What Type Of Particles/sediment Makeup Limestone? What Type Of Particles Makeup Limestone

Stone Fracture and Rock Forcefulness

Zong-Xian Zhang , in Rock Fracture and Blasting, 2016

iii.i.iii Sedimentary Rock

Sedimentary rock is formed by sedimentation of material at the World's surface and inside bodies of water. Particles that form a sedimentary rock by accumulating are chosen sediment. Before being deposited, sediment was formed by weathering and erosion in a source area, so transported to the place of deposition by water, wind, mass movement, or glaciers. Some examples of sedimentary rocks are limestone, sandstone, siltstone, shale, conglomerate, and breccia. About sedimentary rocks contain either quartz or calcite. In contrast with igneous and metamorphic rocks, a sedimentary stone usually contains very few different major minerals, and it has lower strengths and higher porosity.

Most sedimentary rocks are considered to be of anisotropy in their concrete and mechanic properties due to marked bedding structure, as shown in Fig. iii.anec. The term anisotropy refers to a mensurate of the directional properties of a fabric. However, some sedimentary rocks such as sandstone and limestone tin be considered to be isotropy. In full general, the physical and mechanical backdrop in the horizontal management to a bedding airplane are largely different from those in the vertical direction to the bedding plane. The cement or matrix between grains often determines the mechanical properties of sedimentary rocks. The sedimentary rocks containing dirt minerals are particularly sensitive to force per unit area and water. Therefore, water tin can cause a big problem in the stability of a tunnel or an hole-and-corner structure that is surrounded by sedimentary rock containing some clay minerals.

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Sediments, Diagenesis, and Sedimentary Rocks

Westward.B.Northward. Berry , in Treatise on Geochemistry, 2003

7.13.1 Introduction

Sedimentary rocks are commonly organized into discrete strata. The strata are composed of materials, diverse particles of inorganic and/or organic origin, that reflect aspects of the environmental conditions under which they got accumulated. Sequences of sedimentary-stone layers were seen and studied initially in cliffs, human being-fabricated exposures, and sites where the vegetation was non thick enough to obscure the rock layers. It was in mines, yet, that sequences of strata came to exist examined closely. Miner's observations of the succession of sedimentary rock layers they saw and quarried below the Globe'southward surface gave nativity to a domain, viz. stratigraphy, and an understanding of sedimentary rocks. Drupe (1968, 1987) discusses—from a historical perspective, the geological timescale—how an economic imperative became a meaning force in the evolution of chronometry of sedimentary rocks.

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The Oceans and Marine Geochemistry

H.D. Holland , in Treatise on Geochemistry, 2007

vi.21.iv.2 The Mesoproterozoic (one.viii–1.2   Ga)

Sedimentary rocks of the McArthur Basin in Northern Australia provide one of the best windows on the chemistry of the Mesoproterozoic ocean. Some 10  km of 1.vii–1.6   Ga sediments accumulated in this intracratonic basin (Southgate et al., 2000). In sure intervals, they contain giant strata-bound Pb–Zn–Ag mineral deposits (Jackson et al., 1987; Jackson and Raiswell, 1991; Crick, 1992). The sediments accept experienced only low-grade metamorphism.

Shen et al. (2002) accept reported data for the isotopic limerick of sulfur in carbonaceous shales of the lower office of the 1.73–ane.72   Ga Wollogorang Formation and in the lower part of the 1.64–1.63   Ga Reward Germination of the McArthur Basin. These shales were probably deposited in a euxinic intracratonic basin connected to the open sea. The δ³⁴S of pyrite in black shales of the Wollogorang Germination ranges from −1‰ to +6.3‰ with a hateful and standard derivation of four.0±i.9‰ (north=14). Donnelly and Jackson (1988) reported like values. The δ³⁴South values of pyrite in the Lower Reward Germination range from +18.two‰ to +23.iv‰ with an average and standard deviation of 18.iv±1.8‰ (due north=10). The spread of δ³⁴S values within each germination is relatively small. The sulfur is quite ³⁴S enriched compared to compositions expected from the reduction of seawater sulfate with a δ³⁴Due south of +20−25‰ (Strauss, 1993). This is especially true of the sulfides in the Reward Formation. Shen et al. (2002) propose that the Reward data are all-time explained if the concentration of sulfate in the contemporary seawater was between 0.5 and 2.4   mmol   kg^−one. Sulfate concentrations in the Mesoproterozoic bounding main well below those of the present oceans have also been proposed on the footing of the rapid change in the value of δ³⁴Southward in carbonate-associated sulfate of the 1.two   Ga Bylot Supergroup of northeastern Canada (Lyons et al., 2002). However, the value of in Mesoproterozoic seawater is still rather uncertain.

Somewhat of a cross-check on the SO_iv ²⁻ concentration of seawater can exist obtained from the evaporite relics in the McArthur Group (Walker et al., 1977). Upwards to 40% of the measured sections of the Amelia Dolomite consists of such relics in the course of carbonate pseudomorphs after a variety of morphologies of gypsum and anhydrite crystals, chert pseudomorphs afterwards anhydrite nodules, halite casts, and microscopic remnants of original, unaltered sulfate minerals. Muir (1979) and Jackson et al. (1987) have pointed out the similarity of this formation to the contempo sabkhas forth the Persian Gulf coast. The pseudomorphs crosscut sedimentary features such as bedding and laminated microbial mats, suggesting that the original sulfate minerals crystallized in the host sediments during diagenesis.

Pseudomorphs later halite are common throughout the McArthur Group. The halite appears to have formed by most complete evaporation of seawater in shallow marine environments and probably represents ephemeral table salt crusts. The general lack of association of halite and calcium sulfate minerals in these sediments probably resulted in part from the dissolution of previously deposited halite during surface flooding, but besides indicates that evaporation did not always continue across the calcium sulfate facies.

This observation allows a rough bank check on the reasonableness of Shen et al.'s (2002) estimate of the sulfate concentration in seawater during the deposition of the McArthur Group. On evaporating modern seawater, gypsum begins to precipitate when the degree of evaporation is ca. 3.8. As shown in Effigy 13, the onset of gypsum and/or anhydrite precipitation occurs at progressively greater degrees of evaporation as the product ${\mathrm{yard}}_{{\text{Ca}}^{2+}}{m}_{{\text{SO}}_4^{\mathrm{two}-}}$ in seawater decreases. Today, ${\mathrm{1000}}_{{\text{Ca}}^{\mathrm{ii}+}}{\mathrm{thou}}_{{\text{SO}}_4^{2-}}$ =280 (mmol   kg⁻¹)². If this product is reduced to 23 (mmol   kg⁻¹)ⁱⁱ, anhydrite begins to precipitate simultaneously with halite at a degree of evaporation of 10.eight. The presence of gypsum casts without halite in the sediments of the McArthur Grouping indicates that in seawater at that time $m_{{\text{Ca}}^{\mathrm{two}+}}{m}_{{\text{And so}}_{\mathrm{iv}}^{2-}}$ >23 (mmol   kg⁻¹)^two, provided the salinity of seawater was the same equally today. If ${\mathrm{grand}}_{{\text{SO}}_4^{\mathrm{ii}-}}$ was ii.4   mmol   kg⁻¹, the upper limit suggested by Shen et al. (2002), $m_{{\text{Ca}}^{2+}}$ must so have been >10   mmol   kg⁻ⁱ the concentration of Ca²⁺ in modern seawater. An Then₄ ²⁻ concentration of 2.4   mmol   kg⁻¹ is therefore permissible. Sulfate concentrations equally low as 0.five   mmol   kg^−ane require what are probably unreasonably high concentrations of Caⁱⁱ⁺ in seawater to business relationship for the precipitation of gypsum earlier halite in the McArthur Group sediments.

Figure xiii. The relationship between the value of the product $m_{{\text{Ca}}^2}+\cdot m_{{\text{SO}}_4^{\mathrm{two}-}}$ in seawater and the concentration cistron at which seawater becomes saturated with respect to gypsum at 25   °C and 1   atm (Kingdom of the netherlands, 1984; Eugster et al., 1980). Reprinted by permission of Princeton University Press from Holland (1984). © 1984 Princeton University Press.

The common occurrence of dolomite in the McArthur Group indicates that the $k_{{\text{Mg}}^{2+}}/{\mathrm{chiliad}}_{{\text{Ca}}^{2+}}$ ratio in seawater was >i (see below). This is likewise indicated by the common occurrence of aragonite as the major primary CaCO_iii stage of sediments on Archean and Proterozoic carbonate platforms (Grotzinger, 1989; Winefield, 2000). Although these hints regarding the composition of Mesoproterozoic seawater are welcome, they demand to be confirmed and expanded by analyses of fluid inclusions in marine calcite cements.

Perhaps the most interesting implication of the shut association of gypsum, anhydrite, and halite relics in the McArthur Group is that the temperature during the deposition of these minerals was non much above eighteen   °C, the temperature at which gypsum, anhydrite, and halite are stable together (Hardie, 1967). At higher temperatures, anhydrite is the stable calcium sulfate mineral in equilibrium with halite. The coexistence of gypsum and anhydrite with halite suggests that the temperature during their degradation was mayhap lower, but probably no higher than in the mod sabkhas of the Persian Gulf, where anhydrite is the dominant calcium sulfate mineral in association with halite (Kinsman, 1966).

In their paper on the carbonaceous shales of the McArthur Basin, Shen et al. (2002) comment that euxinic conditions were common in marine-connected basins during the Mesoproterozoic, and they suggest that low concentrations of seawater sulfate and reduced levels of atmospheric oxygen at this time are compatible with euxinic deep bounding main waters. Anbar and Knoll (2002) echo this sentiment. They point out that biologically important trace metals would then have been scarce in nearly marine environments, potentially restricting the nitrogen cycle, affecting principal productivity, and limiting the ecological distribution of eukaryotic algae. However, some of the presently available bear witness does not support the notion of a Mesoproterozoic euxinic ocean floor. Figure 14 shows that the redox-sensitive elements Mo, U, and Re are well correlated with the organic carbon content of carbonaceous shales in the McArthur Basin. The slope of the correlation lines is close to that in many Phanerozoic black shales, suggesting that the concentration of these elements in McArthur Basin seawater was comparable to their concentration in modern seawater. Preliminary data for the isotopic composition of Mo in the Wollogorang Germination of the McArthur Bowl (Arnold et al., 2002) suggest somewhat more than extensive sulfidic deposition of Mo in the Mesoproterozoic than in the modern oceans. Their information may, still, reflect a greater extent of shallow water euxinic basins rather than an entirely euxinic ocean floor. Boosted data for the concentration of redox-sensitive elements in carbonaceous shales and more than data for the isotopic limerick of Mo and peradventure of Cu in carbonaceous shales will probably clarify and perhaps settle the questions surrounding the redox land of the deep ocean during the Mesoproterozoic (Poulson et al., 2006).

Figure 14. The concentration of Mo, U, and Re in carbonaceous shales: McArthur Basin, Australia, ane.6   Ga; Finland and Gabon, 2.15–2.0   Ga; South Africa, ≥2.3   Ga.

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Sedimentary: Phosphates☆

Shamim A. Dar , ... Due west.D. Birch , in Reference Module in World Systems and Environmental Sciences, 2017

Distinguishing Characteristics

Phosphate-rich sedimentary rocks may occur in layers ranging from sparse laminae a few millimeters thick to beds a few meters thick. Some phosphate successions such as the Phosphoria Formation of the Idaho-Wyoming expanse may reach several hundred meters in thickness, although such successions are not composed entirely of phosphate-rich rocks. Phosphorites are generally interbedded with shales, cherts, limestones, dolomites, and, more than rarely, sandstones. Phosphatic rocks commonly grade regionally into nonphosphatic sedimentary rocks of the same age. Phosphorites have textures that resemble those in limestones. Thus, they may be made upward of peloids, ooids, fossils (bioclasts), and clasts that are now composed of apatite. Some phosphorites lack distinctive granular textures and are composed instead of fine, micrite-like, textureless collophane. The phosphatic grains may incorporate inclusions of organic thing, clay minerals, silt-size detrital grains, and pyrite. Peloidal or pelletal phosphorites are particularly common; oolitic phosphorites are somewhat less so. Phosphatized fossils or fragments of original phosphatic shells are important constituents of some deposits. Most phosphorite grains are sand size, although particles greater than 2  mm may be present. These larger grains, referred to as nodules, can range in size to several tens of centimeters. Considering the textures of phosphorites have such close resemblance to those of limestones, some geologists suggest using modified limestone classifications to distinguish different kinds of phosphorites. For example, Slansky (1986) advocates using a classification system based to some extent on Folk's (1962) limestone classification, and Cook and Shergold (1986b) and Trappe (2001) advise adapting Dunham'due south 1962 carbonate classification (modified by Embry and Klovan, 1971) for utilise in describing phosphorites. Using these modified classifications thus yields names such equally wackestone phosphorite (Cook and Shergold) and phosdast wackestone (Trappe).

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The Oceans and Marine Geochemistry

T.I. Eglinton , D.J. Repeta , in Treatise on Geochemistry, 2003

6.06.2.two.1 Terrigenous organic matter fluxes to the oceans

OC fluxes from sedimentary rock weathering on land are non well constrained but on geological timescales are believed to match OC burial in sediments (Berner, 1989). Superimposed on this background of relict OC from sedimentary rock weathering are fluxes associated with terrestrial primary production. The global rate of cyberspace terrestrial photosynthesis is estimated to be in the range of 60   Gt   yr⁻¹ (Mail service, 1993). Approximately two-thirds of the resulting full plant litter is oxidized rapidly to CO₂ (Post, 1993), while the residuum enters the soil bike and is subject to further oxidation. Organic thing pools inside soils exhibit different reactivities and turnover times that range from decades to millenia (Torn et al., 1997). Over geologic timescales, notwithstanding, the pervasive and continuous oxidative degradation and leaching and erosion processes on the continents result in picayune long-term storage of organic thing on the continents (Hedges et al., 1997). Withal, some fraction of this terrestrial (vascular plant-derived) OC and sedimentary stone-derived (relict) OC escapes oxidation and is delivered to the oceans. The delivery of terrigenous OC to the oceans is primarily via riverine or atmospheric (eolian) processes.

Riverine fluxes. Approximately 0.two   Gt each of dissolved and particulate OC are carried from land to bounding main annually by rivers (Ludwig et al., 1996). Much of this riverine organic thing appears to be soil derived based on its chemical characteristics (Meybeck, 1982; Hedges et al., 1994), although autochthonous sources may be important for the dissolved fraction (Repeta et al., 2002). It is now recognized that, on a global basis, riverine belch is dominated past depression-latitude tropical rivers. This not only includes major systems such as the Amazon, and Congo, only too includes the numerous smaller rivers draining mountainous tropical regions (Nittrouer et al., 1995), about notably in Papua New Guinea and other parts of Oceania, which are estimated to account for well-nigh 50% of the global flux of river sediment to the oceans (Milliman and Syvitski, 1992). During the present-day high ocean-level stand up, much of the particulate OC associated with riverine discharge is trapped and buried on continental shelves (Berner, 1982; Hedges, 1992). However, some rivers discharge much of their terrestrial OC load beyond the shelf due either to turbidity flows downward submarine canyons (e.one thousand., Congo, Ganges, Brahmaputra), to the presence of a narrow shelf (eastward.chiliad., on the eastern flank of Papua New Guinea), or the influence of water ice-rafting as an additional style of sediment entrainment and export on polar margins (e.thou., Macdonald et al., 1998).

Eolian fluxes.Eolian fluxes of organic matter from land to body of water are much less well constrained than riverine inputs. They have been estimated to be <0.1   Gt   yr⁻¹ (Romankevich, 1984). While these flux estimates imply lesser importance of eolian inputs compared to riverine OC contributions, this fashion of delivery may exist significant in a regional context. In item, marine locations downwind from major dust sources (principally in eastern Asia and western Africa) are influenced profoundly by eolian inputs of OC and other detrital components. In addition, eolian transport can deliver terrigenous materials to remote locations of the oceans, far from the influence of rivers. For such regions (east.g., primal equatorial Pacific Ocean) eolian OC fluxes may be important both in terms of POM in the water cavalcade and underlying sediments (Gagosian and Peltzer, 1986; Zafiriou et al., 1985; Prospero et al., 2003; Eglinton et al., 2002).

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Microbial transformations of organic matter in black shales and implications for global biogeochemical cycles

Due south.T. Petsch , ... T.I. Eglinton , in Geobiology: Objectives, Concepts, Perspectives, 2005

2 Dirt City, Kentucky: the field site

Late Devonian marine sedimentary rocks outcrop equally a five- to 10-km-wide ring extending ~300 km around the edge of the Jessamine Dome in primal Kentucky. Among these rocks are successions of finely laminated pyritic blackness shales upwards of 30 m thick, termed the New Albany, Chattanooga and Ohio Shales (in the w, s, and northeast portions of Kentucky, respectively). In Powell Canton, approximately 50 km ESE of Lexington (KY), a suite of road cuts through these Late Devonian black shales is located near the town of Clay City. 1 detail route cut has been the focus of much attention due to the obvious and well-developed weathering contour exposed ( Fig. 1). A well-developed, vertical weathering front end into this black shale was exposed ~twoscore years ago (private landowner, personal communication) when the hillside was excavated to overstate the roadway and provide fill up for construction. The exposed weathering profile exhibits a distinctly lighter brown color and a friable unconsolidated physical texture compared with unweathered rock in the center of the roadcut. This weathered material is not the outcome of deposition of soil from upslope because private stone strata are discernibly continuous in horizontal layers extending from unweathered rock through the weathering profile. Our efforts have focused on evaluating the chemical and microbiological variations associated with color and textural changes within this and other black shale weathering profiles.

Fig. one. Exposure of ~iv g thick black shale weathering profile in roadcut through 50. Devonian black shale nigh Clay Urban center, Kentucky, USA.

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Lime

Robert L. Zimdahl , in Vi Chemicals That Changed Agriculture, 2015

Agronomical Lime

Limestone is a sedimentary stone composed of different crystal forms of calcium carbonate (CaCO₃). Information technology is about 10% of all sedimentary rock. Near too contains skeletal fragments of marine organisms. Historic uses of limestone included mortar and pulverized limestone used to neutralize acidic soils. Burnt lime (CaO, quicklime) is made by the thermal decomposition of naturally occurring things that contain CaCO₃ (e.k., limestone, seashells), When calcium carbonate is heating above 825°C (1517°F), calcination or lime-burning liberates carbon dioxide and producing quicklime. ¹

CaCO₃(s) → CaO(s)   +   CO_ii(g)

The master agronomical use was and is to raise soil pH. There is no reliable record of when lime was first used as a soil amendment. Lime mortar dated 15,000 to 7000   years BCE has been recovered from terrazzo floors in Turkey. Limestone was used to build portions of the Great Wall of Mainland china and the Great Pyramid of Giza, Arab republic of egypt. Lime has many other uses:

calcium supplement for animal feeds;

construction aggregate every bit a roadway base;

manufacturing of some kinds of glass;

additive to newspaper, plastics, pigment, tiles, and other materials as both white paint and a cheap filler;

toothpaste;

nutrient supplement equally a source of calcium; and an

additive to some pharmaceuticals and cosmetics.

The world'southward soils vary in color, texture, structure, and chemical, physical, and biological composition. Information technology is reasonable to claim that soil is one of the most important things on the earth. Information technology is unquestionably essential to agriculture. Soil is the medium in which food is grown. It is non, as many think, just dirt. Soils are non uniform, although they may appear to be, especially at the local level, only in reality they tin can be very different within a few feet. The diversity of soils is the result of five soil-forming factors. ² Each spans a continuum, which results in their being thousands of unlike soils in the earth.

1.

Climate. The amount, intensity, and timing of precipitation influence soil formation. Seasonal and daily changes in temperature bear upon moisture, weathering, and leaching. Wind (erosion) redistributes sand and other particles. Seasonal and daily changes in temperature decide rainfall's role, its effects, the rate of biological action and chemic reactions, and the resulting vegetation.

2.

Biology. Plants, animals, microorganisms, and humans, independently and collectively, affect soil formation. Establish roots open channels in soil, taproots penetrate deeply gristly roots near the surface easily decompose and add organic matter. Animals and microorganisms mix soils. Microorganisms bear upon chemical exchanges betwixt roots and soil. Humans mix soil, often extensively, and grow the plants they want.

3.

Mural position, topography. Slope and directional orientation affect soil moisture and temperature. Steep slopes facing the sun are warmer. Slopes may lose topsoil every bit they course and be thinner than near level soils that receive deposits from areas above.

4.

Parent cloth. Nearly soil has been created from materials that have moved in from miles or only a few feet away. Loess, an aeolian (windblown) sediment, is common in the good soils of the midwestern United States and some parts of China. It is formed by the gradual accumulation of air current-diddled silt (xx–50   μm particles) and has 20% or less clay. Loess soils are typically most neutral pH.

5.

Time. Soil formation is continuous. Over time, soils exhibit features that reflect the other forming factors.

The primary, if non the just, reason agricultural lime (CaCO_three) is added to soil is to raise soil pH toward neutrality (7.0), thereby reducing soil acidity, increasing food availability, and permitting successful growth of many crops that are pH sensitive. pH is an abbreviation for potential hydrogen. Information technology is the negative logarithm (base of operations x) of the reciprocal of the hydrogen ion concentration in gram atoms/liter of water. It indicates hydrogen ion activeness. It defines acerbity or basicity (pH aq) of a solution on a scale from 0 to 14. Vii is neutral; below is acidic and above is basic. That ways that for each unit pH increases or decreases, basicity or acidity changes past x times. A pH of 5 is x times more acidic than pH 6 and 100 times more than acidic than pH vii.

Because pH controls many soil chemic processes and the chemic form of nutrients, it may enable or inhibit their uptake. Acid soils (below pH v.5) have greatly reduced microbial activity, but release many nutrients, notably fe, which is much less available above pH vii.5. Soils with pH lower than four.6 are too acidic for most plants. Many soils are naturally calcareous ³ (pH in a higher place 7). In some cases, sulfur can be added to make them more acidic through germination of sulfuric acid (H₂And so₄), hydrogen sulfite (HSO_three), and hydrogen sulfide (H₂S). At a pH in a higher place seven, carbonates and oxides are formed and tin can react with many metallic food elements (e.g., iron, copper, molybdenum), which renders them insoluble and therefore unavailable to plants. Almost crops practice not grow well in acidic soil or soil with a pH above eight. Raising the pH of acidic soil improves plant growth, may improve water penetration, and reduce aluminum toxicity. Lime is a source of calcium and magnesium for plants. Because information technology is high in calcium, it can also be beneficial to os growth of foraging animals.

Soils in high rainfall areas go acidic through leaching. Crop growth and livestock grazing remove essential nutrients over time and soil may gradually go acidic. Chemical fertilizers required to achieve maximum yield are major contributors to soil acidity. Therefore, liming acidic soil is essential to attain maximum yield of food crops grown in acidic soils.

In areas of extreme rainfall and high temperature, ⁴ clays and humus may be leached away, which further reduces soil'southward buffering capacity (i.due east., resistance to changes in pH). In low-rainfall areas, unleached calcium may heighten pH to 8.v and if exchangeable sodium levels are loftier, soil pH may accomplish x. Higher up pH 9, most food crops will non grow or their growth and yield will be severely reduced. Loftier pH also results in low micronutrient mobility and availability.

The desirable (optimum) pH range for most food crops is 5–7. Every crop has an optimum pH range inside which production potential peaks. No important food crops have an optimum pH less than 5. Rye, oats, and lupins are acid-tolerant. The optimum for corn and soybeans is 5–7.five. The most important food crops (beans, rice, and wheat) take an optimum pH between 5 and seven. They are acrid sensitive. The 10 most important nutrient crops—the plants that feed the earth—all grow best between pH five.v and 6.5 (see Table 3.1). Potatoes, an exception, grow when soil pH is 4.8–5.5, although they grow well above pH 5.5. Many plants, but not the important food crops, accept adjusted to thrive at pH values outside the optimum range, but do best within the optimum.

Tabular array 3.i. Optimum pH Range for Some Food Plants ^a

pH Range
5.0–five.five	v.8–6.v	six.5–7.0
Blueberries	Edible bean X	Alfalfa
Irish potato X	Cassava X	Barley X
Sweet spud X	Corn/maize X	Carmine
Grasses X	Grapes
Millet 10	Soybean Ten
Oats	Sugarbeet
Rice 10	Wheat X
Rye
Some clovers
Sorghum X
Love apple
Watermelon

a

12 of the world'southward about of import food crops are indicated with X.

Well-nigh 30% as much phosphorus is available when pH is below 6 versus higher up 6.5. Nitrogen and potassium become less bachelor beneath vi. Availability of both decreases by about 30% at pH v.five and seventy% at v. Below pH 5, soil nitrogen, phosphorus, and potassium concentrations may be adequate to support constitute growth, but, considering of formation of insoluble minerals, they are unavailable and of footling use to plants. In contrast, atomic number 26, copper, manganese, and zinc are most available when pH is acidic. It seems contradictory, but when pH is around four, other food levels may be high considering lack of plant uptake leads to their accumulation in soil.

The most constructive way to raise pH is to apply good-quality, finely ground agronomical lime. Lime'southward calcium and/or magnesium carbonate content and the fineness of grinding determine quality. Finely footing lime will enhance soil pH more apace. Incorporating it in soil is more efficient than surface application. Change in pH is more rapid when soil is moist from irrigation or rainfall; h2o is required for the chemical reaction. Action (pH change) may have a year or more in dry soil. However, with good soil moisture, a response may be observed within weeks if pH is extremely depression. When pH is depression, it is important to apply lime immediately later the growing season to allow sufficient reaction fourth dimension before the adjacent crop is planted.

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Meteorites, Comets, and Planets

R.N. Clayton , in Treatise on Geochemistry, 2003

one.06.2.2.1 Archaic nebular materials

Chondritic meteorites are sedimentary rocks composed primarily of chondrules, typically sub-millimeter-sized spherules believed to have been molten droplets in the solar nebula, formed past melting of dust in a brief, local heating outcome. During the high-temperature stage, with a duration of some hours, the aerosol could interact chemically and undergo isotopic exchange with the surrounding gas. Thus, isotopic analyses of individual chondrules can provide information about both the dust and gas components of the nebula. Figure ii shows oxygen isotopic compositions of chondrules from the three major groups—ordinary (O), carbonaceous (C), and enstatite (E) chondrites—and one minor grouping—Rumuruti-type (R)-chondrites. The iii-isotope graph has ii useful properties: (i) samples related to one another by ordinary mass-dependent isotopic fractionation lie on a line of gradient=0.52, like the line labeled terrestrial fractionation (TF), and (ii) samples that are two-component mixtures lie on a straight line connecting the compositions of the end-members.

Effigy two. Oxygen isotopic compositions of chondrules from all classes of chondritic meteorites: ordinary (O), enstatite (E), carbonaceous (C), and Rumuruti-type (R). The TF line and carbonaceous chondrite anhydrous mineral (CCAM) line are shown for reference in this and many subsequent figures. Equations for these lines are: TF—δ¹⁷=0.52δ^xviii and CCAM—δ¹⁷=0.941δ¹⁸−4.00 (sources Clayton et al., 1983, 1984, 1991; Weisberg et al., 1991).

Figure 2 shows that chondrule compositions exercise not lie on a mass-dependent fractionation line, thus indicating isotopic heterogeneity in the nebula. Chondrules from different chondrite classes occupy unlike regions of the diagram, and for each class, they form nearly-linear arrays that are considerably steeper than a mass-dependent fractionation line. For comparing, information technology tin exist noted that analogous three-isotope graphs for iron in diverse meteorite types are strictly mass dependent, and evidence no evidence for nebular heterogeneity (Zhu et al., 2001).

Some other group of primitive objects with a direct link to the solar nebula are the calcium–aluminum-rich inclusions (CAIs), that range in size from a few μm to >1 cm (encounter Chapter 1.08). They are constitute in all types of primitive chondrites simply are rare in all but the CV carbonaceous chondrites. Their bulk chemical compositions stand for to the almost refractory v% of condensable solar matter (Grossman, 1973). They may represent direct condensates from the nebular gas, followed, in many cases, past further chemic and isotopic interaction with the gas. Their radiometric ages accept been measured with high precision (Allègre et al., 1995), and indicate solidification earlier than any other solar arrangement rocks (excluding the presolar dust grains). Thus, the oxygen isotope abundances in CAIs may provide the best guide to the composition of the nebular gas.

CAIs exhibit a specific, characteristic pattern of oxygen isotope abundances. Within an individual CAI, different minerals have different isotopic compositions, with all data points falling on a straight line in the three-isotope graph. This behavior is illustrated in Figure three, based on analyses of physically separated minerals from the Allende (CV3) carbonaceous chondrite (Clayton et al., 1977). Each analysis represents a large number of grains. Figure four shows data obtained by ion microprobe analysis, where each point represents just one grain (Aléon et al., 2002; Fagan et al., 2002; Itoh et al., 2002; Jones et al., 2002; Krot et al., 2002). The line labeled "CCAM" is the same in Figures 3 and 4. Although the ion microprobe data have larger analytical uncertainties, leading to greater besprinkle in the data, it is clear that the same pattern exists at both the microscopic and macroscopic level. Inside individual CAIs, the sequence of isotopic composition, in terms of ^sixteenO-enrichment, is spinel≧pyroxene>olivine>melilite=anorthite. A straight line on the iii-isotope graph is indicative of some sort of ii-component mixture. The fact that the range of variation in the private-grain studies is about the same equally the range in the bulk-sample studies shows that the stop-members do non lie much across the observed range of variation. All studies, as of early on 2003, reveal an ¹⁶O-rich end-fellow member nigh −45‰ for both δ¹⁷O and δ¹⁸O, oftentimes represented by spinel, the almost refractory of the CAI phases. The most obvious interpretation is that this end-member represents the composition of the main nebular gas, from which the CAIs originally condensed. Subsequent reaction and isotopic substitution with an isotopically modified gas could then yield the observed heterogeneities on a micrometer to millimeter scale (Clayton et al., 1977; Clayton, 2002).

Figure 3. Oxygen isotopic compositions of physically separated minerals from several Allende CAIs. These points were used to define the CCAM line (source Clayton et al., 1977).

Figure 4. Ion microprobe oxygen isotope analyses of single grains in several carbonaceous chondrites. Analytical uncertainties are typically about ±2‰. The CCAM line is shown for reference sources: are noted in the figure.

Another argument that the ¹⁶O-rich end-member was a ubiquitous component of archaic solids is that it is found in many different chemical forms (unlike minerals) in many classes of meteorites: CAIs and amoeboid olivine aggregates (AOAs) from Efremovka (CV3) (Aléon et al., 2002; Fagan et al., 2002), AOA from a CO chondrite, Y 81020 (Itoh et al., 2002), and CAI and AOA from CM and CR chondrites (Krot et al., 2002). This isotopic limerick tin as well serve every bit an cease-member for the chondrule mixing line in Figure 2.

If the ¹⁶O-rich composition is indeed the isotopic composition of the primordial solar nebula, the consequences for solar system formation are profound. As noted above, materials with the ¹⁶O-rich composition are ubiquitous, simply they are also rather rare, never amounting to more than a few percent of the host meteorite. The implication is that all the other material in the inner solar system has undergone some procedure that changed its ¹⁶O-abundance by 4–v%. This must accept been a major chemic or physical process that must go out evidence in forms other than the isotopic composition of oxygen.

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Cetacean Fossil Record

R. Ewan Fordyce , in Encyclopedia of Marine Mammals (Second Edition), 2009

II Occurrence, Environment, and Historic period

Fossil cetaceans occur in sedimentary rocks. Originally, remains accumulated in mud, silt, sand, or gravel which, every bit flesh decayed, was cached and turned to rock through compaction and/or deposition of cementing minerals. Sedimentary rocks are recognized every bit discrete formations (genetically unified bodies of strata), and are named formally, eastward.g., the Calvert Formation, Maryland. Marine mammals come from strata including sandstone, mudstone, limestone, greensand, and phosphorite, most of which are marine rocks now exposed on state. Rare fossils have been recovered from the sea flooring. Because broadly similar rock types may form at dissimilar times and places, sedimentary rocks must be dated to establish their fourth dimension relationships.

Two correlated timescales, relative and absolute, are used for the fossil record. The relative timescale has named intervals (epochs; Fig. 1) in an agreed international sequence: Eocene, Oligocene, Miocene, Pliocene, and Pleistocene. These epochs are usually subdivided into early on, middle, and tardily. Stages (due east.g., Aquitanian of Fig. one) may provide effectively subdivision. Typically, distinct historic period-diagnostic fossils are used to recognize time intervals. The almost reliable dates are based on oceanic microfossils with short-fourth dimension ranges, such as foraminifera, which allow correlation between ocean basins. Because of compounded errors of long-altitude correlation, ages are rarely accurate to within ane million years, and many fossils tin can exist placed only roughly inside a phase. Beyond the relative timescale, absolute dates in millions of years are needed to understand rates of processes in phylogeny (involving, east.g., molecular clocks) and in geology (involving, e.thou., rates of sediment accumulation or of climatic change). Absolute dates are usually obtained from radiometric analysis of grains of volcanic rock interbedded with fossiliferous strata.

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The Crust

A.I.S. Kemp , C.J. Hawkesworth , in Treatise on Geochemistry, 2003

3.11.v.3 Relevance for Crustal Differentiation

The fractionated igneous and sedimentary rocks in Figure twenty have Rb/Sr and Eu/Sr ratios that are much higher than those in current estimates for the upper continental crust. The latter can be constrained by the strontium isotope ratios of continental run-off (∼0.712), and its model neodymium age (∼1.viii   Ga). According to this method, a minimum time-integrated upper crustal Rb/Sr ratio of 0.xiv is indicated.

Another striking feature of the information in Figure 20 is that the upper, lower and bulk continental crust compositions all have similar European union/Sr ratios and thus define a singled-out, nigh-vertical tendency that is separate from the igneous and sedimentary arrays. At that place are several potential interpretations for this. Information technology might but reflect the methods employed to judge the upper and lower crustal averages. The upper crust represents a mixture between sediments and intermediate to felsic igneous rocks (annotation how close the upper crustal limerick plots to the reference suite array in Effigy 20), whereas the lower chaff inevitably combines xenolith data from both intraplate and destructive margin settings, not necessarily in representative proportions; the bulk crustal composition is constrained to lie betwixt these extremes. The dispersion in crustal compositions on Effigy 20 could therefore be synthetic and petrogenetically meaningless. Alternatively, if the chemical variation between the crustal components results from the differentiation or "unmixing" of a bulk starting composition, the data suggests that neither igneous fractionation nor weathering processes can be wholly responsible for such differentiation, since strongly increasing Eu/Sr is a signature of those processes. In any instance, it is clear from Effigy 20 that the upper crustal composition has sufficiently depression Rb/Sr and Eu/Sr to preclude a pregnant contribution from the continental sediments, in contrast to the implications drawn from the granite and upper crust trace-element patterns in Effigy 5.

I resolution to this conundrum could exist related to a shift in the oxidation land of europium (i.e., the proportion of Eu^two+ to Eu³⁺) either through fourth dimension, or in different tectonic settings (Carmichael, 1991). The marked increment in Eu/Sr in igneous suites indicates that D_Sr was much larger than D_{European union} in the fractionating plagioclase. All the same, D_Eu in plagioclase is sensitive to oxygen fugacity (f _Otwo ), and it is very low in oxidizing weather condition where europium exists as Eu³⁺ (Drake, 1974). The igneous rocks plotted in Figure 20 can be inferred to have formed nether reasonably oxidizing conditions, since many of them originate above subduction zones, where the mantle has been modified by the introduction of hydrous fluids from the subducted slab, or they derive from, or accept interacted with, recycled sedimentary rocks in the deep chaff.

Yet, in reducing atmospheric condition, D_Eu increases until, at f _Oii ∼10^–12.5bars, it approaches values similar to that of D_Sr. Plagioclase formed under these conditions will therefore not fractionate europium from strontium, and its removal results in vertical arrays on Figure 20, as is shown past the continental chaff. This raises the intriguing possibility that the differentiation of the continental chaff was primarily achieved under relatively reducing weather, such as existed in the Archean menses where a CO₂-rich atmosphere prevailed, or in intraplate settings. The latter would be marked by magmas with distinctive trace-elements patterns, and in particular no negative Nb–Ta anomalies. Such magmas did contribute to the generation of new crust, they cannot be the dominant component, and so the apparent lack of Eu/Sr fractionation in the crustal compositions may largely reflect processes in the Archean. Independent evidence for a reducing surround at this time includes the presence of banded iron formations, uranium placer deposits, high Th/U ratios in igneous rocks, non-mass dependent sulfur isotope fractionations (Farquhar et al., 2000; see also Chapter four.04) and the lead isotope limerick of the mantle (see Elliott et al., 1999). Under such reducing weather condition, intracrustal melting can generate the observed differentiation of the continents, the important bespeak being that the residual plagioclase independent substantially more europium than that of the magmatic reference array in Figure 20. Recycling of pocket-sized amounts of the residues of melting, approximated past the estimated lower crustal composition, can explain the displacement of the bulk chaff from the magmatic array on Effigy 20. The fundamental implication from this reasoning is that melting and weathering processes operating in the outer role of the post-Archean Earth have contributed relatively piffling to the majority differentiation of the continental crust, consistent with its average age of ∼ane.viii   Ga.

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