Geologist’s Exam 2009
Geology Paper – II
Part – 1
1. Write short notes on any ten of the following: 5×10=50
Abyssal zone, portion of the ocean deeper than about 2,000 m (6,600 feet) and shallower than about 6,000 m (20,000 feet). The zone is defined mainly by its extremely uniform environmental conditions, as reflected in the distinct life forms inhabiting it. The upper boundary between the abyssal zone and the overlying bathyal zone is conveniently defined as the depth at which the water temperature is 4° C (39° F); this depth varies between 1,000 and 3,000 m. Waters deeper than 6,000 m are treated separately as the hadal realm by ecologists.
The abyssal realm is the largest environment for Earth life, covering 300,000,000 square km (115,000,000 square miles), about 60 percent of the global surface and 83 percent of the area of oceans and seas.
Abyssal waters originate at the air-sea interface in polar regions, principally the Antarctic. There, the cold climate produces sea ice and residual cold brine. Because of its high density, the brine sinks and slowly flows along the bottom toward the Equator. Abyssal salinities range narrowly between 34.6 and 35.0 parts per thousand, and temperatures are mostly between 0° and 4° C (32° and 39° F). Pressure increases by about one atmosphere (approximately 14.7 pounds per square inch at sea level) with each 10-metre increment in depth; thus, abyssal pressures range between 200 and 600 atmospheres. Pressure presents few problems for abyssal animals, however, because the pressures within their bodies are the same as those outside them.
The concentrations of nutrient salts of nitrogen, phosphorus, and silica are very uniform in abyssal waters and are much higher than in overlying waters. This is because abyssal and hadal waters are the reservoir for the salts from decomposed biological materials that settle downward from upper zones, and the lack of sunlight prevents their uptake by photosynthesis.
The oxygen content of abyssal water depends entirely upon the amounts dissolved into it at its polar site of origin and the absence of photosynthesis, which precludes the introduction of new oxygen at depth. Abyssal waters retain several cubic centimetres of dissolved oxygen per litre, because the sparse animal populations do not consume oxygen faster than it is introduced into the abyssal zone. Abyssal life is concentrated at the seafloor, however, and the water nearest the floor may be essentially depleted in oxygen.
The abyssal realm is very calm, being far removed from storms that agitate the ocean at the air-sea interface. These low energies are reflected in the character of abyssal sediments. The abyssal realm is usually far enough from land that the sediment is composed predominantly of microscopic planktonremains produced in the food chain in the overlying waters, from which they settle. Abyssal sediment in waters shallower than 4,000 m in equatorial to temperate regions is composed primarily of the calcareous shells of foraminiferan zooplankton and of phytoplankton such as coccolithophores. Below 4,000 m, calcium carbonate tends to dissolve, and the principal sediment constituents are brown clays and the siliceous remains of radiolarian zooplankton and such phytoplankton as diatoms.
Abyssal fauna, though very sparse and embracing relatively few species, include representatives of all major marine invertebrate phyla and several kinds of fish, all adapted to an environment marked by no diurnal or seasonal changes, high pressures, darkness, calm water, and soft sediment bottoms. These animals tend to be gray or black, delicately structured, and unstreamlined. Mobile forms have long legs; and animals attached to the bottom have stalks, enabling them to rise above the water layer nearest the bottom, where oxygen is scarce. Abyssal crustaceans and fish may be blind. With increasing depth, carnivores and scavengers become less abundant than animals that feed on mud and suspended matter. Abyssal animals are believed to reproduce very slowly.
(b) Interfacial angle
The law of the constancy of interfacial angles (or ‘first law of crystallography’) states that the angles between the crystal faces of a given species are constant, whatever the lateral extension of these faces and the origin of the crystal, and are characteristic of that species.
Evaporite, any of a variety of individual minerals found in the sedimentary deposit of soluble salts that results from the evaporation of water.
Fig:Three models for deposition of marine evaporites in basins of restricted water circulation. Encyclopædia Britannica, Inc.
Evaporites are layered crystalline sedimentary rocks that form from brines generated in areas where the amount of water lost by evaporation exceeds the total amount of water from rainfall and influx via rivers and streams. The mineralogy of evaporite rocks is complex, with almost 100 varieties possible, but less than a dozen species are volumetrically important. Minerals in evaporite rocks include carbonates (especially calcite, dolomite, magnesite, and aragonite), sulfates (anhydrite and gypsum), and chlorides (particularly halite, sylvite, and carnallite), as well as various borates, silicates, nitrates, and sulfocarbonates. Evaporite deposits occur in both marine and nonmarine sedimentary successions.
Though restricted in area, modern evaporites contribute to genetic models for explaining ancient evaporite deposits. Modern evaporites are limited to arid regions (those of high temperature and low rates of precipitation), for example, on the floors of semidry ephemeral playa lakes in the Great Basin of Nevada and California, across the coastal salt flats (sabkhas) of the Middle East, and in salt pans, estuaries, and lagoons around the Gulf of Suez. Ancient evaporates occur widely in the Phanerozoic geologic record, particularly in those of Cambrian (from 570 to 505 million years ago), Permian (from 286 to 245 million years ago), and Triassic (from 245 to 208 million years ago) age, but are rare in sedimentary sequences of Precambrian age. They tend to be closely associated with shallow marine shelf carbonates and fine (typically rich in iron oxide) mudrocks. Because evaporite sedimentation requires a specific climate and basin setting, their presence in time and space clearly constrains inferences of paleoclimatology and paleogeography. Evaporite beds tend to concentrate and facilitate major thrust fault horizons, so their presence is of particular interest to structural geologists. Evaporites also have economic significance as a source of salts and fertilizer.
All evaporite deposits result from the precipitation of brines generated by evaporation. Laboratory experiments can accurately trace the evolution of brines as various evaporite minerals crystallize. Normal seawater has a salinity of 3.5 percent (or 35,000 parts per million), with the most important dissolved constituents being sodium and chlorine. When seawater volume is reduced to one-fifth of the original, evaporite precipitation commences in an orderly fashion, with the more insoluble components (gypsum and anhydrite) forming first. When the solution reaches one-tenth the volume of the original, more soluble minerals like sylvite and halite form. Natural evaporite sequences show vertical changes in mineralogy that crudely correspond to the orderly appearance of mineralogy as a function of solubility but are less systematic.
Evaporite deposition in the nonmarine environment occurs in closed lakes—i.e., those without outlet—in arid and semiarid regions. Such lakes form in closed interior basins or shallow depressions on land where drainage is internal and runoff does not reach the sea. If water depths are shallow or, more typically, somewhat ephemeral, the term playa or playa lake is commonly used.
Water inflow into closed lakes consists principally of precipitation and surface runoff, both of which are small in amount and variable in occurrence in arid regions. Groundwater flow and discharge from springs may provide additional water input, but evaporation rates are always in excess of precipitation and surface runoff. Sporadic or seasonal storms may give rise to a sudden surge of water inflow. Because closed lakes lack outlets, they can respond to such circumstances only by deepening and expanding. Subsequent evaporation will reduce the volume of water present to prestorm or normal amount; fluctuation of closed lake levels therefore characterizes the environment.
Such changing lake levels and water volumes lead to fluctuating salinity values. Variations in salinity effect equilibrium relations between the resulting brines and lead to much solution and subsequent reprecipitation of evaporites in the nonmarine environment. As a result of these complexities as well as the distinctive nature of dissolved constituents in closed lake settings, nonmarine evaporite deposits contain many minerals that are uncommon in marine evaporites—e.g., borax, epsomite, trona, and mirabilite.
Shallow marine environment
Evaporite deposition in the shallow marine environment (sometimes termed the salina) occurs in desert coastal areas, particularly along the margins of such semi-restricted water bodies as the Red Sea, Persian Gulf, and Gulf of California. Restriction is, in general, one of the critical requirements for evaporite deposition, because free and unlimited mixing with the open sea would allow the bodies of water to easily overcome the high evaporation rates of arid areas and dilute these waters to near-normal salinity. This semi-restriction cannot, in fact, prevent a large amount of dilution by mixing; coastal physiography is the principal factor involved in brine production. Shallow-water evaporites, almost exclusively gypsum, anhydrite, and halite, typically interfinger with tidal flat limestone and dolomite and fine-grained mudrock.
Most of the thick, laterally extensive evaporite deposits appear to have been produced in deep, isolated basins that developed during episodes of global aridity. The most crucial requirement for evaporite production is aridity; water must be evaporated more rapidly than it can be replenished by precipitation and inflow. In addition, the evaporite basin must somehow be isolated or at least partially isolated from the open ocean so that brines produced through evaporation are prevented from returning there. Restricting brines to such an isolated basin over a period of time enables them to be concentrated to the point where evaporite mineral precipitation occurs. Periodic breaching of the barrier, due either to crustal downwarping or to global sea-level changes, refills the basin from time to time, thereby replenishing the volume of seawater to be evaporated and making possible the inordinately thick, regionally extensive evaporite sequences visible in the geologic record.
Debate continues over the exact mechanisms for generating thick evaporite deposits. Three possible models for restricting “barred” evaporite basins are shown in Figure 5. They differ in detail, and none has garnered a consensus of support. The deep-water, deep-basin model accounts for replenishment of the basin across the barrier or sill, with slow, continual buildup of thick evaporites made possible by the seaward escape of brine that allows a constant brine concentration to be maintained. The shallow-water, shallow-basin model produces thick evaporites by continual subsidence of the basin floor. The shallow-water, deep-basin model shows the brine level in the basin beneath the level of the sea as a result of evaporation; brines are replenished by groundwater recharge from the open ocean.
Benthos, the assemblage of organisms inhabiting the seafloor. Benthic epifauna live upon the seafloor or upon bottom objects; the so-called infauna live within the sediments of the seafloor. By far the best-studied benthos are the macrobenthos, those forms larger than 1 mm (0.04 inch), which are dominated by polychaete worms, pelecypods, anthozoans, echinoderms, sponges, ascidians, and crustaceans. Meiobenthos, those organisms between 0.1 and 1 mm in size, include polychaetes, pelecypods, copepods, ostracodes, cumaceans, nematodes, turbellarians, and foraminiferans. The microbenthos, smaller than 0.1 mm, include bacteria, diatoms, ciliates, amoeba, and flagellates.
The variety and abundance of the benthos vary with latitude, depth, water temperature and salinity, locally determined conditions such as the nature of the substrate, and ecological circumstances such as predation and competition. The principal food sources for the benthos are plankton and organic debris from land. In shallow water, larger algae are important, and, where light reaches the bottom, benthic photosynthesizing diatoms are also a significant food source. Hard and sandy substrates are populated by suspension feeders such as sponges and pelecypods. Softer bottoms are dominated by deposit eaters, of which the polychaetes are the most important. Fishes, starfish, snails, cephalopods, and the larger crustaceans are important predators and scavengers.
(e) Suspended load
Suspended load consists of sediment particles that are mechanically transported by suspension within a stream or river. This is in contrast to bed or traction load , which consists of particles that are moved along the bed of a stream, and dissolved load, which consists of material that has been dissolved in the stream water . In most streams, the suspended load is composed primarily of silt and clay size particles. Sand-size particles can also be part of the suspended load if the stream flow velocity and turbulence are great enough to hold them in suspension.
The suspended load can consist of particles that are intermittently lifted into suspension from the stream bed and of wash load, which remains continuously suspended unless there is a significant decrease in stream flow velocity. Wash load particles are finer than those along the stream bed, and therefore must be supplied by bank erosion, mass wasting , and mass transport of sediment from adjacent watersheds into the stream during rainstorms.
Water density is proportional to the amount of suspended load being carried. Muddy water high in suspended sediment will therefore increase the particle buoyancy and reduce the critical shear stress required to move the bed load of the stream.
The ratio of suspended load to bed load in a stream depends on the ratio of the shear velocity (a property of the flowing water that reflects the degree of turbulence) and the fall velocity (a property of the sediment grains). The fall velocity is that at which a sediment particle will fall through still water, and thus depends on both grain size and mineralogy (density). Bed-load transport predominates when the shear velocity is significantly less than half the fall velocity and suspended load transport predominates when the shear velocity is two to three times greater than the fall velocity. Mixed-mode transport occurs when the ratio falls within a range of approximately 0.4 to 2.5.
Fig:Modes of transportation of sediments and dissolved ions (represented by red dots with + and – signs) in a stream.
(f) Pelagic sediments
Pelagic sediment or pelagite is a fine-grained sediment that accumulates as the result of the settling of particles to the floor of the open ocean, far from land. These particles consist primarily of either the microscopic, calcareous or siliceous shells of phytoplankton or zooplankton; clay-size siliciclastic sediment; or some mixture of these. Trace amounts of meteoric dust and variable amounts of volcanic ash also occur within pelagic sediments. Based upon the composition of the ooze, there are three main types of pelagic sediments: siliceous oozes, calcareous oozes, and red clays.
The composition of pelagic sediments is controlled by three main factors. The first factor is the distance from major landmasses, which affects their dilution by terrigenous, or land-derived, sediment. The second factor is water depth, which affects the preservation of both siliceous and calcareous biogenic particles as they settle to the ocean bottom. The final factor is ocean fertility, which controls the amount of biogenic particles produced in surface waters.
- Carbonate Ooze (mostly CaCO3 – Calcite, Aragonite)
- Foraminiferal (mostly Globigerina) – Calcite
- Nannofossil (Coccolithophorids) – Calcite
- Pteropod (planktonic Mollusks) – Aragonite
- Siliceous Ooze (mostly X-ray amorphous SiO2)
- Red (Brown) Clay (Mostly everything else!)
- Clay Minerals (Illite, Smectite*, Kaolinite, Chlorite)
- Zeolites* (Phillipsite, Clinoptilolite)
- Barite (BaSO4)
- Phosphates (skeletal Apatite, Francolite)
- Ferromanganese Oxides (micronodules)
- Volcanic Ash (glass, feldspar, pyroxene phenocrysts)
- Cosmogenic Debris ( tektites, Fe-Ni spherules)
- Refractory Organics
- Other (e.g., Rutile (TiO2) in the insoluble residue of Mnnodules)
- Hemipelagic Muds (blue, green, red, volcanic, coral)
- Glacial-Marine (marine glacial ourwash)
Pelagic sediments are fine grained deep sea sediment composed of largely of biogenicooze that is often rich in foraminifera with 60% pelagic and neritic grains. It can also be a red clay, with less than 40% siliciclastic and volcaniclastic grains. It can also be a silica ooze (often rich in radiolaria).
Pelagic sediments are deposited at such low rates that they tend to be overwhelmed near shore by terrigenous deposits from land. So, pelagic sediments are normally associated with deep sea regions.
Unlike terrigenous sediments, pelagic sediments are classified by composition, not size. The size of pelagic sediments is uniformly pretty small, so their widely varying compositions are more interesting.
(g) Mineral relief
Relief is the term used to describe the degree to which edges and surface imperfections of crystals are visible in plane-polarised light. Minerals with a refractive index very different from that of the mounting glue are said to have high relief as they appear to ‘stand out’ from the slide (as for pyroxene in Figure(a) ): grain boundaries are easily seen and surface imperfections appear pronounced. Minerals with refractive indices similar to that of the mounting glue are said to have low relief: individual grain boundaries are not easily observed and the minerals are featureless and almost invisible in plane-polarised light (as for plagioclase in Figure(b) ). If the mineral is highly anisotropic (e.g. calcite), the relief may vary as the stage is rotated: the transmitted light sampling first one permitted vibration direction, then the other.
Pyroxene, a high-relief mineral that stands out – parallel cleavage traces are clearly visible. (b) Plagioclase, a featureless low-relief mineral.
When examining minerals the results will be:
- Strong relief
- mineral stands out strongly from the mounting medium,
- whether the medium is oil, in grain mounts, or other minerals in thin section,
- for strong relief the indices of the mineral and surrounding medium differ by greater than 0.12 RI units.
- Moderate relief
- mineral does not strongly stand out, but is still visible,
- indices differ by 0.04 to 0.12 RI units.
- Low relief
- mineral does not stand out from the mounting medium,
- indices differ by or are within 0.04 RI units of each other.
A mineral may exhibit positive or negative relief:
- +ve relief – index of refraction for the material is greater than the index of the oil.
- -ve relief nmin < noil
e.g. garnet 1.76
e.g. fluorite 1.433
(h) CCD or calcium carbonate depth
Calcite compensation depth (CCD) is the depth at which the rate of carbonate accumulation equals the rate of carbonate dissolution. The input of carbonate to the ocean is through rivers and deep-sea hydrothermal vents. The CCD intersects the flanks of the world’s oceanic ridges, and as a result these are mostly blanketed by carbonate oozes, a biogenic ooze made up of skeletal debris. Carbonate oozes cover about half of the world’s seafloor and are present chiefly above a depth of 4,500 metres (about 14,800 feet); below that they dissolve quickly. In the Atlantic basin the CCD is 500 metres (about 1,600 feet) deeper than in the Pacific basin, reflecting both a high rate of supply and low rate of dissolution in comparison to the Pacific.Variation in input, productivity, and dissolution rates in the geologic past have caused the CCD to vary over 2,000 metres (about 6,600 feet).
In the geological past the depth of the CCD has shown significant variation. In the Cretaceous through to the Eocene the CCD was much shallower globally than it is today; due to intense volcanic activity during this period atmospheric CO2 concentrations were much higher. Higher concentrations of CO2 resulted in a higher partial pressure of CO2 over the ocean. This greater pressure of atmospheric CO2 leads to increased dissolved CO2 in the ocean mixed surface layer. This effect was somewhat moderated by the deep oceans’ elevated temperatures during this period. In the late Eocene the transition from a greenhouse to an icehouse Earth coincided with a deepened CCD.
Today, increasing atmospheric concentration of CO2 from combustion of fossil fuels may lead to shallower CCD, with zones of downwelling first being affected.
John Murray investigated and experimented on the dissolution of calcium carbonate and was first to identify the carbonate compensation depth in oceans.
(i) Retrograde metamorphism
Metamorphism is mineralogical and structural adjustments of solid rocks to physical and chemical conditions differing from those under which the rocks originally formed. Changes produced by surface conditions such as compaction are usually excluded. The most important agents of metamorphism include temperature, pressure, and fluids. Equally as significant are changes in chemical environment that result in two metamorphic processes:
(1) mechanical dislocation where a rock is deformed, especially as a consequence of differential stress; and
(2) chemical recrystallization where a mineral assemblage becomes out of equilibrium due to temperature and pressure changes and a new mineral assemblage forms.
Three types of metamorphism may occur depending on the relative effect of mechanical and chemical changes.
- Dynamic metamorphism
- Contact metamorphism
- Regional metamorphism
Retrograde metamorphism involves the reconstitution of a rock via revolatisation under decreasing temperatures (and usually pressures), allowing the mineral assemblages formed in prograde metamorphism to revert to those more stable at less extreme conditions. This is a relatively uncommon process, because volatiles must be present.
The recrystallization of pre-existing rocks in response to a lowering of metamorphic grade in the presence of a fluid phase. After reaching a metamorphic climax, lowering of metamorphic grade does not usually cause retrograde reactions to occur because all the water in the rock system has been expelled at the metamorphic climax, thus preserving high-grade mineral assemblages. If some water remains in the system, however, or is introduced as the grade decreases, the water can act as a catalyst to initiate retrograde reactions. The reactions produce hydrated mineral types, in contrast to the dehydration reactions of prograde metamorphism.
Greywacke or Graywacke is a variety of sandstone generally characterized by its hardness, dark color, and poorly sorted angular grains of quartz, feldspar, and small rock fragments or lithic fragments set in a compact, clay-fine matrix. It is a texturally immature sedimentary rock generally found in Paleozoic strata. The larger grains can be sand- to gravel-sized, and matrix materials generally constitute more than 15% of the rock by volume. The term “greywacke” can be confusing, since it can refer to either the immature (rock fragment) aspect of the rock or its fine-grained (clay) component.
The origin of greywacke was problematic until turbidity currents and turbidites were understood, since, according to the normal laws of sedimentation, gravel, sand and mud should not be laid down together. Geologists now attribute its formation to submarine avalanches or strong turbidity currents. These actions churn sediment and cause mixed-sediment slurries, in which the rocks may exhibit a variety of sedimentary features. Supporting the turbidity current origin theory is that deposits of greywacke are found on the edges of the continental shelves, at the bottoms of oceanic trenches, and at the bases of mountain formational areas. They also occur in association with black shales of deep sea origin.
Greywackes are mostly grey, brown, yellow or black, dull-colored sandy rocks which may occur in thick or thin beds along with shales and limestones. They can contain a very great variety of minerals, the principal ones being quartz, orthoclase and plagioclasefeldspars, calcite, iron oxides and graphitic, carbonaceous matters, together with (in the coarser kinds) fragments of such rocks as felsite, chert, slate, gneiss, various schists, and quartzite. Among other minerals found in them are biotite, chlorite, tourmaline, epidote, apatite, garnet, hornblende, augite, sphene and pyrites. The cementing material may be siliceous or argillaceous and is sometimes calcareous.
As a rule greywackes do not contain fossils, but organic remains may be common in the finer beds associated with them. Their component particles are usually not very rounded or polished, and the rocks have often been considerably indurated by recrystallization, such as the introduction of interstitial silica. In some districts the greywackes are cleaved, but they show phenomena of this kind much less perfectly than the slates. Some varieties include feldspathic greywacke, which is rich in feldspar, and lithic greywacke, which is rich in tiny rock fragments.
Although the group is so diverse that it is difficult to characterize mineralogically, it has a well-established place in petrographical classifications because these peculiar composite arenaceous deposits are very frequent among Silurian and Cambrian rocks, and are less common in Mesozoic or Cenozoic strata. Their essential features are their gritty character and their complex composition. By increasing metamorphism, greywackes frequently pass into mica-schists, chloritic schists and sedimentary gneisses.
(k) continuous series
Bowen’s reaction series is a means of ranking common igneous silicate minerals by the temperature at which they crystallize. Minerals at the top have a relatively high crystallization temperature, which means that they will be the first minerals to crystallize from a magma that is cooling. IF they are chemically compatible with the magma as it continues to cool, they will grow larger by addition of external layers of additional material. [They then can potentially become the phenocrysts in a porphyritic igneous texture.] If they are chemically incompatible, they will react with the melt and recrystallize into new minerals. What determines this chemical compatibility is in large part the silica content of the melt.
The Principles that Bowen realized are as follows:
- As a melt cools minerals crystallize that are in thermodynamic equilibrium with the melt (dissolution equals crystallization; if no equilibrium either crystallization will dominate [supersaturation], or dissolution [undersaturated]).
- As the melt keeps cooling and minerals keep crystallizing, the melt will change its composition.
- The earlier formed crystals will not be in equilibrium with this melt any more and will be dissolved again to form new minerals. In other words: these crystals react with the melt to form new crystals, therefore the name reaction series.
- The common minerals of igneous rocks can be arranged into two series, a continuous reaction series of the feldspars, and a discontinuous reaction series of the ferromagnesian minerals (olivine, pyroxene, hornblende, biotite).
- This reaction series implies that from a single “parental magma” all the various kinds of igneous rocks can be derived by Magmatic Differentiation.
Minerals on the left part of the “Y” of the diagram are called ferromagnesian minerals, because they contain iron (Latin: ferrum) and magnesium in their composition. This part of the series is referred to as the discontinuous series, since these minerals, if chemically incompatible with the melt as it cools, will usually completely react to form totally new minerals: an olivine crystallizing in a melt relatively high in silica (e.g., 60%) will completely recrystallize into pyroxene, and that pyroxene may in part or completely recrystallize into hornblende. Because they contain water (as OH – hydroxyl radicals) in their structures, hornblende and biotite in volcanic rocks are almost always phenocrysts that actually crystallized underground before the magma was erupted; they cannot form from crystallization of lava at the surface.
The minerals on the right arm of the “Y” are the plagioclase feldspars, which form a continuous series from 100% Ca-plagioclase (anorthite) with the highest melting point, to 100% Na-plagioclase (albite) with the lowest melting point. The first crystals forming may only partially re-react with the melt, but without destroying the basic feldspar crystal structure. Very often, large plagioclase crystals in an igneous rock will have cores that are more calcium rich than the outer layers, and sometimes this layering (called zonation) can be clearly seen under the microscope, or even with the naked eye for particularly large crystals. In general, melts higher in silica are higher in sodium (Na) and lower in calcium (Ca).
The lower portion of Bowen’s Reaction Series is dictated more by chemistry than is the upper part. Biotite, orthoclase feldspar and muscovite are the only minerals here that contain large amounts of potassium; of these three, only biotite is found in volcanic rocks. Orthoclase is a mineral found in plutonic rocks, those that crystallize entirely underground. Its high temperature (1000°C melting point) volcanic equivalent – with the same formula but a different crystal structure – is the mineral sanidine, which is common in high-silica volcanic rocks. Minerals near the bottom of the series also have much higher silica contents than the minerals at the top (e.g., pure olivine is about 38% SiO2, while pure sanidine is 65% SiO2). It is this increase in silica content that lowers the melting point; note that quartz, at the bottom of the series, is 100% SiO2 and has the lowest melting point (about 700°C).
As a result, rocks that crystallize from mafic melts (45-55% silica) will tend to be made up of minerals that are high in Bowen’s reaction series – such as olivine, pyroxene and Ca-rich plagioclase feldspar, and will crystallize at higher temperatures than more silica-rich melts.
Rocks from felsic melts (>65- 70% silica) will be composed mostly of minerals from the bottom of the series – hornblende and/or biotite, Na-rich plagioclase, sanidine and possibly quartz.
Rocks from intermediate magmas will contain minerals from the middle of the sequence. Worth noting is that these are the major minerals that will appear in the rocks; there will be numerous accessory minerals present that are not in Bowen’s reaction series; these are present in small quantities only in most cases, but can be very informative about fine details of the magma origins, history and properties. Finding minerals in a volcanic rock that shouldn’t be there can also be extremely informative about the magma history, since there has to be a reason for their existence! Low-silica minerals (e.g., calcium-rich plagioclase, olivine, pyroxene) tend to be dark in color – dark gray for the plagioclase, dark green to black for olivine and pyroxene. As a result, low-silica volcanic rocks are commonly dark in color – dark gray to black. In general, the higher the silica content, the lighter the color of the rocks.
(l) Pyroclastic rocks
Pyroclastic rocks or pyroclastics (meaning broken) are clastic rocks composed solely or primarily of volcanic materials. Where the volcanic material has been transported and reworked through mechanical action, such as by wind or water, these rocks are termed volcaniclastic. Commonly associated with unsieved volcanic activity—such as Plinian or krakatoan eruption styles, or phreatomagmatic eruptions—pyroclastic deposits are commonly formed from airborne ash, lapilli and bombs or blocks ejected from the volcano itself, mixed in with shattered country rock.
Pyroclastic rocks may be a range of clast sizes, from the largest agglomerates, to very fine ashes and tuffs. Pyroclasts of different sizes are classified as volcanic bombs, lapilli, and volcanic ash. Ash is considered to be pyroclastic because it is a fine dust made up of volcanic rock. One of the most spectacular forms of pyroclastic deposit are the ignimbrites, deposits formed by the high-temperature gas-and-ash mix of a pyroclastic flow event.
Sub-divisions of Pyroclastic-rock:
- Andesitic lapilli-tuff
- Dacitic lapilli-tuff
- Rhyolitic lapilli-tuff
- Calciocarbonatite lapillistone
- Palagonite tuff
- Andesitic tuff
- Basaltic tuff
- Crystal tuff
- Dactic tuff
- Epiclastic tuff
- Felsic tuff
- Lithic tuff
- Rhyodactic tuff
- Rhyolitic tuff
- Trachytic tuff
- Coarse tuff
- Fine tuff
- Vitric tuff
- Ultramafic tuff
- Welded tuff
- Volcanic breccia