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Reading material for M.Sc (Geology) First Semester

under CBCS

This reading material has been prepared by Dr H N Sinha, Dr Bipin Kumar, Dr B A

Kumar and Y N Jha of the Department of Geology, VBU, Hazaribag strictly for

restricted circulation.

Students are advised to elaborate the reading materials by adding other

information from the journals, books , reference books etc.

References

1.Structural and field geology by James Geikie

2. Text book of Invertebrate Palaeontology by Henry Woods

3. An introduction to igneous and metamorphic petrology by Ohn D. Winter

4. Folding and fracturing of rocks- J G Ramsay

5. Plate tectonics and crustal evolution-K C Condie

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Sedimentry Geology

Facies- Lateraland vertical lithological variation

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Major environment of deposition

Sedimentary Basin

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A depression on the Earth’s surface where sediments have been deposited with the prospect of long-term preservation. To produce a sedimentary basin there must be a mechanism to drive. the depression of the Earth’s surface. There must also be a supply of sediment to at least partially fill the depression. Sedimentation may result from either chemical precipitation or clastic deposition. Some basins on or adjacent to continents have over 10 km of sedimentary fill whilst others are relatively “empty”. Subsidence of Basin Crustal extension – Rifts: evolution to passive continental margins – Strike-slip pull-apart • Thermal relaxation/ cooling – Post rift adjustment to thinned crust – Sag basins • Crustal loading – Foreland basins; crustal loading by nappes – Sediment loading • Others – Growth synclines – Piggy-back basins – Salt withdrawal Minerals in microscope

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Cleavage

A. Augite: Section orthogonal to c axis. The {110} cleavage planes form

angles of 87°and 93°.

B. Hornblende: Section orthogonal to c axis. The {110} cleavage planes

form angles of 56°and 124°.

C. Kyanite: Two sections approximately orthogonal to c show the typical

pattern of very good cleavage {100} and distinct cleavage {010}.

D. Sillimanite: The section orthogonal to c shows the good cleavage {100}.

Fracture

A. Perlite: Due to quenching of the glassy material (obsidian) concentrically

curved tension cracks developed.

B. Pyrope: Radial cracks emanating from coesiteinclusions now largely

transformed to quartz. The cause of the fracturing is the volume increase from the

coesite-quartz transformation, resulting in increased pressure imposed on the

garnet host.

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C. Garnet: Fracture planes oriented parallel to narrow-spaced jointing in a basic

granulite.D. Nepheline: Irregular tension cracks caused by rapid cooling.

Twinning

Twins are generated through crystal-structure-controlled intergrowths of two or

more individual crystal segments with a defined symmetrical relationship.

Twinning can also result from deformation (as in calcite). The individual parts of a

twinned mineral are intergrown such that they either mirror each other's orientation

(the mirror plane being the twin plane), or they are rotated against each other by a

specific angle (the rotation axis being the twin axis), or both. The twin interface

commonly corresponds to the twin plane. For many mineral species twinning is an

important property for identification. There are different kinds of twinning such as

contact twins, penetration twins, simple twins, multiple twins, polysynthetic (or

lamellar) twins (Fig. 3.3-1). In thin section, twinning is commonly easily

recognised under crossed polarizers if the mineral is anisotropic. The individual

parts of twinned crystalsshow different brightness and interference colour, and on

turning the microscope stage different extinction positions are revealed. There are

exceptions, however, as not all types of twins can be recognised under the

microscope. If the indicatrix orientation of the individual parts of twinned crystals

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is identical, they are indistinguishable under crossed polarizers (e.g., the most

abundant quartz twins have a twin axis parallel to c, which means the indicatrices

are in parallel alignment; thus, the twins go into extinction simultaneously).

FELSPAR GROUP Felspar is a general term for a number of closely related minerals which play a very important role as rock-formers. They are the chief constituents of most eruptive rocks, and are met with likewise more or less abundantly in many crystalline schists. They vary in colour, but are usually grey, white, or reddish; occasionally, however, they show yellow, green, or blue tints. As rock-constituents they frequently assume the form of tabular crystals, or appear as long rods or rectangular lath-shaped bodies. All are characterised by two well-marked sets of cleavage-planes (at, or nearly at, right angles) which show usually a glassy or pearly lustre ; further, all have approximately the same hardness (6 to 7), and specific gravity (2-54 to 2-76). Chemically, they are silicates of aluminium with either potassium, sodium, or calcium, or several of these together. Hence we have potash

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felspar, soda felspar, lime felspar, soda-lime felspar, etc. These felspars so closely resemble each other that it is often hard or even impossible to distinguish one from another by the unassisted eye. This, of course, is especially the case when the crystals are small. Usually, however, the particular class or series to which a felspar belongs can be determined by examination in thin slices under the microscope. Two series of felspars are recognised one of these crystallising in monoclinic and the other in triclinic forms. The monoclinic series includes Orthoclase and Sanidine, while the triclinic class is represented by Microcline, Anorthoclase, and Plagioclase the last-named forming a group of felspars which are all more or less closely related, and often hardly to be distinguished from each other without careful microscopical or chemical examination. As a group they are more or less readily differentiated from the monoclinic felspars by the inclination of theircleavage-planes in the monoclinic felspars these planes being directed at right angles to each other, while in the triclinic group referred to they are not at right angles. Hence we have two series of felspars namely, (a) Orthoclase, with rectangular cleavage, and Plagioclase, with oblique cleavage. If felspars always assumed their external crystalline form and were of sufficient size,' it would not be hard to distinguish between orthoclase and plagioclase. As rock - constituents, however, they are often so unsymmetric in shape, or occur as granules so small in size, that the geologist must have recourse to other differentiating characters to distinguish between one felspar and another. Under the microscope, the plagioclase felspars can usually be recognised by their "multiple twinning." A crystal or crystalline granule having this structure appears as if it were composed of a series of parallel plates or lamelke, which show alternately lighter and darker tints when examined between crossed nicols. Thus a section of the mineral, if cut in a proper direction, exhibits a banded or striped appearance. THE AMPIIIBOLE AND PYROXENE GROUP The Amphiboles described here are calcium-magnesium silicates, some being rich in aluminium and iron, others containing little or no trace of either. Of less frequent occurrence are varieties rich in soda. When crystallized they appear as prisms; but they show a marked tendency to assume fibrous and radiated forms. Their specific gravity ranges from 2-9 to 3-5, and their hardness is between 5 and 6. They are usually fusible, more particularly when rich in iron. Amphiboles crystallise in monoclinic, triclinic, and orthorhombic forms, but only the first are

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important rockformers. The monoclinic non-aluminous amphiboles are usually lighter in colour than those rich in aluminium and iron. The most commonly occurring representatives of the nonaluminous class are Tremolite and Actinolite. Tremolite is white, grey, or light green in colour, aud occurs usually in the form of long blade-shaped crystals, striated longitudinally : or it assumes the appearance of thin fibrous crystals radiating from a centre. The crystals have a pearly or silky lustre. This mineral is a constituent of some schistose rocks ; it occurs not uncommonly in crystalline limestone (marble) and dolomite near their point of contact with plutonic rocks. Now and again it is met with as an alteration-product 13 olivine- rocks and serpentine. Actinolite differs from tremolite in containing a considerable percentage of iron ; hence it is generally light or dark green in colour. It usually occurs as long thin columnar crystals and radiate aggregates. It is a common ingredient of many crystalline schists, where it is frequently associated with talc, chlorite, and epidote. In eruptive rocks (as in saussurite-gabbro) it is often met with as an alteration-product. Tremolite and actinolite sometimes assume forms so fibrous that they can be readily separated into thin, soft, cotton-like, or silky threads, and are then known as Amianthus or Asbestos. The fibres are often matted together so as to form felt-like substances, termed "mountain-leather," mountain-cork," etc. Most of the asbestos .of commerce, however, is not amphibole, but fibrous serpentine (chrysottle). Of the monoclinic aluminous amphiboles, by far the most important is Hornblende. This mineral has much the same composition as actinolite, but contains a notable percentage of alumina. Two varieties are recognised namely, Common Hornblende and Basaltic Hornblende. The former is dark leek-green to black in hand specimen, but is green in transmitted light in thin section. The crystals generally show an elongated prismatic habit, but sometimes appear as blade-like, fibrous, radiating aggregates. It is an essential constituent of many plutonic rocks (syenite, diorite, hornblendic granite), occurring now and again also as an accessory ingredient in gabbro. It is a frequent constituent of crystalline schists (amphibole-schist, hornblende-gneiss). It commonly alters to chlorite or epidote, or may be still further broken up by weathering and reduced to the condition of ferruginous clay. Basaltic Hornblende is generally brownish-black to pitchblack, but when viewed in thin sections it usually shows a deep brown or reddish-brown colour. The crystals are commonly short, stout prisms, and are frequently well formed. The mineral occurs as a macroscopic and microscopic ingredient of certain trachytes, andesites, and basalts the larger crystals often showing corroded blackened borders the result of magmatic resorption. The only other monoclinic amphibole that need be

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mentioned is Smaragdite a peculiar grass-green fibrous lamellar form, approaching actinolite in composition, but containing some alumina. It occurs in eclogite, where it forms parallel growths with omphacite a similar green pyroxene mineral. The Pyroxenes have much the same chemical composition, hardness, and specific gravity as the amphiboles, and crystallise like them in monoclinic, triclinic, and orthorhombic forms. They differ amongst themselves as regards fusibility those containing much iron being usually more fusible than the less ferriferous varieties. The monoclinic forms are divisible, like the corresponding amphiboles, into non-aluminous and aluminous types. The non-aluintnous pyroxenes are mostly light- coloured white or, more commonly, some pale shade of green. They occur chiefly in crystalline schists and in crystalline limestones and marbles, but are not such important rock-formers as the corresponding lightcoloured amphiboles. Their alteration-products are usually talc or serpentine and carbonates. Of the aluminous pyroxenes the most notable is Augite. it crystallises in prismatic forms, which are often twinned. As rock-constituents the crystals frequently have their edges and angles rounded off. Augite is dark brown to black, but in thin sections may be almost colourless or show various shades of brown or yellow, and sometimes of green. It is often altered into an aggregate of chlorite, scattered through which may be minute granules of epidote, calcite, and quartz ; or it may be still further changed to a mixture of limonite, quartz, and carbonates. Sometimes it is replaced by secondary hornblende. It is an essential constituent of such basic rocks as basalt, dolerite, etc., but occurs as an accessory ingredient of many other eruptive rocks. Diallage is a brownish, grey, or greenish variety of augite, which rarely assumes a crystalline form, and has a lamellar structure. Numerous platy inclusions occur along the cleavage-planes, so that the mineral exhibits a submetallic lustre on broken surfaces. It is an essential constituent of gabbro, and occurs also as an occasional ingredient of serpentine and olivine-rocks; but appears never to be met with in effusive igneous rocks. THE MICA GROUP The micas, as rock-formers, mostly occur as thin plates and scales, the surfaces of which show a pearly to submetallic lustre. Usually these plates are irregular in shape, but now and again they are six-sided. The micas, however, are really monoclinic with pseudo-hexagonal symmetry. The cleavage is perfect, all micas being readily split up into exceedingly thin, transparent, and elastic leaflets. They

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are all rather soft (2-5 to 3 in the scale), and the specific gravity ranges from 2-7 to 3-2. They are essentially silicates of aluminium and potassium (or sodium), some kinds containing magnesium and iron. Only two micas are important rock-formers,namely, the brown to black Biotite or ferromagnesian mica, and the silver-white Muscovite or potash mica. They are essentialconstituents of many schistose rocks and of granite, and aremet with in a large number of eruptive rocks of all ages. Soft, non-elastic scales of mica are also of common occurrence in many derivative rocks, particularly in fissile sandstones. Biotite (ferromagnesian mica) is usually dark brown to black, but green and red varieties are known. It is decomposed by strong sulphuric acid ; and in nature alters readily to chlorite, with separation of iron-oxide. Not infrequently, however, biotite becomes pale through loss of iron, and then assumes a golden yellow to silver-grey colour, thus sometimes closely resembling muscovite. It is a primary or original constituent of granites, rhyolites, some syenites and diorites, trachytes, etc. In effusive rocks the scales often show blackened borders, which, as in the case of basaltic hornblende, appear to be due to the corrosive action of the igneous magma. Biotite occurs also in certain schistose rocks. Being a less durable mineral than muscovite, it is not so often met with in sedimentary rocks. In thin rock-sections under the microscope, biotite, if cut at right angles to its vertical axis (or, in other words, if the slice be parallel to the cleavage-planes), appears deep brown or deep green to black, and shows little or no change of colour when rotated above the polariser. But when the section cuts across the cleavage-planes, which then appear as a scries of parallel lines traversing the mica, as shown in and the stage of the microscope is rotated, the change of colour is strongly pronounced. The polarisation colours are very brilliant in sections showing cleavage, and cut thin enough. Inclusions are frequently numerous, mostly of apatite and magnetite, and less commonly of zircon and rutile. Muscovite (potash mica) is sometimes colourless, but usually palecoloured or silvery ; occasionally, however, it assumes a light shade of brown or green. It fuses on thin edges to a grey glass or white enamel, but is not attacked by acids, and as a rock-constituent is not so readily altered as biotite. As a primary rock-former its chief habitats are the crystalline schistose rocks (gneiss, mica-schist, phyllite), and the granites. It never occurs as an original constituent in any igneous rocks save granite, certain quartz-porphyries, and syenites. Being a mineral not readily decomposed, it frequently appears in the form of soft, worn looking, non-elastic scales in sedimentary rocks of many kinds. Although muscovite has no great range as a primary constituent of crystalline eruptive rocks, it occurs in many as a secondary ingredient the product

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of the alteration of silicates rich in alumina. Thus it often replaces such minerals as andalusite, felspar, nepheline, etc. Seen in thin sections under the microscope, muscovite is colourless or very faintly yellowish or light green. It shows no change of colour, or at most only a slight difference in the intensity of the colour, when rotated above the polariser. It polarises, however, more brilliantly than biotite. Inclusions are few. Although the micas, as rock-formers, occur most frequently in the form of scales, flakes, or plates of relatively small size, they now and again appear as large rough prisms, often tapering to a point as, for example, in limestones which have been subject to metamorphic action. Very large individuals of muscovite also are met with in the pegmatitic veins (giant granite) associated with so many granitic masses. THE OLIVINE GROUP The minerals of this group are non-aluminous silicates. The only one of importance as a rock-former is Olivine (Pendote) a silicate of magnesium and iron which crystallizes in orthorhombic forms and shows an imperfect cleavage. It has a hardness of 6-5 to 7, and a specific gravity of 3 to 4. The proportion of iron varies specimens containing very little being infusible, while those which are rich in iron are more or less readily fused. The mineral is slowly decomposed by cold hydrochloric acid with gelatinisation. It is usually yellowish-green or olive-green, has a glassy lustre, and breaks with a conchoidal fracture. As a rock-former it sometimes constitutes the whole mass or the larger proportion of a rock, as in dunites (peridotites). It is present also in many other igneous rocks more especially in those of basic composition, as certain gabbros, basalts, and felspathoid rocks. It is readily recognised in such rocks by the naked eye as granules or blebs, usually of a greenish tint with a glassy lustre, and showing its conchoidal fracture. Now and again it occurs in basalts as large granular aggregates resembling nodules, some of which may measure 5 or 6 inches across, but they are generally smaller. Forsterite, a light-coloured variety, is met with as a "contact mineral" in metamorphosed limestones. In nature, olivine alters readily to serpentine; probably, indeed, most serpentines have originated from the alteration of olivine-rocks. The finely coloured (yellow or green) transparent varieties of olivine are used in jewellery, and are known as Chrysolite and Peridote. In thin rock-slices olivine is usually almost colourless, but may show pale yellowish-green or yellowish-brown tints. In basic eruptive rocks

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it appears sometimes in good crystal forms, with lozenge-shaped or long rectangular outlines, but the outlines are frequently rounded as if from magmatic corrosion. It shows high relief, the outlines of the mineral and the cracks traversing it being strongly pronounced. It is not pleochroic, but when viewed between crossed nicol prisms show brilliant interference colours. FOSSILS Modes of Preservation of Organic Remains. Kinds of Rock in which Fossils occur. Fossils chiefly of Marine Origin. Importance of Fossils in Geological Investigations. Climatic and Geographical Conditions and Terrestrial Movements deduced from Fossils. Geological Chronology and Fossils. HITHERTO we have been concerned with rocks mainly as aggregates of mineral matter, and only a passing reference has been made to the fact that certain derivative accumulations contain fossils the remains and traces of formerly living creatures. We have seen, it is true, that some kinds of rock, such as coal and limestone, consist chiefly of the debris of plants and animals, but we have now to realise that almost every variety of derivative rock may be more or less fossiliferous, and that traces of former life have been met with, now and again, even in certain igneous and metamorphic rocks. When a plant or animal, or any portion of either, is buried in sediment, it becomes subject to decomposition. This process usually results in the destruction of all organic compounds of carbon and nitrogen, and even the harder and more durable parts undergo some change, and may eventually become disintegrated, and entirely disappear. Certain chemical changes, however, may supervene before the process of destruction is completed. In many cases, for example, carbonisation takes place various gases are given off, and the organic tissues are gradually transformed into carbon. Of mineral matter may be introduced in solution so as to fill up all the cavities of the original structures, or even to replace completely the substance of the organism. Fossils, therefore, are met with in all states of preservation. Exceptionally, the entire organism has been preserved with little or practically no change of the original substance the bodies having been protected from decomposition by the nature of the materials in which they have been entombed. As examples may be cited the carcasses of the extinct mammoth and woolly rhinoceros which, long ages ago, were sealed up in the frozen earths and ice of Northern Siberia, so that when in recent times they became exposed, owing to the gradual dissolution of the medium in which they

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had been buried, their bodies were in so fresh a state that dogs devoured the flesh. Insects, spiders, and plants have similarly been completely preserved in amber (fossil gum or resin) ; but, in most cases, these would appear to be more or less carbonised. The more common methods of preservation, however, are as follows : Incrustation. The organism under certain conditions is enveloped in a covering of mineral matter. Calcareous tufa, for example, is often precipitated upon plants growing near springs containing much calcium-carbonate. In the case of thermal waters siliceous sinter may be the incrusting substance. Vegetable and insect remains preserved in this manner are often more or less carbonised, or they may be entirely decomposed and -dissipated, leaving merely hollow moulds behind them. Carbonisation. Plant - remains and chitinous animal structures, without having been previously incrusted, frequently undergo carbonisation a deoxydising process which takes place under conditions permitting of only a limited access of air. Thus plants accumulated in marshy ground, or on the floor of lake or estuary, or buried in mud, etc., tend to undergo a kind of distillation whereby the oxygen and other gases are gradually eliminated the carbon in this way becoming concentrated. Moulds and Casts. The substance of a buried organic body may be entirely dissipated, and only a mould of it remain. Should this mould be subsequently filled with mineral matter, a cast showing the external form of the original will be produced. This is a common kind of fossilisation. Many fossil shells, for example, are simply casts, and do not contain a particle of the original substance. Whenan empty bivalve or univalve shell is enclosed in a deposit, the sediment usually at the same time fills the vacuity. Afterwards, the shell itself may be gradually dissolved and removed by percolating water. The cavity thus formed may be subsequently reoccupied by mineral matter, and in this way a perfect cast will be produced. Not infrequently, however, the space left by the shell remains unfilled, containing in its centre the stony kernel which formerly occupied the interior of the original. Should this kernel not adhere to the matrix, it will rattle, like a nut in its shell, when the specimen containing the fossil is shaken. Permeation and Molecular Replacement. Mineral matter has often thoroughly permeated an organic body, and filled up all its pores and cavities a process which has usually been preceded, accompanied, or followed by carbonisation. Not infrequently, under these conditions, the original substance itself is more or less molecularly replaced by mineral matter, with partial or perfect preservation of the internal structure. This kind of fossilisation is well illustrated by some specimens of silicified wood, the minutest

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structures of which have been so completely replaced that a slice of the specimen, viewed under the microscope, reveals as much as a section of the original wood itself could have shown. Permeation and molecular replacement may be exemplified by one and the same fossil, so that the two kinds of fossilisation are frequently hard to distinguish. An organic body which is permeated and molecularly replaced by mineral matter is a true petrifaction. In cases of true petrifaction, the replacing mineral is usually either silica (mainly chalcedony or opal) or calcium carbonate. The same substances also play the most important part in the formation of incrustations and casts, which is just what might have been expected when we remember how very widely calcareous and siliceous solutions are diffused. Other substances, however, not infrequently replace organic remains, such as the compounds of iron (pyrite, marcasite, haematite, limonite, and siderite), and, less frequently, gypsum, barytes, fluor-spar, and various metals and metallic compounds. It is not only the relics and remains of plants and animals which are termed fossils, but any recognisable trace of their former existence any impressions or tokens left behind them whether it be footprints, tracks, or trails, burrowings, castings, or coprolites, or even the markings traced on sediment by the waving to-and-fro of sea-weeds, etc. are all equally fossils. Kinds of Rock in which Fossils occur. As a rule, the best preserved fossils are met with in the finer grained sedimentary rocks, as in marls, limestones, clay and shale, and fine argillaceous or calcareous sandstones. Calcareous Rocks. Argillaceous limestones and marly shales are often highly fossiliferous, and the fossils are usually well preserved. But pure limestones, which have become more or less crystalline, frequently appear to be poor in organic remains, so that when a fresh fracture of the rock is obtained, few or no traces of any structure may be visible. On surfaces which have been for some time exposed to the weather, however, fossils not infrequently project in bold relief the limey matrix in which they are embedded offering less resistance to atmospheric action. The same phenomena characterise many dolomitic limestones. Argillaceous Sliales. Not infrequently these are rich in fossils their impervious character having doubtless tended to the preservation of the remains. Some shales, however, are very barren, or the few fossils present may be included in nodular concretions of calcium-carbonate, siderite, or other substance. Sandstones are not so frequently fossiliferous as shales, for which there are at least two reasons. First, a sandy sea-floor, owing to frequent or constant movement of the sediment, is not favourable to sedentary forms of life, and is therefore avoided by organisms which cannot shift

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for themselves. An ordinary siliceous sandstone might therefore be expected to be somewhat barren. Second, the permeable character of sandstones must favour the subsequent passage of percolating water which so frequently dissolves and removes organic bodies. Massive thick-bedded quartzose sandstones and red sarfdstones are, as a rule, singularly poor in organic remains. Certain thin-bedded argillaceous and thick-bedded calcareous sandstones, however, are not infrequently highly fossiliferous and this is especially the case when the sandstones occur in beds alternating and interosculating with dark carbonaceous or lighter coloured calcareous shales. Conglomerates are generally unfossiliferous, or, if fossils are present, these are usually more or less rolled and waterworn. For example, we may obtain, in some Carboniferous and Jurassic conglomerates, worn fragments of the trunks and branches of trees but the more delicate twigs and leaves are absent. So, again, in gravels and conglomerates of Pleistocene and Recent age, only the more resistant large bones and teeth of mammals are ever met with, and they are often rolled and broken. There are exceptions to every rule, however, for, now and again, tolerably well-preserved shells do occur in conglomerates. Volcanic Tuffs. In certain bedded volcanic tuffs fossils occur, but this is not common. Plant-remains have even been encountered in the coarse tuffs and agglomerates that occupy the throats or necks of certain ancient Carboniferous volcanoes in Scotland. Probably these represent trees, etc., which grew upon the slopes of the old cones after the volcanoes had become extinct. More rarely still, charred fragments of trees have been met with enclosed in the lower portion of an ancient lava. Schistose Rocks. It need hardly be said that these rocks are usually destitute of organic remains. Nevertheless, fossils are occasionally present in schists, as in certain metamorphic Silurian rocks in the neighbourhood of Christiania, and in the highly crystalline schists of Liassic age which enter into the structure of the Central Alps. Fossils differ much not only in regard to the state of preservation of their internal structure, but also of their external form. In many cases, they have been much compressed what were formerly cylindrical branches, for example, have often been flattened, so as to give lenticular sections when cut across. In limestones, marly shales, and calcareous sandstones, shells, corals, etc., usually retain their original shapes ; while in argillaceous shales, fossils of all kinds are apt to be more or less flattened a rule, however, which has many exceptions. In clay-slates and rocks which have rarely been preserved. As these tracks and castings occur chiefly in marine sedimentary rocks, it is very doubtful if

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any of them indicate earthworms. The " tubes " formed by many marine annelids are often met with as fossils. Molluscoidea. These are among the commonest and most abundant fossils. One great division (Polyzoa) comprises the lace-corals and seamats, which are chiefly marine, and, as fossils, often occur associated with other marine organisms. The other great division (Brachiopoda) is exclusively marine, and includes the lamp-shells, etc. one of the most important types of life with which the student of historical geology has to deal. Mollusca. The same holds true with the marine mollusca, which are more or less abundantly represented in every great system of strata. Not only are they of prime importance by reason of their abundance as regards genera, species, and individuals, but their shells, like those of the brachiopods, appear often in a comparatively perfect state of preservation. Freshwater shells and land-snails are of much less frequentoccurrence as fossils. Arthropoda. This phylum embraces lobsters, crabs, scorpions, spiders, centipedes, and insects, and is of great value to the geologist the crustaceans more especially, for a large proportion of these being marine, they are well represented by fossils. Some of the extinct types, as Trilobites, for example, are very characteristic fossils of the older systems in which they occur. Freshwater and terrestrial forms are not so commonly encountered, since they are largely confined to freshwater deposits and to lignite- and coal-bearing strata. Vertebrata- great phylum is most numerously represented by marine fishes. Marine types of reptiles and mammals also occur now and again, but with the exception of the fishes vertebrate remains of any kind are sparingly met with. Remains of birds and land-mammals are almost confined, as might have been expected, to freshwater and terrestrial accumulations. Importance of Fossils in Geological Investigations. It need hardly be said that the study of fossils to the biologist is of surpassing importance. Such study, indeed, cannot be ignored by him if he would understand the life-history of existing types. But it is not with that side of palaeontological inquiry that the practical or field-geologist is mainly concerned. He values fossils chiefly for the help they yield him in his endeavours to realise the conditions under which sedimentary rocks were formed, and to ascertain the chronological sequence of the strata. Climatic Conditions deduced from Fossils. Individual fossils, if of existing species, and occurring in situt may give valuable evidence as to former climatic conditions. Two examples may be cited. Certain relatively recent accumulations of calcareous

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-tufa, occurring at La Celle near Paris, have yielded numerous remains of the Canary laurel (Laurus Canariensis). There is no doubt, therefore, that this plant formerly flourished in Northern France. It is no longer a native of that country, however, its headquarters being in the Canary Islands, where it is found flourishing luxuriantly in the woody regions with a northern exposure, between a height of 1600 feet and 4800 feet above the sea regions which are nearly always enveloped in steaming vapours, and exposed to the heavy rains of winter. The temperature there keeps above 69 F. during the greater part of the year, rarely falling in the winter months below 59 or 60, and only on the coldest days reaching 49. The presence, therefore, of this variety of laurel in the Pleistocene tufa of La Celle shows that the winter climate of Northern France must formerly have been very mild. The laurel in question is most susceptible to cold, and as it flowers in the winter season, it is obvious that repeated frosts, such as are now experienced in the north of France, would prevent it reproducing its kind. Another and more familiar example of the important evidence which is sometimes afforded by fossil remains of existing types is that of the Polar willow (Salix polaris) a characteristic arctic plant, living in Northern Lapland, Spitzbergen, etc. This dwarf willow has been met with again and again in Pleistocene deposits in Southern Sweden, Denmark, England, etc., and in various parts of Central Europe, as far south as Bavaria and the low-lying parts of Switzerland. It cannot be doubted, therefore, that the appearance of the Polar willow so far south of its present habitat, points to a very considerable climatic change arctic conditions would seem to have prevailed at a relatively recent period in what are now the temperate latitudes of our continent. It is obvious, however, that the evidence of fossils as to climatic conditions must be much stronger when a whole assemblage of organic remains tells the same tale. In the case of the tufa of La Celle, for example, the Canary laurel is accompanied by the remains of many other plants, as well as by shells of land-snails, each of which is indicative of a milder and more equable climate than now characterizes Northern France. And the same is the case with the Polar willow, the evidence supplied by it being fortified by that of other high northern plants, and by the relics of such animals as lemming, arctic fox, etc. Great caution must be exercised in deducing climatic conditions from the occurrence of extinct forms of life. For these, even when they very closely resemble living types, need not have existed under similar conditions. For example, so long as the mammoth and woolly rhinoceros were only known from their skeletal remains, they were

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generally supposed to have existed under the same climatic conditions as their living representatives. We know now, however, that each was provided with a thick woolly and hairy covering, and was capable, therefore, of withstanding the rigours of a northern winter. In dealing with fossils consisting largely of extinct species, it is the general facies of a flora and fauna, and not individual forms, that are to be specially considered. For example, the London Clay (Eocene) has yielded a large number of types having a tropical or subtropical aspect. Amongst the plants are forms of sarsaparilla, aloe, amomum, fan-palms, fig, liquidambar, magnolia, eucalyptus, cinnamon, various proteaceous plants, etc. ; while the animals include turtles, tortoises, crocodiles, tapir-like pachyderms, and certain birds with affinities to living tropical types. Associated with these are many forms of molluscan life which have their nearest living representatives in warm latitudes, such as cones, cowries, volutes, nautilus, etc., together with sword-fish, saw-fish, sharks, and rays. All this is good evidence that a warm climate prevailed during the deposition of the London Clay. The land was clothed with tropical or subtropical vegetation, while corresponding types of animal-life haunted the rivers and flourished in the sea of the period. In the older geological systems we may say that all the species and nearly all the genera are extinct, so that any general resemblance which an assemblage of Palaeozoic fossils may have to those of some particular groups of living plants and animals may have no climatic significance whatsoever. We may feel sure, indeed, that the abundant flora of the Carboniferous period could not have flourished under arctic or even cold temperate conditions of climate; and we may be equally convinced that the abundant corals and cephalopods of Palaeozoic times, with their numerous congeners, were not denizens of cold seas. Existing conditions might even lead us to believe that the massive limestones of those early ages were most likely formed in genial waters. For at the present day it is in warm seas that lime-secreting organisms, such as corals, pelagic molluscs, and foraminifera, flourish most abundantly, and are there giving rise to widespread and thick accumulations of calcareous matter. But it would be rash to conclude that the climatic conditions of Palaeozoic times were similar to those of our present warm latitudes. When the geographical distribution of Palaeozoic floras and faunas, however, is kept in view, we may advance our inferences a step further. Should the fossils or groups of fossils of some particular formation be known to occur over vast areas of the earth's surface, in arctic, temperate, subtropical, and tropical latitudes, and even in similar latitudes of the Southern Hemisphere, we should be justified in the inference that the climatic

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conditions indicated by the fossils in question must have been singularly equable. The mere fact that in the earlier stages of the world's history, cosmopolitan forms of plant- and animal-life abounded, affords good ground for believing that the climatic conditions of those far past times differed considerably from the present. The climate of the globe in those days could not have been differentiated into such distinct zones as is now the case. Geographical Conditions deduced from Fossils. Fossils naturally yield evidence as to terrestrial, freshwater, and marine conditions. (a) Land-surfaces. These are seldom preserved. Nevertheless, they do occur in strata belonging to widely separated periods. Now and again, for example, the stools and roots of trees penetrating ancient soils, occur intqrbedded with sedimentary strata, a good example being furnished by the " dirt-bed " of Portland. This dirt-bed is simply an old soil containing the roots and stumps of extinct forms of cycads and conifers. It is intercalated between beds of freshwater origin, a succession which shows that, after the deposition of a wide area of fluviatile mud, dry land prevailed and eventually became covered with forests. Subsequently, owing probably to subsidence, the forest was submerged and buried under newer accumulations of fluviatile mud and silt. Many of the coal-seams of the Carboniferous period, with their underclays, tell a similar tale, and the same history is repeated by not a few of the lignites belonging to later geological periods, [Certain coals and lignites, however, appear to represent masses of vegetable matter which have probably been drifted from the land into estuaries and shallow bays of the sea.] The not infrequent occurrence of arachnids, insects, lizards, and land-snails associated with beds ofcoal and lignite, is additional evidence of terrestrial conditions. Amber, again, is an abundant product of the lignite-bearing beds of Germany, and unquestionably represents the gum and resin which exuded from some of the forest trees of Tertiary times. (b) Lacustrine conditions. These are indicated by the presence of numerous freshwater molluscs and small crustaceans which are sometimes so abundant as to form beds of marl and limestone. Plant-remains, insects, and other relics of land-life, such as reptiles or mammals, often occur in lacustrine deposits. It is from lacustrine and estuarine deposits, indeed, that we obtain our fullest information as to the life of former land-surfaces. (c) Marine conditions. Relatively deep or, at least, clear water is indicated by thick masses of limestone, more or less abundantly charged with corals, sea-lilies, and other marine organisms. This inference is based partly on the fact that these

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limestones are comparatively pure that is, they contain relatively little insoluble matter, and this is usually in a very finely divided state. In short, it is evident tftat such limestones have accumulated over parts of the sea-floor not reached by ordinary sediment conditions which, as a rule, can obtain only at a considerable distance from the shore, and often, therefore, in somewhat deep water. Further, we judge from the analogy of the present, that, as existing corals only flourish in clear water, their predecessors probably demanded similar conditions. This inference is strengthened by the fact that when, towards the top of a bed of limestone, the rock becomes more and more iippure, the corals, and certain of their congeners, often begin to diminish in size, and even to become somewhat distorted, as if the influx of muddy sediment had acted prejudicially upon their growth and development. Shallow-water conditions and proximity of the land are often evidenced by trails, burrows, and castings of annelids tracks of crustaceans, etc., footprints of reptiles, amphibians, birds, or mammals. Along with these the strata may yield more or less well-preserved plants, insect-remains, and other relics of land-life. Beds containing such fossils are not infrequently estuarine deposits, and often exhibit ripplemarks, rain-prints, and sun-cracks. Terrestrial Movements deduced from Fossils. The presence of marine fossils in a rock obviously indicates oscillations of the sea-level. The appearance, for example, in our maritime districts, at various heights above the present sea-level, of terraces of sand and gravel, crowded with seashells of still living species, is proof positive of some recent crustal movement either the land has risen or the sea-floor has subsided. Again, the existence at various depths on the sea-bottom of peat overlying the stools of trees belonging to kinds that still flourish in these islands, is evidence sufficient of a recent subsidence of the land. Geological Chronology and Fossils. In many cases it is quite impossible to correlate the formations occurring in separate regions by means of lithological characters alone. Within limited areas these may be reliable, but strata begin to change in character as they extend in various directions. Limestones, for example, may become gradually more and more argillaceous until at last they merge into shales, while these last may in their turn eventually pass into or interosculate with sandstones. Now, unless such changes could be followed in continuous open section, we could not possibly be sure that certain given beds of limestone, shale, and sandstone were exactly contemporaneous all laid down on one and the same sea-floor. These rocks are so dissimilar that,

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unless we actually traced the connecting passages, we could not tell how one was related to another. So far as lithological character is concerned, they might each have been formed at a different time. But if the separate sections of strata contained fossils having the same general fades and especially if several species were common to the limestones, the shales, and the sandstones, we could no longer doubt that all these rocks were accumulations formed in one and the same sea. Fossils are thus of paramount importance in the correlation of strata.

FAULTS Normal Faults. Dip-faults and Strike-faults their effect upon Outcrops. Oblique Faults. Systems of Faults. Step-faults. Trough- and Ridge-faults. Shifting of Faults. Reversed Faults. Transcurrent Faults. Origin of Faults. HAVING now learned that rocks of all kinds are more or less fissured, and that no small proportion of the joints by which they are thus traversed appear to owe their origin to crustal movements, we must next make the acquaintance of fissures of another kind, known as Faults or Dislocations. These are doubtless due likewise to crustal movements, but they differ from joints in being not mere cracks or rents, but fissures of displacement. The rocks on one side of a fault are thus abruptly truncated and brought against younger or older rocks on the other side. Three types of faults are recognised, namely, Normal Faults, Reversed Faults ,and Transverse Fault. NORMAL FAULTS - These dislocations are rarely, if ever, quite vertical, although in natural exposures they sometimes appear to be so. But when they are followed downwards, as in mining operations, they are invariably found to be inclined, the degree of inclination varying, it may be, from point to point, so that in places they occasionally show verticality. The general inclination of a fault from the vertical is termed the hade, and this, in the case of normal faults, is always in the direction of the downthrow. The degree of deviation from the vertical is quite indeterminate; but, as a general rule, the larger are more steeply inclined than the smaller faults. But to this rule many exceptions occur. The amount of vertical displacement is known as the throw of a fault, and is measured by protracting a line in a horizontal direction, across the fault from the truncated end of some particular bed until a perpendicular dropped from the protracted line can reach the other end of the selected stratum on the opposite side of the fault. Miners

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seldom use the term fault, but speak of downthrows or downcasts, and upthrows or upcasts according to the NORMAL FAULTS IN HORIZONTAL STRATA. direction in which they are working. Thus the faults would be described as downcasts or downthrows if they were encountered by a miner working in the direction representing faulted horizontal strata, the amount of throw is equal to the thickness. Strata cut across by an inclined fault are not only dropped to a lower level on the downthrow side, but the fault. The amount of displacement varies indefinitely. Some faults are mere slips of a few feet or inches; others are downthrows of several thousand yards. Between these extremes all gradations are met with.

Normal fault

Horst & Graban

Crystallography

Crystal Lattices

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One of the consequences of the way in which we have defined the unit cell is that the corners of the cell have absolutely identical chemical and structural environments (in addition, in some crystals there may be locations other than at the corners of the unit cell whihc also have identical environments). It is useful to imagine a hypothetical crystal made up of numerous unit cells stacked along the three repeat direction, x, y and z, with the corners of each individual cell marked by a point. These points together with any other points in the crystal with identical environments are termed lattice points:

This notion of a crystal lattice provides a wonderful insight into many of the fundamental properties of crystals such as shape and symmetry. For example a crystal growing free of interference will tend to develop faces parallel to lattice

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planes, and in particular lattice planes with the greatest density of lattice points. The concept of the crystal lattice therefore provides us with a very useful way of understanding the seemingly infinite variety of crystal shapes found in nature. In order to make full use of the lattice concept we will need to (stretch the neurons in order to) develop some technique enabling the systematic description of the orientation of lattice planes and rows. Miller indices of lattice planes: A set of lattice planes intercept each of the crystallographic axes a finite number of times per unit length of axis. By convention, the number of intercepts of planes per unit length along the x-axis is termed h, per unit length along the y-axis is termed k and per unit length along the z-axis is termed l. The unit length along each axis is taken to be the length of the unit cell parallel to that axis, that is a for the x-axis, b for the y-axis and c for the z-axis. The Miller Indices of the plane are then (hkl). Thus the set of lattice planes (hkl) divides a into h parts, b into k parts and c into l parts. An alternative way of conceptualising Miller indices is to consider the first plane out from the origin in any set of planes. This plane will make the intercepts a/h, b/k, c/l (where h, k and l are integers), on the x-axis, y-axis and z-axis respectively.

Symmetry in crystals Each of the seven crystal systems is characterised by different symmetries. By

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inspection of the above table we can see that there must be a general decrease in the symmetry from the cubic system down through to the triclinic system. Consequently, the triclinic system is said to have lower symmetry than the cubic system and, conversely, the cubic system higher symmetry than the triclinic system.. In order to understand the basic symmetry elements in crystallography we will initially consider examples of the symmetry elements in a cube (a = b = c,

re: Rotation axes of symmetry: Rotation axes of symmetry are axes which upon rotation reproduce the exact configuration of the crystal. An n-fold rotation axis of symmetry repeats the structure n times in one complete 360° rotation. In crystals rotation axes can be sixfold (termed hexad), fourfold (tetrad), threefold (triad), twofold (diad), or onefold (monad). The monad is a trivial since it merely states that upon rotation through 360° the crystal returns to the initial position. A cube contains a number of different tetrads, triads and diads (try and determine the total number in each case):

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Inversion axes of symmetry: A symmetry type involving a rotation about a line plus an inversion through a point (on the line) is known as an inversion axis of symmetry. In any crystal the operation of an inversion axes of symmetry can always be achieved by a combination of the other symmetry operators in that crystal. As already stated each of the seven crystal systems have characteristic symmetries. Recognition of these symmetry elements allows us to classify any crystal into the appropriate class. For example the presence of a tetrad indicates either cubic or tetragonal; more than one tetrad and it must be cubic (however, the presence of a tetrad is not the characteristic symmetry element of the cubic system which is, rather, the resence of three triads). Hexads are diagnostic of the hexagonal system, while traids preclude triclinic and monoclinic systems. The triclinic system has no rotational axes of symmetry, except of course the trivial monad. The characteristic symmetry elements in each of the seven groups are listed below: Cubic Three triads Hexagonal One hexad (// z) Tetragonal One tetrad (// z) Trigonal One triad (// [111]) Orthorhombic Three perpendicular diads (// x, y and z) Monoclinic One diad (// y) FOLD

A geological fold occurs when one or a stack of originally flat and planar surfaces, such as sedimentary strata, are bent or curved as a result of permanent deformation. Synsedimentary folds are those due to slumping of sedimentary material before it is lithified. Folds in rocks vary in size from

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microscopic crinkles to mountain-sized folds. They occur singly as isolated folds and in extensive fold trains of different sizes, on a variety of scales.

Folds form under varied conditions of stress, hydrostatic pressure, pore pressure, and temperature gradient, as evidenced by their presence in soft sediments, the full spectrum of metamorphic rocks, and even as primary flow structures in some igneous rocks. A set of folds distributed on a regional scale constitutes a fold belt, a common feature of orogenic zones. Folds are commonly formed by shortening of existing layers, but may also be formed as a result of displacement on a non-planar fault (fault bend fold), at the tip of a propagating fault ,by differential compaction or due to the effects of a high-level igneous intrusion e.g. above a laccolith.

Fold terminology in two dimensions

A fold surface seen in profile can be divided into hinge and limb portions. The limbs are the flanks of the fold and the hinge is where the flanks join together. The hinge point is the point of minimum radius of curvature (maximum curvature) for a fold. The crest of the fold is the highest point of the fold surface, and the trough is the lowest point. The inflection point of a fold is the point on a limb at which the concavity reverses; on regular folds, this is the midpoint of the limb.

Fold terminology in three dimensions

The hinge points along an entire folded surface form a hinge line, which can be either a crest line or a trough line. The trend and plunge of a linear hinge line gives you information about the orientation of the fold. To more completely describe the orientation of a fold, one must describe the axial surface. The axial surface is the surface defined by connecting all the hinge lines of stacked folding surfaces. If the axial surface is a planar surface then it is called the axial plane and can be described by the strike and dip of the plane. An axial trace is the line of intersection of the axial surface with any other surface.

Finally, folds can have, but don't necessarily have a fold axis. A fold axis, “is the closest approximation to a straight line that when moved parallel to itself, generates the form of the fold.”. A fold that can be generated by a fold axis is called a cylindrical fold. This term has been broadened to include near-cylindrical folds. Often, the fold axis is the same as the hinge line.

TRILOBITA

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Trilobites (meaning "three lobes") are a fossil group of extinct marine arthropods that form the class Trilobita. Trilobites form one of the earliest known groups of arthropods. The first appearance of trilobites in the fossil record defines the base of the Atdabanian stage of the Early Cambrian period (521 million years ago), and they flourished throughout the lower Paleozoicera before beginning a drawn-out decline to extinction when, during the Devonian, all trilobite orders except the Proetids died out. Trilobites finally disappeared in the mass extinction at the end of the Permian about 250 million years ago. The trilobites were among the most successful of all early animals, roaming the oceans for over 270 million years.

By the time trilobites first appeared in the fossil record, they were already highly diversified and geographically dispersed. Because trilobites had wide diversity and an easily fossilized exoskeleton, an extensive fossil record was left behind, with some 17,000 known species spanning Paleozoic time. The study of these fossils has facilitated important contributions to biostratigraphy, paleontology, evolutionary biology, and plate tectonics. Trilobites are often placed within the arthropod subphylum Schizoramia within the superclass Arachnomorpha (equivalent to the Arachnata), although several alternative taxonomies are found in the literature.

Trilobites had many lifestyles; some moved over the sea bed as predators, scavengers, or filter feeders, and some swam, feeding on plankton. Most lifestyles expected of modern marine arthropods are seen in trilobites, with the possible exception of parasitism (scientific debate continues). Some trilobites (particularly the family Olenidae) are even thought to have evolved a symbiotic relationship with sulfur-eating bacteria from which they derived food.

Sea floor spreading

Seafloor spreading is a process that occurs at mid-ocean ridges, where new oceanic crust is formed through volcanic activity and then gradually moves away from the ridge. Seafloor spreading helps explain continental drift in the theory of plate tectonics. When oceanic plates diverge, tensional stress causes fractures to occur in the lithosphere. Basaltic magma rises up the fractures and

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cools on the ocean floor to form new sea floor. Older rocks will be found farther away from the spreading zone while younger rocks will be found nearer to the spreading zone.

Earlier of continental drift were that continents"ploughed" through the sea. The idea that the seafloor itself moves (and carries the continents with it) as it expands from a central axis was proposed by Harry Hess from Princeton University in the 1960s.The theory is well accepted now, and the phenomenon is known to be caused by convection currents in the asthenosphere, which is the plastic, relatively weak part of the upper mantle]

In the general case, sea floor spreading starts as a rift in a continental land mass, similar to the Red Sea-East Africa Rift System today. The process starts with heating at the base of the continental crust which causes it to become more plastic and less dense. Because less dense objects rise in relation to denser objects, the area being heated becomes a broad dome (see isostasy). As the crust bows upward, fractures occur that gradually grow into rifts. The typical rift system consists of three rift arms at approximately 120 degree angles. These areas are named triple junctions and can be found in several places across the world today. The separated margins of the continentsevolve to form passive margins. Hess' theory was that new seafloor is formed when magma is forced upward toward the surface at a mid-ocean ridge.

If spreading continues past the incipient stage described above, two of the rift arms will open while the third arm stops opening and becomes a 'failed rift'. As the two active rifts continue to open, eventually the continental crust is attenuated as far as it will stretch. At this point basaltic oceanic crust begins to form between the separating continental fragments. When one of the rifts opens into the existing ocean, the rift system is flooded with seawater and becomes a new sea. The Red Sea is an example of a new arm of the sea. The East African rift was thought to be a "failed" arm that was opening somewhat more slowly than the other two arms, but in 2005 the Ethiopian Afar Geophysical Lithospheric Experiment reported that in the Afar region last September, a 60 km fissure opened as wide as eight meters. During this period of initial flooding the new sea is sensitive to changes in climate and eustasy. As a result, the new sea will evaporate (partially or completely) several times before the elevation of the rift valley has been lowered to the point that the sea becomes stable. During this period of evaporation large evaporite deposits will be made in the rift valley. Later

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these deposits have the potential to become hydrocarbon seals and are of particular interest to petroleum geologists.

Sea floor spreading can stop during the process, but if it continues to the point that the continent is completely severed, then a new ocean basin is created. The Red Sea has not yet completely split Arabia from Africa, but a similar feature can be found on the other side of Africa that has broken completely free. South America once fit into the area of the Niger Delta. The Niger River has formed in the failed rift arm of the triple junction.

Continued spreading and subduction

Spreading at a mid-ocean ridge

As new seafloor forms and spreads apart from the mid-ocean ridge it slowly cools over time. Older seafloor is therefore colder than new seafloor, and older oceanic basins deeper than new oceanic basins due to isostasy. If the diameter of the earth remains relatively constant despite the production of new crust, a mechanism must exist by which crust is also destroyed. The destruction of oceanic crust occurs at subduction zones where oceanic crust is forced under either continental crust or oceanic crust. Today, the Atlantic basin is actively spreading at the Mid-Atlantic Ridge. Only a small portion of the oceanic crust produced in the Atlantic is subducted. However, the plates making up the Pacific Ocean are experiencing subduction along many of their boundaries which causes the volcanic activity in what has been termed the Ring of Fire of the Pacific Ocean. The Pacific is also home to one of the world's most active spreading centres (the East Pacific Rise) with spreading rates of up to 13 cm/yr. The Mid-Atlantic Ridge is a slow-spreading centre, while the East Pacific Rise is used as an example of fast spreading. The differences in spreading rates affect not only the

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geometries of the ridges but also the geochemistry of the basalts that are produced.

Since the new oceanic basins are shallower than the old oceanic basins, the total capacity of the world's ocean basins decreases during times of active sea floor spreading. During the opening of the Atlantic Ocean, sea level was so high that a Western Interior Seaway formed across North America from the Gulf of Mexico to the Arctic Ocean.

Texture of Sedimentary rocks

The word 'texture' refers to the size, shape, packing and fabric of the components of the rock. Since the sedimentary rocks are broadly classified as (1) exogenetic or clastic rocks and (2) endogenetic rocks or the chemically precipitated amorphous or crystalline rocks, accordingly their texture is also classified into two broad categories.

1. Clastic texture:

It includes elements like:

(I) Size, (II) Shape, (III) Sphericity, (IV) Packing, (V) Fabric.

(a) Size. The grain size is dependent on the (i) mode of weathering, (ii) nature of the source rock, and (iii) kind and distance of transport and the nature of deposition.

Broadly, the size characters of the sediments are described as either coarse, medium or fine. The size grades of the clastic particles, in the went worth scale are indicated as follows:

The sire analysis results are represented in form of tables. Histograms or frequency curves as well as by statistical methods." In statistical methods the following measurements are made:

(a) Measurement of central tendency:

Which determines the average size of the distribution and refers to the overall competency of the transporting medium.

(b) Measurement of dispersion:

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It determines the turbulence of the transporting medium or the amount of reworking of the sediment has undergone prior to final burial.

(c) Measurement of skewness:

It determines whether the coarser and finer admixture occur in same proportion, in a, sediment.

(d) Measurement of kurtosis:

It determines the peakedness or flat-toppedness of a distribution.

Geological Significance:

The size analysis indicates the following:

1. Provenance:

(i) Composition of the source rock is an important factor that determines the extent to which the component minerals are susceptible to weathering and liable to pass on to the products and reduced in size and shape.

(ii) Besides the coarse or fine texture is also a function of the source area.

2. Transport:

As we know, more the distance of transport, finer is the grain size. Besides, the character of the sediments are also governed by the mode of transport, i.e., traction, saltation, suspension, which is a function of the kinetic energy of the transporting medium. Higher the energy, the coarser particles can be transported.

3. Depositional environment. 4. Palaeo current:

The coarser sediments carried in rolling are deposited in basin margin, whereas the finer sediments are gradually carried to the centre. As such there is a regular variation in grain size from the margin to the centre of the basin.

5. Transporting medium:

Graded sediments are the result of long continued transport, while ill-sorted sediments of a rapid and confused deposition like glacial deposits. Aeolian deposits are apt to be well graded and uniform.

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6. Tectonics of depositional site:

With an increase in the rate subsidence of the area of deposition, the grainsize decreases and the average sorting is poor. Under stable condition, the grainsize is determined by the texture of the material available for reworking sorting steadily improves.

II. Shape:

It is defined as the sharpness of corners and edge of a clastic fragment. Accordingly the shape may be angular, sub angular, sub-rounded, rounded, well rounded, etc. The shape of the sedimentary grain is determined by:

(i) Original shape of the mineral, (ii) Stability of minerals and (iii) amount and nature of transport.

III. Sphericity:

It is defined as the extent to which a particle approaches a sphere. It depends on (i) distance of transport. (ii) Mode of transport, and (iii) provenance.

(i) Longer the distance of transportation, more chances of being reworked and therefore more the degree of roundness. Besides, wind produces perfect rounding, glacier does the least.

(ii) The mode of transport like traction, saltation and suspension produces particles of variable roundness.

(iii) Generally the elongation quotient is maximum in metamorphic rocks, less in igneous rocks and very less in sedimentary rocks. Thus when the source rock is of metamorphic origin, the sphericity of the clastic grains is not that much pronounced.

IV. Packing:

It is the manner of aggregation of sedimentary grains, which are held together in place in the earth's gravitational field. There are six methods of packing out of which the rhombohedfal packing is the most compact and tight whereas, the cubic packing is the loosest possible packing. It determines the porosity and permeability of sedimentary rocks.

V. Fabric:

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It is the arrangement of the clastic particles in sediments. It is defined as the orientation of the grains or lack of it with which the sedimentary rock is composed. Pebbles, sand grains, mica-flakes etc. are the most useful fabric elements, also, some fossils like gastroped shells etc. It determines the palaeo-current direction.

2. Non-clastic textures:

It is formed as a result of deposition through chemical reaction. They are transported chemically by getting dissolved in the transported media but reappear due to precipitation or evaporation. It is of two types:

(a) Crystalline texture.

(b) Non-crystalline tenure.

(a) Crystalline texture:

They are formed due to direct precipitation from a saturated solution, and the result is an interlocks aggregate of crystals.

(b) Non-crystalline texture:

When colloids coagulate they form a gelatin like mass. This gelatinous mass may lose some of, he water in it and eventually harden to form an amorphous mass Nodular, oolitic, spherulitic textures are the examples. Many concretionary, botryoidal, reniform, nodular, oolitic and pisolitic textures are believed to be of colloidal origin and they show noncrystalline textures as described above.

Classification of Limestone

INTRODUCTION

Something like about one-fifth of all sedimentary rocks are carbonate rocks. The two main kinds of carbonate rocks, limestones and dolostones, together with sandstones and shales, are what might be called the “big four” of sedimentary rock types. I’m reluctant to try to guess what percentage of all sedimentary rocks those “big four” account for, but the figure must be in the upper nineties.

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Moreover, carbonate rocks are economically important because together with sandstones they constitute reservoirs for most of the world’s petroleum and gas reserves (and let’s not forget that they are the source of all of the world’s portland cement—not a jazzy, exciting resource, but a very important one for our modern civilization). 1.2 Up until about fifty years ago, the petrologic study of carbonate rocks lagged far behind that of siliciclastic rocks. Since that time, however, there has been great progress, as it has become realized that to a great extent carbonate rocks can be treated as clastic deposits analogous to sandstones and shales (the main exception being reef limestones). Great progress has also been made in recent years on the geochemistry of carbonate precipitation, the role of organisms in carbonate deposition, and the diagenesis of carbonate sediments. 1.3 Carbonate sediments are often described as chemically precipitated. In one sense, that’s true: they are formed by precipitation of one or another carbonate mineral in various sedimentary environments. But don’t let the term “chemically precipitated” fool you: they don’t form in the same way that rock candy does from a sugar solution on your windowsill. Some carbonate sediments are indeed precipitated directly from seawater, in the form of fine crystals in the water column, which then settle to the seafloor, or as successive spherical shells deposited around a nucleus particle in a warm, shallow marine environment. Most, however, are biochemically precipitated, in the tissues of organisms, mainly marine invertebrates of various phyla. 1.4 The title of this chapter could have been “carbonate rocks”, but it seems somewhat more natural to restrict it to limestones. You will learn that dolomite is almost invariably a secondary rather than a primary sedimentary mineral, meaning that carbonate sediments don’t start their lives as dolomite. Much dolomite is precipitated very early, however, at shallow depths in the originally calcium carbonate sediment, although much is also precipitated late, after deep burial has produced solid limestone. Material on dolostone (a carbonate rock consisting mainly of the mineral dolomite) is postponed until the later chapter on diagenesis. 2. CARBONATE MINERALS 2.

About sixty minerals have the carbonate ion in their composition. But there are only three really important carbonate minerals: calcite, aragonite, and dolomite. (In the parlance of mineralogy, the first two are said to be polymorphs.) And aragonite is unimportant in ancient rocks, because it reverts to calcite with time. Other sedimentary carbonates of non-negligible importance are magnesite (magnesium carbonate) and siderite (ferrous iron carbonate). Dolomites containing some percentage of Fe2+ are called ferroan dolomite. The middle

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member of the range between dolomite and hypothetical (Fe, Ca)(CO3)2 is called ankerite. There are also a few significant carbonate evaporite minerals (trona, natron) we won't consider here. 2.2 Figure 5-1 is a composition triangle showing the range of carbonate minerals stable at the low temperatures at or near the Earth’s surface. It’s in terms of the three divalent positive ions, Ca2+ (0.99 Å), Mg2+ (0.66 Å), and Fe2+ (0.74 Å), that are abundant and of about the right size to fit into carbonate structures. (Remember that one angstrom is equal to 10-10 meters, or 0.1 nanometers.) 2.3 The two most important minerals are calcite, CaCO3, and dolomite, Ca(Mg,Fe)(CO3)2. Crystals of calcite and dolomite have rhombohedral symmetry. Think in terms of the simple crystal structure of halite, NaCl, in which effectively spherical Na+ and Cl- ions alternate along each of three mutually perpendicular directions to form a cubic structure with each Na+ placed between six Cl- ions and vice versa. The structure of calcite is similar, with Ca2+ and CO3 2- ions alternating in three directions. Ca2+ is about as big as Na+ , but CO3 2- is larger and has an effectively triangular shape and so takes up more space than the Clions. The CO3 2- ions have to be cocked at an angle to fit between the six nearest Ca2+ ions. So the whole array can be thought of as squeezed along one of the inner diagonals of the cube until the three lines of alternating ions meet at about 106° instead of at 90°. This results in rhombohedral rather than cubic symmetry. 2.4 In dolomite, about half of the Ca2+ ions are replaced by Mg2+ (and Fe2+) ions. If these ions were of the same size, there could be unlimited substitution of one for the other, and there would be complete isomorphism between calcite and magnesite. But the effective radius of Ca2+ is 36% larger than that of Mg2+, so the presence of more than a few percent Mg2+ in the calcite lattice would cause so great a distortion that the structure would be unstable. But it turns out that Ca2+ and Mg2+ can be present in about equal numbers, alternating regularly between the CO3 2- ions along each of the three directions, thus forming separate sheets of Ca2+ and Mg2+ in the structure; that’s the mineral dolomite. This results in slightly different angles and different symmetry. 113 2.5 Most of the calcite precipitated by marine organisms contains a certain percentage of magnesium. Such calcite is called magnesian calcite; it’s subdivided into low-magnesium calcite and high-magnesium calcite at 4% MgCO3 content. Generally the more advanced the organism, the less magnesium in the calcite. In the case of red algae, an important sediment producer, the percentage is as much as 25%. Magnesian calcite is unstable, and it eventually expels its magnesium, but that takes a long time. 2.6 The effective ionic radii of Fe2+ and Mg2+ are almost the same, so these two ions substitute for

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each other in any proportion. There is complete isomorphism between magnesite and siderite, and also between dolomite and the iron equivalent, called ferroan dolomite. So you should expect most dolomite to contain at least some iron. This is why dolostones are typically tan-weathering, in contrast to usually gray-weathering limestone. 2.7 The other calcium carbonate mineral, aragonite, has an entirely different structure with orthorhombic symmetry. Aragonite can’t tolerate even a few percent Mg2+ or Fe2+, but it can take some Sr2+ and Ba2+, which have much larger effective sizes. Also, some SO4 2- sulfate ions can replace the carbonate ions. 2.8 Aragonite is the high-pressure form of calcium carbonate, and calcite is the theoretically stable form under all sedimentary conditions. But in most sedimentary environments aragonite is precipitated rather than calcite. The exception is purely fresh-water inorganic precipitation, which is of minor sedimentological importance. Supersaturation, warm water, and the presence of sulfate and magnesium ions tend to promote precipitation of aragonite instead of calcite. Aragonite eventually reverts to calcite, but it sometimes takes a long time; where sealed in sulfate-rich impermeable rocks, aragonite has in some cases been found in rocks even as old as the late Paleozoic. 3. THE CARBONATE BUDGET OF THE OCEANS 3.1

Calcite is more stable in pure water than its polymorph aragonite, so if the carbonate system in the oceans obeyed thermodynamics, calcite rather than aragonite should be precipitated. But it’s observed that aragonite is precipitated instead. In the case of inorganic precipitation of aragonite, probably this has to do with the presence of other kinds of ions, like sulfate and magnesium, in sea water. These ions act as kinetic inhibitors and serve to impede the growth of calcite. In the case of biogenic precipitation, organisms also seem to disregard the laws of thermodynamics, and deposit either calcite or aragonite, or both at the same time, or one at one time and one at another. Some details follow. calcareous algae: Halimeda and other green algae, aragonite; Lithothamnium and other reef-making red algae, calcite modern corals (hexacorals): aragonite 114 Paleozoic corals (rugose corals, tabulate corals): probably calcite, by good preservation of structures brachiopods, bryozoans, foraminifera: calcite then and now echinoderms: calcite then and now; large single-crystal skeletal components, very durable trilobites: probably calcite, by good preservation of structures mollusks: mostly aragonite, sometimes partly or entirely calcite 3.2 At present, most of the open surface waters of the oceans, except at high latitudes, are about saturated or even supersaturated with respect to CaCO3, so there must be at least a broad inorganic control on precipitation of CaCO3. Figure 5-2 shows the degree of

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saturation of aragonite and calcite with depth for both the Atlantic Ocean and the Pacific Ocean. It's not known how long things have been this way, but judging from the overall similarity of present deposits and past deposits, the situation has probably been the same at least back into the Precambrian. In the Precambrian, the saturation of seawater with calcium carbonate may have been the highest of all time, because of much higher levels of CO2 in the atmosphere. 3.3 By river contributions of dissolved load from rock weathering on the continents, the doubling time of Ca2+ in the oceans should be about a million years. (By the doubling time is meant the time it would take for the concentration of a given substance to increase by a factor of two if no other process is at work to remove the substance at the same time.) Obviously something must be taking it out at about the same rate. The same is true for Mg2+ ion, but the doubling time is 115 Figure 5-2: Degree of saturation with respect to calcite and aragonite with depth in the Atlantic arid Pacific Oceans ATLANTIC OCEAN PACIFIC OCEAN ATLANTIC OCEAN PACIFIC OCEAN DEPTH Km DEPTH Km 6 5 4 3 2 1 0 .5 .6 .8 1.0 1.5 2.0 3.0 4.0 5.0 7.0 6 5 4 3 2 1 0 .3 .4 .5 .6 .8 1.0 1.5 2.0 3.0 4.0 5.0 DEGREE OF SATURATION, CALCITE DEGREE OF SATURATION, ARAGONITE Figure by MIT OCW. about an order of magnitude greater, something like twenty million years. In the case of calcium ion, the obvious process by which it’s removed from solution is precipitation of carbonate minerals. For magnesium ion, however, the removal process is less obvious—because, as you will see presently, dolomite is being precipitated as a primary sedimentary mineral almost not at all in the modern oceans. It’s known that Mg is removed from seawater very efficiently by reaction of the seawater with freshly crystallized basalt at and near mid-ocean ridges. Also, in shallow-water carbonate-producing environments the magnesium in solution is buried with the calcium carbonate sediments, where, with time, during shallow-burial diagenesis, it dolomitizes the calcium carbonate minerals—about which you will learn more in the later chapter on diagenesis. 3.4 As for the entire ocean bottom, about one gram of CaCO3 is deposited per square centimeter of ocean bottom per 1000 years on the average, on the basis of balance considerations. Over large areas, however, no carbonate is being deposited at all, whereas in other places the rates of accumulation are far greater than the average. 3.5 The rate of deposition of CaCO3 is much greater in shallow ocean than in the deep ocean, but the volume of newly deposited shallow-water CaCO3 is far smaller than that of deep- sea CaCO3. But this is true only for the present time: planktonic carbonate-secreting organisms evolved late in geologic history, in the Mesozoic, so the volume of shallow-water CaCO3 must have been much greater

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in the Precambrian and the Paleozoic than now (on the reasonable assumption that the rate of supply of dissolved calcium ion was about the same). 3.6 Another point worth noting here is that, from the standpoint of long-term geologic history, deep-sea storage of carbonate is only temporary: sea-floor spreading and subduction of oceanic lithosphere at continental margins provides a very satisfying way of reincorporating deep-sea carbonate deposits into the geological record of the continents, albeit usually in unrecognizable form. Before the plate-tectonics revolution in the 1960s, geologists believed that what was put into the deep ocean stayed there forever! 4. GEOCHEMISTRY OF CARBONATE PRECIPITATION 4.1 Precipitation of carbonate in natural waters is more complicated than that of, say, halides or sulfates, because of the dissolution of carbon dioxide in natural waters. Here are the reactions that are relevant to carbonate precipitation: CO2 (gas) + H2O ⇔ CO2 (aqueous solution) + H2O CO2 + H2O ⇔ H2CO3 (carbonic acid, about 1%) H2CO3 ⇔ H+ + HCO3 - K = 4.3 x 10-7 116 HCO3 - ⇔ H+ + CO3 2- K = 4.8 x 10-11 CaCO3 (calcite) ⇔ Ca2+ + CO3 2- K = 3.8 x 10-9 CaCO3 (aragonite) ⇔ Ca2+ + CO3 2- K = 6.0 x 10-9 BACKGROUND: EQUILIBRIUM CONSTANTS You probably can recall from some earlier chemistry course that for a chemical reaction aA + bB ⇔ cC + dD where, if the reaction goes to the right, A and B are called the reactants and C and D are called the products (whereas if the reaction goes to the left, A and B are the products and C and D are the reactants), and the coefficients a, b, c, and d are the numbers of atoms or ions involved in the reaction, the equilibrium constant K is ([A]a [B]b )/([C]c [D]d ) where the square brackets signify the concentrations (technically, the activities, but in dilute solutions the two are not greatly different) of the various substances. If the equilibrium constant for a given reaction is small, that means that at equilibrium (that is, the reaction is proceeding just as fast to the left as to the right, so the concentrations of the various substances remain the same through time) the concentrations of the substances on the right side of the reaction are smaller than the concentrations of the substances on the left side of the reaction. 4.2 There are various ways of adding these reactions, but the way that’s most relevant to carbonate precipitation in the oceans is CaCO3 (solid) + H2O + CO2 ⇔ Ca2+ + 2CO3 2- + 2H+ So anything that increases the concentration of dissolved CO2 tends to cause dissolution of calcium carbonate, 117 and anything that decreases the concentration of dissolved CO2 tends to cause precipitation of calcium carbonate. The two most important effects are: temperature: as the water temperature increases, the equilibrium solubility of CO2 decreases, so as sea water is warmed there is a tendency for CO2 to be released back into the

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atmosphere and for CaCO3 to be precipitated. photosynthesis: in photosynthesis, plants take up CO2 from the environment and fix it in organic compounds in their tissues, thereby releasing oxygen. 4.3 Another important factor is this: carbonate-secreting marine invertebrates live in greatest numbers in the warm shallow parts of the oceans. These organisms can secrete CaCO3 even from water that is not saturated in CaCO3, but they do it best and most abundantly where the water is saturated. 4.4 So it makes sense that most of the CaCO3 precipitated in the oceans today is in warm, shallow water, where the water is warmed, so that the concentration of dissolved CO2 is lowered and saturation with respect to CaCO3 is thus enhanced, and where aquatic plants (largely algae) flourish. 4.5 In summary, there is a broad inorganic control on carbonate precipitation in the oceans, but the specific controls have to do with local water temperature and photosynthesis. 4.6 So far in this section we have addressed only the precipitation or dissolution of the calcium carbonate minerals. How about the mineral dolomite? To deal with precipitation of dolomite, we have to think about undersaturation and oversaturation (also called supersaturation). Suppose you put a piece of calcite in a beaker of distilled water, and for the sake of simplicity you arrange that no carbon dioxide is dissolved in the water. You know what will happen: the calcite dissolves slowly, and, as it does, the concentrations of dissolved Ca2+ ions and CO3 2- ions increase The solution is said to be undersaturated with respect to calcite. Eventually (and it would take months), the reaction reaches equilibrium, whereupon the concentrations of the Ca2+ ions and the CO3 2- ions reach constant values. If, however, we somehow pump Ca2+ ions into the solution, we drive the reaction toward precipitation of calcite. The solution is said to be oversaturated. 4.7 We can think separately about the saturation state of calcite (or aragonite) in seawater and about the saturation state of dolomite in seawater. The conventional way of doing that is to define a quantity variously called the solution quotient or the mass action quotient or the ion activity product (the last designated IAP), which has exactly the same form as the equilibrium constant but the concentrations are those that exist at a given time in the solution, whether the 118 solution is in equilibrium (the state of saturation) or undersaturated or oversaturated. If the solution is oversaturated, then the IAP is greater than the equilibrium constant K; if the solution is undersaturated, then the IAP is less than the equilibrium constant K. 4.8 It turns out that in warm, shallow seawater (the part of the oceans that is most relevant to precipitation of carbonate minerals, as you saw above), all three of the important carbonate minerals—calcite, aragonite, and dolomite—are in a state of oversaturation: that is, the situation should be

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conducive to precipitation. You have seen that calcite or aragonite is indeed precipitated, but dolomite is not. Yet, perhaps surprisingly, the ratio IAP/K for the three minerals is as follows: Aragonite: 2.1 Calcite: 3.3 Dolomite: 55 119 4.9 Seawater is far more oversaturated with respect to dolomite than it is with respect to aragonite and calcite! How, then, can we account for this seemingly paradoxical situation? We need to think not just about the equilibrium of chemical reactions but also about the kinetics of chemical reactions. Ionic crystals are built ion by ion, at some rate. Building of crystals of minerals with a greater degree of ionic ordering is inherently slower—more difficult— than that of crystals with a lesser degree of ordering. Dolomite has a much more highly ordered structure than does calcite or aragonite, inasmuch as in dolomite, in contrast to the two calcium carbonate minerals, the calciums and the magnesiums need to be in a regularly alternating array. Reaction kinetics is the basic reason why dolomite, despite its great oversaturation, is not precipitated from seawater: it loses out in the kinetic competition. And this must have been true in the geologic past as well as at present. 4.10 Another way of looking at dissolution of calcite is to write a reaction that pertains to rapid dissolution in the presence of hydrogen ions: CaCO3 (solid) + 2H+ ⇒ Ca2+ + CO2 + H2O What’s going on here is that the calcite dissolves in acidic waters to release calcium ions and carbon dioxide gas. That’s what happens when we add groundup limestone to our lawns and gardens to decrease the acidity of the soil. It’s also what happens when we use our little bottle of dilute hydrochloric acid in the field to test whether a given rock is a limestone. When I was an undergrad, I worked in my advisor’s paleontology laboratory, where we dissolved large blocks of limestone containing silicified fossils by painting the bottom surface of the block with a plastic resin and then putting the block in a tub of dilute hydrochloric acid in a fume hood. After the entire block was dissolved, we rinsed the residue and picked the silicified fossils with tweezers under a microscope. The topic of silicification will appear is a later chapter. (As a final note, the above reaction can also be used to account for the gradual dissolution of marble monuments and gravestones by acidic rainwater.)

5. MODERN MARINE CARBONATE SEDIMENTS 5.1 Carbonate sediments are forming today on many parts of the ocean bottom, according to definite but complex controls. Presumably this has been so throughout much of geologic history, but because organisms do most of (or, at least, are ultimately responsible for) the deposition, the controls must have changed through time as the organisms evolved. 120 5.2 Carbonate deposits are most abundant between

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about 30° N and 30° S. (They are not restricted to that zone, though: the topic of cold-water carbonates is an active area of research nowadays.) The low-latitude regions are where the surface waters tend to be saturated with respect to CaCO3, and where the water is warm enough the year round for carbonate-secreting organisms to flourish. But this picture is obviously complicated in detail, by patterns of ocean currents, nutrient supply, and dilution by siliciclastic sediments. 5.3 There are three main classes of modern marine carbonate sediments: • calcareous oozes in the deep ocean • carbonate buildups • calcareous sand and mud on platforms (shelf and ramp carbonates)

6. CALCAREOUS OOZE 6.1 More than a third of the present deep ocean bottom is covered with sediment containing more than 30% CaCO3. The carbonate in these sediments is in the form of tiny shells or tests of various carbonate-secreting planktonic organisms that live in the warm shallow waters above. Such deposits are called ooze; there is calcareous ooze and siliceous ooze, depending upon what the organisms secrete. 6.2 The most abundant kind of calcareous ooze is foraminiferal ooze. Foraminifera are single-celled protozans. In the oceans of today there are about thirty species of foraminifera (forams, for short) belonging to two families, Globigerinidae and Globorotaliidae. These are the only two of more than fifty foram families that are adapted to a planktonic rather than a benthic mode of life. But these planktonic forams grow in enormous numbers in the warm shallow waters of the ocean. At subdividing time the protoplasm of each foram subdivides into zoospores that swarm out to develop into new forams, leaving the empty test to sink to the bottom. So very little organic matter goes down with the tests. Brasier, M.D., 1980, Microfossils: George Allen & Unwin, 193 p Brasier, M.D., 1980, Microfossils: George Allen & Unwin, 193 p. Boersma, A., 1978, Foraminifera, in Haq, B.U., and Boersma,., Introduction to Marine Micropaleontology: Elsevier, 376 p. shows what some planktonic forams look like. 6.3 Because of the greater solubility of CaCO3 in the colder deeper waters, which are derived from cold polar regions and thus contain more dissolved CO2, below a certain depth the foram tests are completely dissolved before they have a chance to be covered by more tests. is a graph of CaCO3 content of bottom sediments vs. depth. 121 6.4 Calcareous oozes are abundant down to about 4000 m, but below that they become much less abundant; there is very little carbonate below about 4500 m. Oceanographers call the depth between 4000 and 4500 m where carbonate becomes less abundant the carbonate compensation depth, or CCD; it’s actually a narrow zone rather than a single depth. 6.5 The oceans are undersaturated with respect to CaCO3 below

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about 500 m everywhere; the reason carbonate is present down to thousands of meters is basically a time-lag effect. Presumably the CCD has to do with what happens on the bottom, not what happens on the way down, because it doesn’t take long for the tests to get to the bottom, but they sit there exposed to the water for a long time. 6.6 The distribution of calcareous ooze is very irregular. Calcareous ooze is present mainly on high areas in low latitudes, and it’s present to greater depths where surface productivity is high. There’s not much under areas of low productivity, where nutrient concentrations in the surface waters are low, or at high latitudes, where surface waters are unfavorable. 6.7 Forams are the major constituent of calcareous ooze, but not the only one. Two other kinds of carbonate-secreting planktonic organisms also contribute to calcareous ooze: • a minor group of tiny gastropods, called pteropods and heteropods (mostly the former); these are larger than forams but they secrete aragonite, so they are not found below about 3500 m; • two kinds of calcareous algae, called coccoliths and rhabdoliths (mostly the former). 122 6.8 In a few areas pteropods or coccoliths predominate in the ooze (then the sediment is called pteropod ooze or coccolith ooze), but usually they are just a minor admixture in foram ooze. Figure 5-5: (left) Brasier, M.D., 1980, Microfossils: George Allen & Unwin, 193 p. (Figure 8.3a, p. 48). (upper center) Brasier, M.D., 1980, Microfossils: George Allen & Unwin, 193 p. Haq, B.U., 1978, Calcareous nannoplankton, in Haq, B.U., and Boersma, A., Introduction to Marine Micropaleontology: Elsevier, 376 p. Herman, Y., 1978, Pteropods, in Haq, B.U., and Boersma, A., Introduction to Marine Micropaleontology: Elsevier, 376 p. (shows what pteropods and coccoliths look like. 6.9 Calcareous ooze is a tan-ccolored to cream-colored sediment, gritty and clean-feeling; it’s not at all as repulsive as its name suggests. In the deep ocean it grades laterally into the noncalcareous brown clay of the deepest ocean.

7. PLATFORM CARBONATES 7.1 Introduction 7.1.1 In many places in the world today, carbonate sediments are accumulating on platforms along continental margins where siliciclastic deposits are absent. The two most important factors that control this accumulation are • lack of siliciclastic input, and • high biogenic carbonate productivity. Unfortunately the range of sedimentary environments represented by such areas is less extensive than must have prevailed at many times in the geologic past. That makes it more difficult for carbonate sedimentologists to interpret ancient carbonate rocks than it is for siliciclastic sedimentologists to interpret ancient siliciclastic rocks—because they don’t have a wide a range of modern sedimentary processes to use as a basis for interpreting the ancient rocks. 7.1.2 Because the precipitation of carbonate is easiest in warm,

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shallow seawater, most carbonate production takes place on tropical platforms within a fairly restricted range of shallow subtidal water depths. These environments have come to be called, picturesquely, by carbonate sedimentologists the carbonate factory. Although the most of the sediments that are produced in the carbonate factory remain in the source area, some are transported landward and some are transported basinward (Figure 5-6). Thus, there are three zones of accumulation: 123 • the subtidal open shelf and shelf margin, characterized by in-place accumulations of carbonate sands, carbonate muds, and reefs; • the shoreline, where sediments are transported from the open shelf onto beaches and tidal flats; and • the slope and basin, where shelf-edge sediments are transported seaward, often by mass movements, and redeposited at depth. 7.1.3 The shallow-water carbonate factory is very sensitive to sea-level change. At various times in Earth history, sea level has fluctuated, on time scales of tens of thousands to hundreds of thousands of years, and with magnitudes ranging from meters to hundreds of meters. Most of the major carbonate-secreting organisms flourish when the water is shallow. If sea-level rise is slow, and the concomitant increase in water depth is slow, the carbonate factory can keep up production; this is called keep-up mode. But if sea level increases fast enough, and water depth thereby increases fast enough, the carbonate factory has a strong tendency to shut down. This is called give-up mode. 7.2 Regional Geometry 7.2.1 The term platform carbonate is in common use for all accumulations of carbonate sediments in tectonically stable shallow-water environments. This includes reefs, but I’ll deal with them later. Because in areas of high productivity carbonate sedimentation can be so rapid, carbonates tend to mold their own environment, even on the scale of entire shelves. So the regional bathymetry of carbonate areas tends to be more varied than that of siliciclastic shelves, which are more familiar to most geologists. Below is an outline of the regional geometry of shallow-water carbonate bodies. Refer to.2.2 Carbonate ramps are major carbonate bodies built far outward from land to have gentle regi onal sea-floor slopes. Carbonate shelves are large flat topped carbonate bodies extending seaward from land areas. Ramps tend to evolve into platforms where carbonate productivity is high. Isolated platforms (or offshore carbonate banks) are large carbonate accumulations built up over local high areas offshore of continental land areas; they typically develop approximately flat tops and steep margins. Carbonate workers use the term shelf a little differently from others to be the flat upper surface of carbonate platforms or banks. Carbonate buildups are smaller-scale or more local carbonate bodies showing topographic relief; they are

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commonly present at shelf margins, but they are formed in other places as well. They usually result from the rapid growth of a community of calcareous organisms and include the well-known group of coral reefs. 7.3 The Bahamas 7.3.1 Few places in the world today are very representative of the vast areas of shelf carbonate deposition at various times in the geologic past. Some of the carbonate-producing regions are compared in Figure 5-11. An excellent place to study shelf carbonates today is on the Bahama Banks. The Bahamas are an extensive and outstanding example of warm, shallow, extensive carbonatedepositing seas. The following is a brief account of the Bahama Banks. 126 7.3.2 The Bahama Islands are located on several submerged shallow carbonate platforms that lie just off the North American continental shelf. The banks cover 60,000 square miles, but the land area of the islands is only about 4400 square miles. Water depths over most of the banks are less than ten meters! There are several large islands and thousands of very small islands called cays. The exposed land surface is very pure Pleistocene limestone with little soil development. In places elevations are over a hundred meters; these are fossil subaerial dune ridges of oolitic sediment. Keep in mind that the entire surface area of the banks was exposed during the low stands of sea level during the Pleistocene. 127 7.3.3 The largest of the Bahama Banks, Great Bahama Bank, is split by two troughs, Tongue of the Ocean and Exuma Sound, which extend down to deep ocean depths, and each of the banks is separated from its neighbors by deep-water channels. The shelf margins are well defined. The upper few hundred meters of the slope is very steep, greater than the angle of repose; this is probably controlled by reef development in the past. 7.3.4 The Bahamas lie in the belt of northeast trade winds. This governs the position of islands on the windward margins of the platforms. Hurricanes are common, and they do important geological work. Throughgoing currents are warmed as they pass over the banks, leading to supersaturation with respect to calcium carbonate. 7.3.5 Bahama reefs play a minor part in the carbonate sediment picture today, presumably because of the Pleistocene history of fluctuating sea level. But deep borings show that they were much more important in the past. Reef construction is most important along the outer edges of the banks facing deep ocean water, and they are much better developed on the windward sides than on the leeward sides. 7.3.6 Great deposits of pure calcium carbonate sand and mud are forming and accumulating on the banks. With stable sea level, the carbonate sediment is transported to deep water as fast as it is produced, but slow crustal subsidence has caused about 5 km of carbonate sediment to have been deposited since at

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least the early Cretaceous, so the banks date from not long after the opening of the Atlantic. 7.3.7 All of the various kinds of shelf-carbonate constituents described in the next section are found on the Bahama Banks, most of them in abundance. Areas of ooids and of aragonite muds in particular are shown on the sketch map in Figure 5-12. 128 7.4 Constituents 7.4.1 A great variety of sedimentary carbonate particles are produced in carbonate depositional environments. Here’s a summary. I’ll leave the broader patterns of depositional facies aside for now. Skeletal Grains Many kinds of marine invertebrates precipitate calcium carbonate to form their skeletons. There is an enormous range in size, shape, internal structure, and composition. Crystal size within the skeletal material ranges from microscopic to large single crystals. Some is aragonite, and some is calcite; the calcite itself ranges from magnesium-free to as much as 25% Mg (in the case of red algae). Grain shape depends on the skeletal geometry of the particular species, and, in the case of colonial organisms, on the style of colonial development as well. In the words of Dunham (1962), “Carbonate grains are shaped more like twigs or potato chips than marbles”. Some of this skeletal material is buried in place; some is transported, and can undergo various degrees of breakage and wear. I won’t even attempt to illustrate the enormous range of geometry and structure of skeletal grains; you will be seeing many in both hand specimen and thin section. 129 Ooids Ooids are nearly spherical grains consisting of a grain of calcareous or noncalcareous material serving as a nucleus around which successive layers or shells of calcium carbonate are precipitated or accreted while the particle is moved in flowing water that is supersaturated with respect to calcium carbonate: Tucker, M., 1991, Sedimentary Petrology; An Introduction to the Origin of Sedimentary Rocks: Blackwell, 260 p. Williams, H., Turner, F.J., and Gilbert, C.M., 1982, Petrography; An Introduction to the Study of Rocks in Thin Section: W.H., Freeman, 626 p. Size ranges mostly from a fraction of a millimeter to about 2 mm— although, especially in the Neoproterozoic, they can be on a centimeter scale. If the coating is thin relative to the size of the nucleus, the ooids are called superficial ooids or coated grains. Commonly there are two kinds of concentric spherical shells in the ooid structure: tangentially oriented aragonite needles, and non-oriented cryptocrystalline aragonite. Ooids tend to contain organic matter in the form of algal mucilaginous matter; this is seen as dark brown areas in thin section. Ooids can also show radial magnesian calcite. This seems to result from purely inorganic precipitation. Ooids are common in the ancient and are known but not common in modern carbonate sediments. The processes of ooid growth are still a subject of discussion. Pellets Pellets are

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rounded grains of very fine-grained aragonite and calcite, a few tenths of a millimeter to about a millimeter in size. They are usually elliptical or ovoid in shape, but they may be broken to form beehive-shaped grains. Some are clearly fecal pellets excreted by such organisms as worms, gastropods, mollusks, and crustaceans. These are soft and friable at first, but they soon become well cemented. Other pellets seem to be formed by cementation and rounding of friable irregular aggregates of aragonite silt. Because it is usually difficult or impossible to distinguish among the various possible processes that form such objects, the term peloid is in common use. Intraclasts Intraclasts are fragments of carbonate sediment, usually fine-grained, that was deposited and then later ripped up by strong currents to be redeposited with other carbonate sediment. The derivation of the word implies that the ripping up took place within the environment of carbonate deposition, geologically soon after the depositional of the sediment; don’t confuse these with carbonate rock fragments in a largely siliciclastic conglomerate. The stage of cementation varies considerably. Commonly the intraclasts are tabular, reflecting breakage of semiconsolidated sediment along stratification planes. 130 Carbonate Mud Carbonate mud consisting mostly of needle-shaped aragonite crystals is common in areas with weak currents. Some of this carbonate mud is produced by abrasion of larger grains, but most seems to be precipitated directly from seawater. The nature of this precipitation has been controversial: is it purely inorganic, or is it caused by algae? (You can imagine precipitation of aragonite next to the bodies of photosynthesizing algae, because the CO2 content of the water right next to their bodies is decreased, which favors carbonate precipitation.) The answer seems to be that both processes operate, but that algal precipitation is generally much more important than inorganic precipitation. 8.

REEFS 8.1 Introduction 8.1.1 A reef, rising above the sea floor, is an entity of its own making—a sedimentary system within itself. The numerous, large calcium-carbonatesecreting organisms stand upon the remains of their ancestors and are surrounded and often buried by the skeletal remains of the many small organisms that once lived on, beneath, and between them At present, far more shallow-water carbonate sediment is produced in, or in connection with, reefs than by any other means. Judging from the presence of reefs in the stratigraphic record, reefs have been important sites of carbonate sedimentation from the Archean. Throughout geologic history, they have 131 experienced various periods of importance and decline in terms of their absolute abundance. Furthermore, as organisms have evolved through time the record of this is reflected in the history

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of reefs that have been dominated by different communities at different times. However, in most cases, these organisms have had very similar functional morphologies. 8.2 Terminology 8.2.1 In the most common sedimentological usage, the term reef is used for a marine structure built by organisms and with a framework strong enough to withstand attack by waves. But the usage is confused; some use the term for all kinds of carbonate buildups, the above kind being just one of many possibilities. And the nonscientific meaning of the term reef is just a shoal of any kind on which a ship can go aground. So it’s best to preface the term with a modifier: organic- 132 framework reef is probably the best. And there are several kinds of such reefs, depending upon geographic setting. 8.3 General Stuff 8.3.1 Reefs are produced by growth of colonial carbonate-secreting organisms. But there’s more to it than that. Marine life is abundant in warm shallow seas, but generally it has to adapt itself to the physical environment. But under certain conditions, certain kinds of organisms can create their own environment by building major structures that alter the local marine environment fundamentally. 8.3.2 Reefs need the strong waves and currents that act upon them: without the strong water movements to clean them now and then of loose sediment and to provide them with nutrients from the open sea, reefs would not develop in the first place. The geological significance of this is that fragmentation by the destructive forces of storms is a natural concomitant of reef growth. 8.3.3 Reefs have been of special interest to petroleum geologists because reef rocks in the stratigraphic record have often ended up being petroleum reservoir rocks: they tend to have high porosity, and they usually become encased in finer and much less permeable sediment, so they make good stratigraphic traps. 8.3.4 You probably think of reefs as being formed by corals; to many, the term “reef” is synonymous with “coral reef”. Corals are indeed the chief reefbuilding organisms today, but calcareous algae are major contributors, and in the past algae, bryozoans, and archeocyathids were at times the major reef builders. It’s better to think in terms of organic reefs than of coral reefs. 8.3.5 The coral polyp, the coelenterate that sits in and on a cup-shaped calcareous skeleton, feeds by filtering plankton. The polyps of reef-forming corals contain algae called zooxanthellae, mainly in the cells of the covering of the polyp. The polyps and the algae are symbiotic: the algae receive nutrients and CO2 from the polyp, and the polyp receives oxygen from the algae. The polyps feed at night, and during the day they are partly closed to put the algae in the best position for photosynthesis. 8.3.6 Colonies of polyps develop by budding or fission. All of the polyps stay in contact with one another, so the colony is like a quasi-animal. The growth forms

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of coral colonies vary widely, even with the same genus: dome-shaped, mushroom-shaped, and branching in all sorts of ways. 8.3.7 Important: not more than a few tens of percent of the reef volume is the rigid coral framework. There are abundant tunnels and cavities in the framework that are partly filled with other calcareous biogenic sediment. Most reef sediment is produced by the post-mortem disintegration of organisms that are segmented (crinoids, calcareous green algae) or non-segmented (bivalves, 133 brachiopods, foraminiferans) and that grow in the many nooks and crannies between the larger skeletal metazoans. The rest of the sediment is produced by the various taxa that erode the reef: boring organisms (bivalves, worms, sponges) produce lime mud; rasping organisms that graze the surface of the reef (echinoids, fish) produce copious quantities of lime sand and silt. 8.4 Controls on Reef Growth 8.4.1 Reef growth is mainly a function of the factors that determine whether the coral colonies can flourish and increase rapidly in volume: Water Temperature. Winter water temperatures are rarely below 18° C over reefs; corals don’t grow in profusion until the temperature is 25–30° C. Temperatures above 30° C are above the optimum for coral growth, and temperatures above about 35° C kill corals. But temperatures above about 30° C are rare in the open ocean, so the chief temperature factor is absence of cold water. Reefs are rare outside the tropics, and they are more common in the western than in the eastern parts of all three major oceans, because of cold equatorward-flowing currents in the eastern parts. Water Depth. The depth to which reef-forming corals can thrive is a function of light penetration, because the symbiotic algae need light for photosynthesis. Therefore it's not depth alone that limits coral growth, but light penetration. In exceptionally clear water a few species of reef-forming corals can live at 100 m, but for most the limit is 70–80 m. With the usual turbidity around coral reefs, reef corals don’t flourish deeper than 50 m, usually less. Probably the depth at which the greatest volume of new coral sediment or rock is added per year is only a few meters below mean sea level. There are deep-sea corals that are independent of light, but these don’t develop colonies and build reefs. The distribution of reef building also relates to water depth; platy forms develop at greater depths, massive and branching forms at shallower depths, and encrusting forms face the reef front. Nutrients. The coral colony needs a steady supply of food. Plankton are most abundant where nutrients in solution are carried in by currents from the open ocean. Algae produce oxygen during the day, but corals need oxygen at night too, and that has to be transported in by currents as well. Currents are thus favorable to reef growth. Salinity. More or less normal salinity is required for reef

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growth. Reef corals live within the salinity range 27 to 40‰. Reefs can be killed by great floods of fresh water sweeping over them from land. Fine Sediment. Fine sediment restricts penetration of light, and it also hampers growth when it settles on the colonies. Probably the effect of light is more important than smothering, because corals can readily clean themselves off. 134 Also, corals need a suitable substrate on which to develop: coral colonies can’t develop on a uniformly muddy floor. 8.5 Kinds of reefs Fringing Reefs: these grow directly against a rocky coast. Width varies up to a few hundred meters.. Atolls: ring-shaped reefs, unconnected with land. These presumably developed as fringing reefs around oceanic islands that later subsided while reef growth kept pace with subsidence.. Barrier Reefs: these, the largest and most important, are reefs separated from the continental coast by a lagoon that is too deep for coral growth and up to a hundred kilometers wide, floored at least in part by siliciclastic sediment. (The Great Barrier Reef of Australia, following the coast of Queensland for 1500 miles and as much as 150 miles offshore, is of this kind.). 8.6 Morphology and Sedimentology 8.6.1 shows a typical cross section through a reef, normal to the reef front. Lagoon: The lagoon is the relatively protected and shallow area behind the reef. Smaller lagoons are floored entirely by finer carbonate sediment; lagoons behind barrier reefs can have mostly siliciclastic sediment. Lagoons often have numerous coral knolls, or isolated coral colonies, each a miniature reef itself. Depths in lagoons are seldom more than 50 m. There is usually a great variety of 136 sediment types in lagoons: typical sediments are coral debris from the living reef, foraminiferal and algal sediment, skeletal fragments of larger invertebrates, and aragonite mud in the deepest parts. Reef Flat: Behind the growing reef front is a broad expanse of dead reef rock with a flat surface, partly or wholly emergent at low tide. There are patches of sand and coral rubble and scattered small coral colonies, as well as shallow pools, irregular gullies, and potholes. Islands are common on the reef flat; these could be built by storms or be left from a higher stand of sea level. Most of the major islands on reef flats today are a few meters above sea level and are undergoing erosion; they probably were formed by the higher stand of sea level during times of milder climate within the last few thousand years. Reef Front: The heart of the reef is the growing reef front, a narrow zone of flourishing coral colonies. Waves and storms beating against the coral colony keep breaking material away, throwing it onto or over the reef flat or carrying it down the outer reef slope. It’s here that most of the reef growth takes place. Outer Slope: Often the reef descends almost vertically for a hundred meters, then at slopes as much as 30° down to deep ocean depths. The reason for

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the steepness of the slopes is that the coral colonies grow fastest in their uppermost part, tending to produce an overhang. The lower slopes are a jumble or talus of jagged broken reef material. 8.7 Bioherms and biostromes 8.7.1 Whenever you see large carbonate skeletal deposits in the sedimentary record, you have to think about whether they were part of a reef, or, more specifically, whether they formed a mass that stood up from the sea floor or whether they had no relief but built up in thickness along with the surrounding sea floor. A bioherm is the general term for carbonate buildups, large or small, produced either by in-place production of large numbers of individual organisms or by colonial framework building or encrustation, as in reefs. These structures had substantial positive relief on the sea floor. They have low width-to-height ratios (W:H < 30) 8.7.3 A biostrome is the general term for blanketlike deposits of carbonate sediment produced by in-place organisms and surrounded by other sediment types. There are shell biostromes (beds of unsorted and nontransported skeletal remains that grew and died in place) and algal biostromes (soft, sticky mats of sediment-binding algae, which catch fine carbonate sediment and then grow upward through that sediment to reestablish the mat). They have larger width-toheight ratios (W:H > 30). 8.7.4 Finally, the general term in use for any carbonate body that develops in such a way as to stand above the surrounding seafloor is called a buildup. The essential feature of a buildup is that it has topographic relief above the seafloor.

9. LIMESTONES 9.1 Definition 9.1.1 The official definition of a limestone is a sedimentary rock that contains at least 50% carbonate minerals, of which at least 50% is a calcium carbonate mineral. Most limestones you will deal with have far more carbonate material than siliciclastic material: approximately equal mixtures are not as common as dominantly carbonate or dominantly siliciclastic—although approximately equal mixtures are not vanishingly uncommon.

9.2 General Stuff

9.2.1 Limestones are a highly varied group of rocks. All they have in common is their composition. Some are grown in place, as reefs, and others are deposited as particles, analogous to shale, sandstone, and conglomerate. Many geologists have thought of limestone as a "wastebasket" term, and have tried to get rid of it by coining new and more specific words for the various kinds--but they have not been successful. The keys to understanding most limestones are these: • Most carbonate sediments started out as accumulations of particles, large and small,

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analogous to siliciclastic gravel, sands, and muds. These particles are produced locally in the carbonate environment, and they are deposited with or without subsequent transportation. So you can think about most carbonate rocks in terms of framework, matrix, and cement, just as with sandstones, but keep in mind that most if not all the material is just carbonate in composition. • Cementation of carbonate sediments is usually very early; little burial is needed. So reduction of pore space is much more a matter of filling by carbonate 138 cement than by deformation of framework grains. Cementation commonly occurs before compaction of sediments. • Carbonate rocks are much more susceptible to diagenesis, early and late, than siliciclastic rocks. Diagenesis obliterates primary structures and textures in carbonate rocks long before siliciclastic rocks. This is due to their greater solubility in diagenetic pore fluids. 9.3 Classification 9.3.1 In the late 1950s, Folk revolutionized the classification of limestones by proposing that most limestones can be treated as particulate rocks involving framework, matrix, and cement just like siliciclastic rocks. Although the terms he coined for the various classificatory categories are not widely used these days, his ideas have been of great significance.

9.3.2 Folk called the coarse clastic constituents allochems. These are the kinds described in an earlier section: intraclasts, skeletal grains, ooids, and pellets. He used the term micrite for the carbonate mud (short for microcrystalline calcite mud; something of a misnomer because it’s now known that most of the mud is aragonite, not calcite) and spar for the sparry calcite cement. where most limestones fall in a ternary composition diagram. one version of Folk’s classification, Folk, R. L., 1959, Practical petrographic classification of limestone: American Association of Petroleum Geologists, Bulletin, v. 43, Folk, R.L., 1962, Spectral subdivision of limestone types, in Ham, W.E., ed., Classification of Carbonate Rocks: American Association of Petroleum Geologists, Memoir 1, p. 62-84. are two of Folk’s original figures, showing about the same thing as above but in more detail. 139 Sparry rx Micritic rx Intrasparite Oosparite Intra micrite Oomicrite Oolites > 25% F>P P>F Micrite No allochems Pelmicrite Biomicrite Oolites < 25% Intraclasts < 25% Intraclasts Intraclasts > 25% Oolites Fossils Milits No allochems Pelsparite Biosparite Figure by MIT OCW. 9.3.3 In the early 1960s Dunham developed a simpler and broader classification using Folk’s concepts; the names are in very wide use today. shows Dunham’s original classification table, and a later modification by Embry and Klovan. Depositional Texture recognizable Original components not bound together during depositions Contains mud (particles of clay and fine silt size) Mud-supported Grainsupported More than 10%

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grains Less than 10% grains Mudstone Wackstone Packstone Grainstone Lacks mud and is grainsupported Depositional texture not recognizable Crystalline carbonate Original components were bound together during deposition... as shown by intergrown skeletal matter, lamination contrary to gravity, or sediment-floored cavities that are roofed over by organic or questionably organic matter and are too large to be interstices. (Subdivide according to classifications designed to bear on physical texture or diagenesis.) Boundstone Figure by MIT OCW. 140 Figure 5-22: Folk’s classification of limestones Dunham’s classification of limestones The key terms here are grainstone, packstone, wackestone, and mudstone. Get used to trying to apply these names when you look at a limestone. They are by far the most common terms used by practicing sedimentary geologists to describe limestones. Dunham’s term boundstone, also important to remember, applies to a carbonate rock, typically a reef rock, whose original constituents were bound together during deposition and subsequently remained mostly in position of growth. A few others, however, not covered by Dunham’s classification, are also common: a rudstone is a carbonate rock composed of bioclasts or other carbonate fragments, over 2 mm in diameter, that are close-packed and in physical contact. Rudstones are the coarse-grained equivalents of grainstones and packstones— terms that are customarily restricted to sand-size carbonate clasts. A floatstone is a carbonate rock containing a small percentage of bioclasts or other carbonate fragments over 2 mm in size, that are widely spaced, embedded in (“floating in”) finer carbonate sediment. Floatstones are the coarse-grained equivalents of wackestones. A bindstone is a kind of boundstone that is composed of sheetlike colonies that encrust large fossil fragments or finer carbonate sediment, forming a layered mass that is partly in-place skeletal, like boundstone, and partly bioclastic, like packstone or wackestone.

9.4 Microbial Carbonates

9.4.1 The foregoing material neglects a very important kind of carbonate depositional environment and the limestones and dolostones deposited therein: microbial mats. 9.4.2 Since far back in the Precambrian, certain kinds of microbes (mainly filamentous and coccoid cyanobacteria) have developed extensive mats in very shallow marine environments, ranging from very shallow subtidal up through intertidal. These mats present a dense network of growing filaments exposed to the surface. Fine carbonate sediment (aragonite needles, or dolomite crystals to which the aragonite has been changed within the sedimentary environment) is now and then washed onto the mat and adheres to the sticky surface. The

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filaments then grow upward among the sediment particles to reestablish the mat. The result is a laminated fine carbonate deposit. In some cases, the microbes cause precipitation of carbonate directly on the surface, which accretes upward over time. 9.4.3 Depending upon environmental conditions that are still not well understood, the surface geometry of the microbial mat can either be planar or nearly so (these deposits are called microbial laminites) or take on a variety complicated geometries involving heads or domes of various shapes. These structures are called stromatolites. Lamination in these stromatolites is characteristically convex upward, although there are some striking exceptions to this. At the margins of the stromatolites the laminae dip at angles greater than the angle of repose, which is a tip-off for the cohesive or precipitated nature of the deposition. In some cases the laminae end abruptly at the margins of the 141 stromatolites. It was originally thought that stromatolites could be treated as actual organisms, and were given names like Cryptozoon, Anabaria, and Tungussia, based on variations in their form and lamination fabrics. Although such names are still used, it is recognized that the range of morphologies and lamination fabrics is due more to environmental factors than to biologic factors. Some stromatolites have poor lamination and a clotted texture; these are called thrombolites. Collectively, microbial laminates and stromatolites are termed microbialites (pronounced mike-ROW-bee-al-ites, not mike-row-BUY-al-ites!).

9.4.4 In slightly different environments, especially with strong and regular currents, microbes bind or precipitate fine carbonate sediment around small spheroids that roll around while accreting concentrically. These spheroids are called oncolites (also called oncoids or microbial accretionary grains. They are analogous to ooids, but they are different in origin, and usually they are larger. They are usually from a few millimeters to more than a centimeter in size. When larger, up to ten centimeters, they are called microbial biscuits. 9.4.5 Telling microbial laminites from non-microbial laminites in thinly laminated carbonates is a tricky business. Microbial laminites tend to show less regular lamination, and they sometimes show a characteristic structure known as fenestrae: little irregular cavities elongated parallel to bedding, filled with carbonate cement or sometimes even geopetal fine clastic sediment. These cavities seem to have developed during gas production in response to photosynthesis or respiration by the microbial community, or during occasional drying and shrinkage of the mat.

10. STRUCTURES IN LIMESTONES

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10.1 Many limestones have bed configurations (ripples and dunes) and cross stratification much as in siliciclastic rocks. That should not surprise you, inasmuch as carbonate grainstones, especially, consist of particles that were transported by water currents. Certain other structures, however, are specific to limestones. Below are descriptions of two of these. (There are others that could be added here as well.) hardgrounds: Hardgrounds are surfaces or horizons with synsedimentary cementation that develops at, or slightly below, the sediment surface. It’s a feature that develops underwater, on the seafloor. The surface is typically encrusted by sessile benthic organisms (that is, organisms that live attached to the seafloor). Species of various phyla of marine invertebrates have that mode of life. The sediment lying immediately below such surfaces also tend to be bored by organisms that make tunnels into the sediment. The burrows are manifested by slight differences in the nature of the sediment that fills the burrows. Hardgrounds are significant because they record times when the rate of sedimentation was low, or zero. They often mark unconformities. 142 paleokarst: You probably have heard of karst. Karst, a geomorphic feature, is a kind of topography that develops on limestone terranes by solution when the land surface in a region with abundant rainfall lies well above the groundwater table. The surface manifestation of karst is a distinctive pattern of sharp hills and bowl-shaped depressions, that latter commonly floored by sinkholes. Below the surface, karst is associated with networks of caves and caverns. In a limestone succession, karst (paleokarst, that is) can be recognized by the presence of a sharp, irregular surface that caps a zone of dissolution marked by solution breccias and irregular solutional passageways extending often far down into the underlying succession. The great significance of paleokarst is that it marks a time of emergence: a transition from a submarine depositional setting to a subaerial erosional setting.