magma mixing, degassing. Following their ascent through the crust, magmas may feed a volcano or solidify underground to form an intrusion. While the study of magma has 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, once in Hawaii.
Most magmatic liquids are rich in silica. Silicate melts are composed of silicon, aluminium, magnesium, calcium and potassium; the physical behaviours of melts depend upon their atomic structures as well as upon temperature and pressure and composition. Viscosity is a key melt property in understanding the behaviour of magmas. More silica-rich melts are more polymerized, with more linkage of silica tetrahedra, so are more viscous. Dissolution of water drastically reduces melt viscosity. Higher-temperature melts are less viscous. Speaking, more mafic magmas, such as those that form basalt, are hotter and less viscous than more silica-rich magmas, such as those that form rhyolite. Low viscosity leads to less explosive eruptions. Characteristics of several different magma types are as follows: Ultramafic SiO2 < 45% Fe–Mg > 8% up to 32%MgO Temperature: up to 1500°C Viscosity: Very Low Eruptive behavior: gentle or explosive Distribution: divergent plate boundaries, hot spots, convergent plate boundaries.
At any given pressure and for any given composition of rock, a rise in temperature past the solidus will cause 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 under mid-ocean ridges and volcanic arc environments. It is 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 be considered to include volatile phases such as water and carbon dioxide. The presence of volatile phases in a rock under pressure can stabilize a melt fraction; the presence of 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 freeze or solidify.
A major portion of all magma is silica, a compound of silicon and oxygen. Magma contains gases, which expand as the magma rises. Magma, high in silica resists flowing, so expanding gases are trapped in it. Pressure builds up until the gases blast out in a dangerous explosion. Magma, poor in silica flows so gas bubbles move up through it and escape gently. Melting of solid rocks to form magma is controlled by three physical parameters: temperature and composition; the most common mechanisms of magma generation in the mantle are decompression melting and lowering of the solidus. Mechanisms are discussed further in the entry for igneous rock; when rocks melt, they do so and because most rocks are made of several minerals, which all have different melting points. As a rock melts, for example, its volume changes; when enough rock is melted, the small globules of melt 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.
Melts can stay in place long enough to melt to 20% or 35%, but rocks are melted in excess of 50%, because the melted rock mass becomes a crystal-and-melt mush tha
Great Artesian Basin
The Great Artesian Basin, located in Australia, is the largest and deepest artesian basin in the world, stretching over 1,700,000 square kilometres, with measured water temperatures ranging from 30–100 °C. The basin provides the only source of fresh water through much of inland Australia; the Basin underlies 22% of the continent, including the states and territories of Queensland, the Northern Territory, South Australia, New South Wales. The basin is 3,000 metres deep in places and is estimated to contain 64,900 cubic kilometres of groundwater; the Great Artesian Basin Coordinating Committee coordinates activity between the various levels of government and community organisations. This area is one of the distinct physiographic provinces of the larger East Australian Basins division, includes the smaller Wilcannia Threshold physiographic section; the water of the GAB is held in a sandstone layer laid down by continental erosion of higher ground during the Triassic and early Cretaceous periods.
During a time when much of what is now inland Australia was below sea level, the sandstone was covered by a layer of marine sedimentary rock shortly afterward, which formed a confining layer, thus trapping water in the sandstone aquifer. The eastern edge of the basin was uplifted; the other side was created from the landforms of the Central Eastern Lowlands and the Great Western Plateau to the west. Most recharge water enters the rock formations from high ground near the eastern edge of the basin and gradually flows toward the south and west. A much smaller amount enters along the western margin in arid central Australia, flowing to the south and east; because the sandstones are permeable, water makes its way through the pores between the sand grains, flowing at a rate of one to five metres per year. Discharge water exits through a number of springs and seeps in the southern part of the basin; the age of the groundwater determined by carbon-14 and chlorine-36 measurements combined with hydraulic modelling ranges from several thousand years for the recharge areas in the north to nearly 2 million years in the south-western discharge zones.
Prior to European occupation, waters of the GAB discharged through mound springs, many in arid South Australia. These springs supported a variety of endemic invertebrates, supported extensive Aboriginal communities and trade routes. After the arrival of Europeans, they enabled early exploration and faster communications between southeastern Australia and Europe via the Australian Overland Telegraph Line; the Great Artesian Basin became an important water supply for cattle stations and livestock and domestic usage, is a vital life line for rural Australia. To tap it, water wells are drilled down to a suitable rock layer, where the pressure of the water forces it up without pumping; the discovery and use of water held underground in the Great Artesian Basin opened up thousands of square miles of country away from rivers in inland New South Wales and South Australia unavailable for pastoral activities. European discovery of the basin dates from 1878 when a shallow bore near Bourke produced flowing water.
There were similar discoveries in 1886 at Back Creek east of Barcaldine, in 1887 near Cunnamulla. In essence, water extraction from the GAB is a mining operation, with recharge much less than current extraction rates. In 1915, there were 1,500 bores providing 2,000 megalitres of water per day, but today the total output has dropped to 1,500 megalitres per day; this included just under 2000 flowing bores and more than 9000 that required mechanical power to bring water to the surface. Many bores abandoned, resulting in considerable water wastage; these problems have existed for many decades, in January 2007 the Australian Commonwealth Government announced additional funding in an attempt to bring them under control. However, many of the mound springs referred to above have dried up due to a drop in water pressure resulting in extinction of several invertebrate species; the Olympic Dam mine in South Australia is permitted to extract up to 42 million litres of water daily from the Great Artesian Basin under the Roxby Downs Act 1982.
The underground copper and uranium mine commenced operations in 1988 and is expected to continue operating until 2060. In addition, the basin has provided water via a 1.2 km deep bore for a geothermal power station at Birdsville. The heated water provides 25 % of the town's needs. Ergon Energy is expanding the 80 kW plant to meet Birdsville's electricity requirements; as the Great Artesian Basin underlies parts of Queensland, New South Wales, South Australia and the Northern Territory, which each operate under different legislative frameworks and resource management approaches, a coordinated "whole-of-Basin" approach to the management of this important natural resource is required. The Great Artesian Basin Coordinating Committee provides advice from community organisations and agencies to State and Australian Government Ministers on efficient and sustainable whole-of-Basin resource management and to coordinate activity between stakeholders. Membership of the Committee comprises all State and Australian Government agencies with responsibilities for management of parts of the Great Artesian Basin, community representatives nominated by agencies.
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