Introduction To Geology

Sedimentary rocksIntroductionThe article constitutes a brief introduction to the various kinds of sedimentary rocks. Further information about the sources of sediment, its transport, and it deposition, will be covered in further articles; indeed, in much of the rest of this textbook.Types of sedimentary rocksSedimentary rocks can be divided into three main classes:* Clastic sedimentary rocks are formed from sediments created by rocks being physically broken down into small particles (clasts).
* Chemical sedimentary rocks are formed from sediments created by dissolved chemicals being precipitated out of the water they're dissolved in.
* Biochemical sedimentary rocks are formed from sediments consisting of dead organisms, or parts of dead organisms.In some schemes of classification, biochemical sediments are treated as a sub-class of chemical sediments, but this leaves one with the awkward question of what to call chemical sediments which aren't biochemical. For this reason we shall treat them as two non-overlapping classes.Before we review the main types of sedimentary rocks, it is worth mentioning the process by which they turn into rock: this is known as lithification. In some cases, such as shale, mere compaction, along with the resulting loss of water, is sufficient. Coarser sediments, such as sandstone, are both compacted and cemented, as can be seen under a microscope. The cements are minerals precipitated out of the water in which they are dissolved: silica and calcium carbonate are the commonest forms of cement, with iron oxides and hydroxides coming a distant third.In the sections below we shall list the main types of sedimentary rock.Clastic sedimentary rocksAs we have said, sedimentary rocks are formed from broken pieces ("clasts") of pre-existing rocks.* Sediment: rounded gravel. Rock: conglomerate.Gravel is defined as clasts with diameter 2mm or more. Conglomerate is formed of rounded gravel. Note that when geologists speak of rounded clasts, they do not necessarily mean that they are round like a ball, but merely that the sharp corners and edges have been worn off them by erosion. Conglomerates are rocks formed mainly from rounded gravel which has been compacted and cemented together.* Sediment: angular gravel. Rock: breccia.Breccia is like conglomerate except that the gravel is angular: that is, it has not been rounded. This reflects a different history, since gravel that has been transported any appreciable distance by water, or which has been rolled about by waves on a beach, will quickly have its corners and edges worn away.* Sediment: sand. Rock: sandstoneSand is defined as clasts less than 2mm and more than 1/16mm in diameter. Sandstone is sand that has been cemented together.Most sandstone is quartz sandstone; that is, it consists of grains of quartz. This is because the process known as chemical weathering dissolves many rock-forming minerals, or, in the case of feldspar minerals, converts them to clay, leaving behind only the quartz from the original rock.Arkose sandstone is sandstone with an appreciable proportion of feldspar minerals in it. This reflects a somewhat different history to quartz sandstone, in that it must have been formed when mechanical weathering (the physical process of breaking rock into small fragments) has predominated over the chemical weathering that would otherwise have converted the feldspar minerals to clay minerals.Greywacke is sandstone that, in addition to quartz and feldspar, also contains sand-sized fragments of igneous or metamorphic rocks. Similar remarks apply to greywacke as to arkose sandstone.* Sediment: silt and/or clay. Rock: mudrock.Silt is defined as clasts between 1/16mm and 1/256mm in diameter.The term "clay" is a little ambiguous. In the classification of sediments, it is defined as particles less than 1/256mm in diameter. However, in mineralogy, clay is a class of minerals (technically, hydrous aluminosilicates). In practice, this need cause no confusion, because what is clay by size will be overwhelmingly clay by composition.Mudrocks can then be divided into siltstone (formed from silt sediments); mudstone (from sediments that are a mixture of silt and clay); and claystone (from clay sediments).* Sediment: bedded mud or clay. Rock: shale.Most mudstone and claystone is bedded (see the entry above on sandstone for a definition of this term). When it is, it is referred to as shale.Chemical sedimentary rocks* Sediment: salt. Rock: halite.Halite, also known as rock salt, is an evaporite, formed by the evaporation of salt water.It can be formed by complete evaporation of salt water, as seen, for example, in desert salt flats. However, complete evaporation is not necessary; it is sufficient that enough water should evaporate that the remaining water can't hold all of the salt in solution; so halite can also form in shallow seas or salt lakes in a hot environment.* Sediment: hydrated calcium sulfate. Rock: gypsum.Calcium sulfate is another substance to be found dissolved in salt water, and gypsum, like halite, usually forms as an evaporite under pretty much the same circumstances.* Sediment: silica. Rock: chert.Dissolved silica can precipitate out of the water in which it's dissolved to form chert. Note, however, that chert is more usually formed as a biochemical sedimentary rock.* Sediment: calcium carbonate. Rock: limestone.Calcium carbonate, like silica, can precipitate out of water to form limestone. Sometimes it forms tiny spheres called ooids, which form around grains of sand or fragments of shell, which are then cemented together by the further precipitation of calcium carbonate; such limestone is known as oolitic limestone.Most limestone, however, is biochemical sedimentary rock, formed from shells or coral.Biochemical sedimentary rocks* Sediment: shells of calcium carbonate. Rock: limestone.Most limestone is formed from tiny hard parts of marine creatures which build their shells out of calcium carbonate; these settle on the sea floor to form a calcareous ooze. Chalk is an example of such a rock: the tiny fossils that compose it can be clearly seen and identified under a microscope.The hard parts of coral reefs are sometimes preserved intact, giving us reef limestone.* Sediment: shells of silica. Rock: chert.While calcium carbonate is the most popular substance to make shells out of, some organisms such as diatoms and radiolarians build their shells out of silica; these settle on the sea floor to form siliceous ooze which, when compacted and cemented, forms chert.* Sediment: peat. Rock: coal.Peat is plant material laid down in oxygen-poor conditions, so that it doesn't entirely decompose. Pressure, and the higher temperatures which come with deep burial, can then convert it into coal.Modes of depositionIn the sections above we have principally divided sedimentary rocks by their composition. We can also classify them by their modes of deposition: for example, aeolian (deposited by the wind), or fluvial (deposited by rivers) and so forth. We shall have a lot more to say about this in subsequent articles.BeddingSedimentary rock often exhibits bedding: that is, the rock has distinct layers in it and is fissile: that is, it splits more easily at the divisions (bedding planes) between the layers. In cross-bedded rocks, the layers are not flat but lie at an angle to the horizontal, as a result of the original sediment being formed into dunes or ripples by the action of wind or water.The picture below shows a particularly large-scale example of cross-bedding in sandstone. (To give you an idea of the scale, the small patches of green visible in the top right corner are not moss, as you might guess, but bushes.) How do we know?How do we recognize sedimentary rocks as sedimentary? How do we recognize the sediments that compose them and the manner of their deposition?Such questions will be answered later in this course one type of sediment at a time. At present we shall content ourselves with sketching out a general answer.In the first place, the rocks look just like we would expect if sediments became lithified; for example sandstone looks like it's made of sand: everything about the size, the composition, and the erosion of the grains of which it's composed is in agreement that what we're looking at is grains of sand cemented together.Secondly, we can drill down and take cores of sediments, and we can see, as depth increases, how sloppy muddy ooze on the surface grades into hard mudstone with no sharp dividing line between them; similarly we can see calcareous ooze grade into solid limestone.Then again, all types of sedimentary rocks can contain fossils (including, as we have remarked, those rocks which consist of fossils). This is consistent with the processes of burial of organic remains in sediment which we can see going on today.Trace fossils are also a strong argument: when we find, in shale, the recognizable fossil footprints of land animals, it is hard not to conclude that what we are looking at once lay on the surface and was soft enough to take impressions such as we can see being made in mud today.These considerations also allow us to figure out where the sediments were deposited: on land, in fresh-water, or in the sea.The sedimentary structures within the rocks, such as bedding and cross-bedding, can be seen today in oozes forming on the sea floor; in sand-dunes; in ripples caused by tidal action, and so forth. Again, consideration of these structures will allow us to make the deduction, not merely that the rock is sedimentary, but also about the method of its deposition; if, for example, we find sedimentary structures such as can only be formed by tidal action, we are forced to infer that we are looking at lithified nearshore sediments.Also, we may observe the topographic patterns of deposition. For example, when we see sedimentary rocks which because of their structures and fossils we associate with the land divided from sedimentary rocks which because of their structures and fossils we associate with the sea by a long thin strip of sedimentary rocks which because of their structures and fossils we associate with the nearshore, then this observation confirms our identification of the rocks as being terrestrial, marine, and nearshore rocks. If, on the other hand, we found alternating bands of marine and nearshore rocks, this would tend to falsify our theories. The fact that the topography of sediments is always consistent with our theories is therefore a point in favor of their correctness.So the conclusion that sedimentary rocks are, indeed, sedimentary in origin, is a safe one; and we are certainly not without clues as to the manner of their deposition.Note on vocabularyConglomerates and breccias are sometimes called rudaceous rocks; sandstones are sometimes called arenaceous rocks or arenites; and mudrocks are sometimes called argillaceous rocks.The rocks which we have called clastic are sometimes called detrital.As usual, we shall employ a consistent vocabulary in this text; these terms have only been supplied for the benefit of the reader who wishes to pursue a course of further reading.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.

Metamorphic rocksIntroductionMetamorphic rocks are those in which a pre-existing rock (the parent rock) has been altered chemically or texturally by heat and/or pressure. Types of metamorphismThere are two main types of metamorphism; contact metamorphism and regional metamorphism.In contact metamorphism, the heat of magma intruding through rocks causes metamorphism to the surrounding rocks, leaving an aureole of metamorphic rocks around the igneous rocks formed from the magma.Regional metamorphism is caused by tectonic events involving both heat and pressure, and will affect large, elongated regions of rock.Besides these broad classifications, we can also class metamorphic rocks according to their grade of metamorphism; the temperature to which they have been exposed. As we shall see, this determines the chemical changes that metamorphic rocks undergo. Grades of metamorphism range from very low (below 300C) to low (300C - 500C) to medium (500C - 600C) to high (600C and upwards). Chemical changesMetamorphism often causes chemical changes to the minerals of which the affected rocks are composed. For example, at low temperatures clay minerals will be converted into chlorite; at higher temperatures the chlorite will itself be transformed into other minerals. Hence, if we find a rock with chlorite in it, we know that it has undergone low-grade metamorphosis. Minerals such as this, which reveal the grade of metamorphosis, are known as index minerals.Geologists can correlate index minerals with grades of metamorphism because it is simple enough to repeat the processes of metamorphism in the laboratory; that is, they can take a piece of non-metamorphic rock, subject it to various regimes of temperature and pressure, and see which characteristic minerals form at which temperatures. So in shale, for example, we see a sequence (as temperature increases) from unaltered shale to rocks containing chlorite; then biotite; then garnet; then staurolite; then kyanite; and finally silmanite. The image below show some particularly large garnets embedded in metamorphic rock. We see this same sequence arranged geologically as we approach the center of an area of metamorphism: from unaltered shale through the "chlorite zone", the "biotite zone", the "garnet zone" the "staurolite zone", the "kyanite zone", and the "silmanite zone". We may not get all the way up to silmanite, that depends how intense the metamorphism was at the center of metamorphism.The sequence, of course, depends on the parent rock; the sequence given above is specific to shale, and we would see a different sequence if we were looking at (for example) mafic igneous rocks.MetasomatismIn the section above on chemical changes we dealt with the case where the parent rock reacts with itself as a result of heat or pressure acting on the rock. But in the case of contact metamorphism, we can also see metasomatism taking place: the parent rock mixes and/or reacts with the intrusive igneous rock and the hot fluids associated with its eruption. The picture above shows a close-up view of skarn, a very distinctive product of metasomatism. Dark gray intrusive igneous rock has mixed with white recrystalized limestone; index minerals have also been produced, such as garnet (reddish-brown) and epidote (green). Textural changesBesides chemical changes, rocks that undergo metamorphosis suffer textural changes, such a recrystallization, foliation, and lineation.In recrystallization, the original texture of the rock is lost as the minerals, under the effect of high temperatures, reform as a collection of interlocking crystals of similar size. The effect of this is seen most dramatically in sedimentary rocks.So, for example, quartz sandstone loses its sedimentary structure of cemented grains to become quartzite, with a smooth texture consisting of interlocking crystals. As a result, except at very low grades of metamorphism, the bedding of the rock will be destroyed, as will any fossils that the rock contains. Similar textural changes produce marble from limestone, and hornfels from mudrock.When rocks are metamorphosed by pressure as well as heat, they undergo foliation, in which sheet silicates, if they are present in the rock, rearrange themselves so that the sheets are at right angles to the direction of pressure. The picture below shows a view of foliation under a microscope. Lineation is a similar phenomenon affecting silicates with the structure of a chain or double chain; the direction of the chain ends up, again, at right angles to the direction of pressure.Not every metamorphic rock will display foliation or lineation: some rocks simply don't contain any sheet or chain silicates: an example would be limestone, which metamorphoses to marble. Also some metamorphic rocks are formed by heat without any significant pressure, as is usually the case with contact metamorphism; so, for example, mudrocks, which will form foliated shale or schist under pressure, will produce non-foliated hornfels without pressure.Foliation comes in several varieties:Slatey foliation is caused by the alignment of sheet silicates such as clay minerals and chlorite (which is produced by chemical changes to clay minerals). It results in rocks which cleave easily into thin layers.Schistosity is caused by sheet silicates such as biotite and muscovite. Not only do they align, but they tend to separate out from the non-sheet silicates such as quartz, producing a rock that breaks easily into thicker leaves than those found in slatey rocks.Finally, we come to gneiss. At high grades of metamorphism, sheet silicates tend to break down, and dark-colored chain silicates such as hornblende and pyroxene begin to appear. These are separated out into dark bands, again at right-angles to the direction of pressure, giving gneiss a distinctive streaky appearance, as shown in the photograph. How do we know?First of all, as we have observed, we can reproduce metamorphic processes in the laboratory. Marble, for example, is what we get if we heat limestone; quartzite is what we get if we heat quartz sandstone; schist is what we get if we heat mudrock and apply pressure. It would seem downright perverse to maintain that metamorphic rocks should have been produced by other processes not as yet discovered. (Note that the textures of metamorphic rocks excludes the possibility that they are sedimentary rocks, and their chemical composition usually excludes the possibility that they are igneous rocks.)We can also look at the patterns we find in the rocks. I shall give some examples of the kind of predictions we can make from the theory of metamorphism; the reader will doubtless be able to think of other examples.To take a simple example: when we look at an aureole of marble, we should expect to find it embedded in an outer ring of limestone, and not of (for example) sandstone, which would go with an aureole of quartzite.Then again, according to our notion that metamorphic rocks are indeed produced by metamorphosis, we should not (and we do not) find, looking horizontally at sequences of rocks, alternating bands of unaltered rocks and high grade metamorphic rocks. Instead, as we have noted above, we find concentric zones of rocks with high-grade metamorphic rocks at the center of metamorphism, progressing to lower and lower grades of metamorphism until we reach unaltered rocks.If we find a foliated rock like schist, then according to our interpretation of schist as produced by temperature and pressure, we should find other evidence of the pressure; we should expect to find the beds of rock buckled and deformed. And this is indeed what we see.If, on the other hand, we find hornfels, which, laboratory experiments show, requires temperature without significant pressure (otherwise it would be foliated) then we expect to find (and do) that it forms an aureole around igneous rock, with progressively lower grade metamorphism in concentric zones around the igneous rock.Other patterns are discernible: for example, we would not expect to find schist overlying limestone, because the events that created the schist would also have turned the limestone into marble.In summary, the chemical composition and texture of the rocks that we have classed as metamorphic, together with their arrangement and relation to other rocks in the geological record, is just what we should expect to see if they are indeed produced by metamorphosis.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.Edited by Adminnemooseus, : Add "Metamorphic rocks" subtitle.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.

Mechanical weathering and erosionIntroductionGeologists make a distinction between weathering and erosion: weathering breaks rocks but leaves them in place; whereas the processes of erosion are capable both of breaking rocks and transporting the broken fragments (clasts). Most, though not all texts will make this distinction. Those that do will still use "erosion" as a catch-all term: that is, when you see a geologist saying that a rock has been "eroded" s/he does not mean to imply that it has not also been weathered.In this article we shall present a brief overview of erosion and of mechanical weathering. We shall be brief because the erosional processes involved will be discussed at length in subsequent articles, so there is no need to do more than sketch out the topic and its vocabulary. Mechanical weatheringMechanical weathering is sometimes referred to as physical weathering. In both cases, the purpose of this nomenclature is to distinguish it from the processes collectively known as chemical weathering, in which chemical action breaks down the rocks. Chemical weathering will be covered in the next article.Mechanical weathering involves mechanical processes that break up a rock: for example, ice freezing and expanding in cracks in the rock; tree roots growing in similar cracks; expansion and contraction of rock in areas with high daytime and low nighttime temperatures; cracking of rocks in forest fires, and so forth.Mechanical weathering is probably the least important process we shall mention in this text, in that the history of the Earth, and the resulting geological record, would probably have been very similar if there was no such thing as mechanical weathering. Mechanical erosionThe main agents of mechanical erosion are: gravity; aeolian processes (i.e. those caused by the wind); ice in the form of glaciers; and water in the form of rivers, waves, turbidity currents, and runoff caused by rainfall.The reader will be familiar with most of the processes described, but we should provide a brief introduction to the concept of turbidity currents. A current of water that is turbid (that is, which contains a lot of sediment) is denser than clear water, and will flow along the bottom of a lake or the ocean, often over large distances and at high speeds, before failing and dispersing its load; such current occur when a turbid river discharges into the clear waters of a lake, or they can be initiated by a mudslide on a continental shelf. A dust storm may be considered the aeolian equivalent of a turbidity current. Modes of erosionAbrasion of rocks is caused by the sediments carried by wind and water: waves, for example, can hurl their seaload of sand and shingle against a cliff; sandstorms can literally sand-blast rocks; the sand and silt carried by rivers or turbidity currents have the same effect.Attrition is the effect these same forces have on the sediments themselves, breaking them into smaller fragments or rounding the clasts into smooth pebbles or rounded grains of sand. The efficiency of this process can be observed anywhere you can find beach glass, which originates as sharp-edged shards; the process of tumble-polishing semi-precious stones artificially emulates this process and will render most pebbles well-rounded in a matter of days.The simple mechanical force of water or ice can break off chunks of rock, as when glaciers quarry rocks from the surfaces over they move, or when the pounding of waves hammers against a cliff.Gravity can break off the overhang of a cliff undercut by abrasion and wave pounding, when the rock at the top of the cliff is unable to bear the mechanical strain. It is also instrumental in causing such things as rockslides and mudflows; such downhill motion is known collectively as mass wasting. Transport of sedimentErosional processes, as we have said, are defined by their ability to transport sediment as well as to create the clasts of which it is composed.Currents of wind or water can transport sediments in three ways: in suspension, where light particles are carried along in the current above the ground, sea bed, or river bed; by saltation, where particles too heavy to be carried in suspension are bounced along the ground (or river-bed, or whatever); and creep, where particles are rolled along the ground. The size of the particles susceptible to these processes will depend, of course, on the velocity of the current.By contrast, the size of the clasts that can be carried by glaciers is under no such limitation. One of the characteristic results of glacial action is the transport of huge boulders, up to the size of a house, known as erratic boulders. Gravity, too, is obviously under no such limitation; it is possible for entire layers of rock to slide down a hillside. How do we know?In this particular case, asking "how do we know?" seems almost superfluous, for the processes involved are neither hidden nor subtle. We can observe a sandstorm; we can see how the head of a waterfall shifts year on year, or how a river shifting its course scours out a new bed; we can see how cliffs crumble and the effects of landslides. The fact that glaciers carry boulders is evident, and the distance they travel each year can be measured; as can the quantity of sediment discharged at the mouth of a river.A more interesting question is, how do we know that these processes have happened when the agency that caused them is no longer present? How, for example, do we identify the courses of glaciers long since melted, or of rivers that have dried up or shifted their beds? These are questions that we shall review in subsequent articles.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.

BTW, is anyone still reading this? You could indicate as much by "liking" this post. Pressie and I were expecting more feedback.Edited by Dr Adequate, : No reason given.

Chemical weatheringIntroductionChemical weathering is the breakdown of rocks into sediment by chemical processes.The reader who does not appreciate chemistry may skip the chemical formulas in this article and simply note the results of the reactions, as described in the summary section below. The results themselves cannot be skipped over: understanding chemical weathering is essential to answering such basic questions as: "Why is sand mostly made of quartz?" and "Where does clay come from?" Agents of chemical weatheringThe main agents of chemical weathering are: Water. Some minerals, such as rock salt, will dissolve readily in water; others such as pyroxene, will also do so, though at a much slower rate. Carbonic acid. Rainwater and groundwater are not pure water; some of the molecules of water react with the carbon dioxide in the atmosphere (in the case of rainwater) or produced by bacteria and plant roots (in the case of groundwater) producing carbonic acid, as follows: Oxygen. This is a highly reactive chemical, and the only reason that there's so much of it about in the atmosphere is its constant production by biological processes and a shortage of things that it hasn't reacted with already. Chemical weathering of common mineralsIn this section we shall look at how some common minerals are affected by chemical weathering. We have arranged the list more or less in order from the minerals most susceptible to chemical weathering to the most resistant. Halite. Salt, of course, dissolves in water. This is why you are unlikely to see rock salt on the surface except in desert environments. Gypsum. This, like halite, is soluble in water; similar remarks apply to it. Calcite. This, you will recall, is the mineral forming limestone and its metamorphic counterpart, marble. It can just dissolve in water; it also reacts with carbonic acid as follows: The ease with which limestone dissolves (relative, at any rate, to other minerals) produces the distinct topography of a region built on limestone rocks, with underground caves full of stalactites; sinkholes where the land has subsided; streams disappearing into the ground or rising out of it as springs. This is known as karst topography.This is also the reason why a marble tombstone, though handsome in appearance, is not a good long-term investment. Silicate minerals. Of silicate minerals in general we may observe that the more mafic minerals with higher melting points (those to be found to the right of our diagram in the article on igneous rocks) are more susceptible to chemical weathering than felsic low-temperature minerals. This will be reflected in the order in which we list them below. Olivine. This mafic mineral has the formula (Mg,Fe)₂SiO₄. Recall that the (Mg,Fe) in the formula means that it is a solid solution in which varying quantities of magnesium or iron can play the same chemical role. It reacts with carbonic acid as follows: As with limestone, the constituent parts of the mineral are now entirely dissolved in water, leaving no residual mineral.On the other hand, iron olivine can react with water and atmospheric oxygen like this: And the hematite can further react with water as follows: Hematite and goethite are both very insoluble in water: they remain as residual minerals. It is these iron oxides that give many soils their reddish or yellowish color. Calcium feldspar. Chemical weathering turns calcium feldspar into clay (a residual mineral, since clay does not dissolve in water) and leached ions, which dissolve in water and which usually will end up contributing to the dissolved mineral content of seawater. The reaction is as follows: Pyroxene. The typical rock-forming pyroxenes have the formula (Mg,Fe)SO₃. This can react with carbonic acid as follows: Again, as with olivine, the constituent parts of the mineral are dissolved in water. However, as with olivine, iron pyroxene can react with oxygen and water to produce the residual mineral hematite: Again, the hematite can turn to goethite. Mica and amphibole, minor constituents of felsic and intermediate rocks, undergo rather more complicated reactions. (Details may be found here for biotite, a mica, and here for amphibole.)In summary, the residual minerals produced are clay minerals; iron oxides; and, in the case of biotite, the mineral gibbsite (Al(OH)₃), which is usually found in association with clay. Potassium and sodium feldspars produce residual clay minerals. Here, for example, is the reaction by which potassium feldspar produces kaolinite (plus various dissolved substances): As sodium is chemically almost indistinguishable from potassium (since they both have the same valance) potassium feldspar reacts in a similar way. Quartz is the most stable of the silicate minerals. This is why quartz sand is so common as a sediment; when all the other constituents of igneous rocks have either dissolved or been converted to clay, grains of quartz will remain. This is why there is no such thing as a large grain of sand: the maximum size of such grains is limited by the size of the quartz crystals that form in granite and similar rocks.Note, however, that quartz sandstone is vulnerable to chemical weathering, because although the grains of quartz themselves are resistant, the minerals cementing them together may not be. Clay minerals are very resistant to chemical weathering, because they are, as we have seen, a product of chemical weathering, and, like all minerals, they are stable under the conditions under which they were formed. Iron oxides. These, as we have seen, are a product of chemical weathering of iron-bearing forms of such mafic minerals as olivine and pyroxene. We have noted that hematite can be converted to goethite by oxidation: Once iron oxide has formed, there is very little that can happen to it except conversion to another sort of iron oxide; these are regarded as the most stable of all common classes of minerals. SummaryThe residual minerals left after chemical weathering has done its work are quartz, clay minerals, a scattering of ion oxide, and sometimes a little gibbsite. The other constituents of minerals are dissolved; their usual fate is to be carried by rivers to the sea, where they contribute to the dissolved mineral content of seawater.We may note that most land-formed sediments are in fact quartz sand, clay, or a mixture of the two. This demonstrates the predominance of chemical weathering over mechanical weathering and erosion. If sand or mud were produced simply by mechanical crushing of granite, then they would be 60% feldspar; but they are not. When we find any appreciable amount of feldspar in sand (as in arkose sandstone), we may infer that there has been a higher than usual ratio of mechanical to chemical processes. How do we know?The processes of chemical weathering are sufficiently slow that it is reasonable to wonder how we know that they take place at all.In the first place, we know that they ought to take place. According to the theory of chemistry, in the chemical equations given above, under the conditions in which chemical weathering takes place, the situation described by the right-hand side of the equation are more thermodynamically stable than the situations described on the left hand side; so the reactions described should happen.We can speed up reactions which, in nature, involve weak, highly-diluted carbonic acid, and instead use something stronger, such as hydrochloric acid (HCl). In principle the only difference this should make (besides, obviously, the substitution of chloride ions for bicarbonate ions) is that as hydrochloric acid gives up its hydrogen ions more readily, the reaction will go faster. So, for example, we can use hydrochloric acid to convert potassium feldspar to clay as follows: And this happens fast enough for us to observe it happening. Alternatively, to experiment under more natural conditions, you can bury samples of minerals in a nice moist acidic soil, where chemical weathering occurs fastest, leave them for a few years, and see what's happened to them; they will not weather completely, but if the process involves conversion of one mineral to another, rather than just dissolution, then the chemical changes are observable on the surface of the mineral.Without any of these artifices, we can find naturally occurring rocks which appear to be in the process of weathering: for example, the exterior of such mafic rocks as basalt can often be seen to be rusty, as a result of the iron pyroxene being converted to iron oxides. In the same way, granite boulders can be found with a weathering rind, where on the outside of the boulder the feldspar on the outside has been partly converted to clay, while the feldspar on the inside is still relatively intact. In tropical soils, it is possible to find granite boulders in which the feldspar has become so "rotten" with clay that it is literally possible to kick the boulder to pieces. Under a microscope, the feldspar crystals will appear corroded and pitted. Or if a road cut or railroad cut goes through a hill (an event which always delights geologists) then since the rocks near the top are more weathered, we can take a series of samples going up through the vertical section from unaltered rock to completely weathered rock (saprolite) as in this study of the weathering of biotite.We can observe the effect of weathering on old tombstones or on dressed stones used for building; as you would expect, this is most noticeable in the case of limestone or marble. In such cases we can usually be quite certain that the stones in question have suffered little from merely physical processes such as abrasion by sandstorms, tidal action, transport in rivers, and so forth.Finally, we may note that the processes we have described explain the nature of sediment: they explain why so much of it is quartz sand and clay; they explain, as we have seen, the size of sand grains; they explain the origin of the minerals that cement together the grains in sandstone; and they explain the origins of the minerals found in sea water as dissolved ions.In summary: these processes ought to happen; we can simulate them happening; and what we see in nature is just what we should see if they did happen.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.

Quite so. I have in fact gone into much more detail than is usual in an introductory textbook. I feel that this is in line with my intention to explain "how do we know". I could have been much shorter, at the cost of making it seem that primary rocks turned into clay and quartz sand by magic --- but I think it's important to explain why this is so. Here Pressie is being needlessly modest: I know, and the readers of this thread should know, that he worked very hard behind the scenes to help check the facts for this particular article.

I did briefly mention weathering rinds in the "how do we know" section.Dating is a topic that I'll come to much later, in fact practically at the end.

GlaciersIntroductionA glacier is a mass of ice moving on the surface of the Earth. In this article, we shall discuss how glaciers form and move; we shall discuss the geological features associated with glaciers; and we shall show how recognizing these features can allow us to tell where glaciers have been in past ages of the Earth's history. Formation and motion of glaciersA glacier forms at an accumulation point, that is, a place where more snow accumulates than melts. This snow then piles up and compacts under its own weight to form ice.Even if this happened on a perfectly level surface, as the ice mounted up it would eventually start to squidge outwards under the pressure of its own weight; and often glaciers will form on mountaintops, where gravity is also a factor. Under the effects of pressure and/or gravity, the ice will flow. A glacier flows in two ways: by sliding along its base, and by "plastic flow" of the molecules of ice within the glacier.You may recall from science class that ice tends to melt under pressure; this means that the base of a glacier is often lubricated by water.The overall speed of a glacier can be measured by simple methods: hammer a stake into a glacier, wait a while, come back, and see how far it's moved. The speeds so measured range from centimeters to meters per day, depending on the glacier.One significant different between the flow of ice and the flow of water is this: a river is pulled downwards by gravity. This happens to glaciers too, when flowing downhill; but glaciers are also pushed by the pressure behind them: as a result, glaciers can and do flow uphill.Once in motion, the ice in the glacier will keep flowing until it reaches a point where the ice ablates: either it reaches the sea, breaking up into bergs, or it reaches a zone where the climate is warm enough to melt the advancing glacier. In the latter case, the end of such a glacier represents an equilibrium state at which the rate of melting is just sufficient to balance the rate of flow of the glacier.Now, while this equilibrium is maintained, the glacier as a whole will stay still. The ice in the glacier moves, starting off at the accretion point and ending up at the ablation point, but the glacier as a whole stays in one place: it is like a conveyor belt of ice moving from accumulation to ablation.The length of the glacier will change with the climate: for example, if the climate gets warmer around the ablation end of the glacier, then the glacier won't be able to progress as far before reaching a zone in which the rate of melting equals the rate of flow, so the glacier will retreat (note that the ice in the glacier will still be moving forward while this is going on). Conversely, of course, a drop in temperature will let the glacier get further from the accumulation point. Changes at the accumulation site of the glacier will also affect its length: the more it snows at the accumulation point, the greater the volume of flowing ice, and the further it will get before it melts. It follows that global cooling will cause glaciers to extend further from accumulation points, and global warming will see them extend less far, or vanish entirely if they temperature rises so much that the snow melts at the former point of accumulation.Types of glacierA glacier originating where the snow accumulates on a mountaintop and flows down the mountain is known as a valley glacier or an alpine glacier: the terms are synonymous.The larger glaciers, such as can be found today covering the surface of Greenland and Antarctica, are known as continental glaciers, or sheet glaciers or ice sheets. Again, these are synonymous and the diversity of terms does not indicate that some sort of distinction is being drawn. Whereas valley glaciers flow downwards from the point of accumulation, the sheet glaciers of Greenland and Antarctica flow outwards in all directions from the point of accumulation. Erosion associated with glaciersA glacier moving over a landscape will sweep up topsoil and loose rocks as it goes, transporting them towards the ablation end of the glacier and revealing the bedrock beneath. This bedrock will be polished by the passage of the ice over it; it will also be grooved and scored in the direction of travel by the rocks contained in the glacier: such groves are known as striations or striae, and the rocks are said to be striated. The picture below shows an example of striation: The passage of a glacier will produce debris of all sizes, from enormous chunks of rock plucked from the bedrock to very fine rock flour produced by the grinding action of the glacier. A valley glacier will also carry along any rock fragments that fall from the valley walls. Some glaciers, known, reasonably enough as "rock glaciers" consist mostly of rocks cemented together by ice.One erosional feature commonly associated with glaciers is the roche moutone, caused when a glacier slides over a hummock of bedrock. As the glacier slides up the hill, it polishes and striates it; flowing down the other side, it plucks fragments from the rock, leaving a steeper and more ragged face in its wake.A valley glacier will created a bowl known as a cirque at the accumulation point, with about a quarter of the rim of the bowl missing in the direction in which the glacier eaves the cirque. A mountain eroded by glaciers will have a rugged, jagged topography, with knife-edge ridges where two cirques or glacial valleys adjoin. The valleys carved out by a valley glacier will have a characteristic U-shaped cross-section quite different from that produced by a river, which produces V-shaped valleys. Deposition associated with glaciersThe rock fragments transported and deposited by glaciers are known as till. It is unrounded and unsorted by size. This may not sound very remarkable, until we reflect that the action of wind or water cannot produce sediments like this: such an unsorted jumble of shards is characteristic of glacial action.In a valley glacier, there will be concentrations of till along the border of the glacier, where it has been plucked or ground from the valley sides. Such an accumulation of till is known as a lateral moraine. Where two valley glaciers meet, lateral moraines will merge into medial moraines in the middle of the larger glacier so formed, as shown in the aerial photograph below. Ground moraine is till deposited over a wide area either when the till on the underside of the glacier lodges against something, or when a glacier retreats in response to climatic changes. Ground moraine is often found in small hills, shaped somewhat like the back of a spoon, known as drumlins. No-one is really sure how these form, but the fact that they are composed of till and found in conjunction with other signs of glacial action confirm that they are glacial in origin; also, they are invariably found oriented with their long axis in the direction of the flow of the glacier (as determined by study of striations, roches moutones, etc). At the ablation end of a glacier, the sediments transported by the glacier will be dumped to form an end moraine, resulting in a ridge of till in the same convex shape as the lobes typically found at the end of a glacier.Beyone the ablation zone, where the glacier melts, the water from it will be carried away, typically in a braided stream (a term which we shall explain in depth in our article on rivers). This will carry with it the lighter sediments, known as outwash, which will be deposited in front of the glacier as an outwash plain.One interesting sedimentary feature can be seen in lakes fed by glacial waters (proglacial lakes). In summer, relatively coarse outwash of sand and gravel will be deposited in these lakes; in winter, when the lake freezes over, the calm conditions below the ice allow fine particles of clay and of organic material to settle. The result is the formation of couplets of sedimentary material, one fine, one relatively coarse, repeated over and over, each couplet being known as a varve. Because the deposition of varves is an annual event, the study of varves is of interest in dating.Kettles are another feature we associate with glaciers. When a glacier retreats, we often observe that it leaves a large block of ice behind it. The outwash sediment from the retreating glacier will then build up around the orphaned block of ice. When it has melted, which may take many years, the result is a depression in the outwash plain: this is a kettle. If it lies below the water table, it will fill up to produce a small lake with outwash banks. Former glaciers: how do we know?It is not difficult to detect the passage of a glacier, even if it is no longer present: for if you take away the glacier, you are still left with the patterns of deposition and erosion that glaciers produce; and these are highly distinctive and cannot be produced by other mechanisms. We may note that we can see glaciers disappearing today: for example, Glacier National Park in Montana has at the time of writing only 26 named glaciers, down from 150 in the year 1850; so our statements about what evidence former glaciers leave behind are by no means hypothetical, but rather are based on direct observation.With all former glaciers, whether valley glaciers or continental glaciers, we see characteristic patterns of erosion: we see such things as bedrock polishing, striation, roches moutones, and so forth. We also see till. As we have remarked, till is a very distinctive sort of sediment, which cannot be produced by the action of wind and water, as shown by its unsorted, unrounded nature. The arrangement of till can also be quite distinctive; a pile of till in a crescent-shaped end moraine admits of no other explanation except that a glacier deposited it there.The picture below shows a scene which any geologist would instantly recognize as glacial in origin: we are looking at the shallow (downstream) side of a striated roche moutone liberally scattered with glacial till. Erratic boulders, when we find them, present another blatant clue. An erratic boulder is one which, in terms of the rock of which it is composed, has nothing in common with the geology of its surroundings, and which must have been transported to its present location over some distance; in some cases, hundreds of kilometers. Ice, as we can observe, can transport such enormous rocks; water and wind do not.We are therefore left with some unambiguous signs of former glaciers. When we see smoothed bedrock marked with striations leading to a semi-circular moraine of unrounded and unsorted rocks, many of them way out of their original geological context, beyond which is what looks suspiciously like an outwash plain, there is really no other conclusion that we can draw, except that we are looking at where a glacier once flowed and terminated.Besides the erosional and depositional features already mentioned, glaciers leave behind some highly distinctive landforms. Consider the for example the photograph below, taken in North Cascades National Park. There is no longer a glacier present, and, indeed, a lake has formed at the former accumulation point. (Such a lake is known as a tarn.) Despite the absence of any actual ice, the reader should have no trouble in recognizing the landforms associated with a valley glacier, which are as clear and distinctive as an elephant's footprint. Here, surely, is the great bowl of a cirque, and just where we would expect to find it, near the peak where the temperatures are lowest; in the foreground, where the rim of the cirque is open, we see a valley with the distinctive U-shaped cross-section of a glacial valley, where the glacier once exited the cirque.Sheet glaciers do not carve out such distinctive forms, but they do leave large-scale clues behind. They often expose large expanses of bedrock. They also erase the drainage systems that were present in the landscape before their arrival, so that after they retreat, the landscape is poorly drained: such features are a sign to geologists to look for other indications of glaciation.We should mention a couple more signs of glaciation. The first of these is isostatic rebound. Ice is heavy: for reasons that we shall go into more thoroughly in our articles on tectonics, the weight of a continental glacier should press the crust of the Earth down into the mantle, and, when the glacier is gone, the Earth's crust should slowly "bounce" back up. This happens rapidly enough as to leave its marks over mere centuries: so, in parts of Scandinavia, we can see former harbors now standing uselessly distant from the sea. Today, the rate of rebound in post-glacial areas is measured directly by a GPS monitoring system called BIFROST: the maximum rate of rebound is about 1 cm/year. This on its own would not prove the former presence of glaciers, but in combination with the less ambiguous signs of erosion and deposition, the phenomenon of rebound does confirm the hypothesis of glaciation.There is one more prediction that we can make and confirm. We can use geological dating methods, of which more will be said in later articles, to establish the times at which sheet glaciers covered northern North America and Eurasia. Now, if we are correct in attributing these striations, moraines, and so forth to glacial action, then we ought to find that at the same time, we have other evidence of a colder climate, such as flora and fauna adapted to colder climates; lower sea levels caused by water being locked up in continental glaciers; temperature-dependent changes in the composition of shells; and so forth. And this is exactly what we do find, providing an independent confirmation of an ice age.We shall deal more fully with these topics in later articles on geological dating methods and on paleoclimatology; for now we shall simply note that these techniques can be used, and that they confirm what we can learn from studying landforms, sediments, and erosional features.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.

I had the impression that it was much harder to do paleoclimatology in more distant geological periods (apart from saying: well, it must have been cold, look at all the evidence of glaciation) so when I was talking about the tie-in between dating and paleoclimatology, I restricted my remarks to the most recent Ice Age.I shall be mentioning glaciation on other continents when I discuss continental drift.

This comment doesn't relate to my post. Of course we have evidence for a recent Ice Age, I've made that clear.Let's start again.My point at the end of the article was that we can look at things which are (claimed to be) evidence of glaciation and find some support for this claim by finding independent evidence that it really was very cold when this (supposed) glaciation took place. This is admittedly a minor argument when compared with the fact that it looks exactly like glaciation took place, but I thought it was worth mentioning as another converging line of evidence.When I raised this point, I specifically mentioned North America and Eurasia, i.e. I was talking about the most recent Ice Age(s).I know that there were other Ice Ages affecting other continents, but I didn't mention them.Why not? Because I have the impression (I may be wrong) that when we're looking at earlier geological periods, it is for technical reasons harder to find evidence of the cold independent of glaciation --- the best evidence that it was cold in (for example) East Africa during the Permian would be the evidence of glaciers. But I couldn't in this context adduce that as evidence that it was cold without involving myself in circular reasoning.

Yes, you're right, that sentence is a mess. I shall amend it, thank you.

The ice at the top may have the lowest rate of flow, but doesn't it also have the highest velocity?---In any case, I've just removed what I said for now.On reflection, I probably shouldn't have said it in the first place. I want to keep this series of articles as compact as possible, so I really shouldn't be talking about the physics or chemistry of any process unless it relates to the evidence that that process has taken place. Now whichever way the rod bends, it doesn't help us to determine whether some area of land was once subject to glaciation, so I have no idea why I brought it up.Edited by Adminnemooseus, : Put the "code" code in the quote box, to preserve raw formatting.

But isn't it the total movement at a given depth that determines the displacement of the rod at that depth? I thought I'd made that quite clear.P.S: I did. It says: "if the climate gets warmer around the ablation end of the glacier, then the glacier won't be able to progress as far before reaching a zone in which the rate of melting equals the rate of flow, so the glacier will retreat (note that the ice in the glacier will still be moving forward while this is going on)".Edited by Dr Adequate, : No reason given.

DesertsIntroductionIn this article we shall discuss the forces that shape deserts, and discuss how geologists can use this knowledge to identify rocks that are the lithified remains of former deserts. We shall use some terms that have been introduced in previous articles on Sedimentary Rocks and Mechanical Weathering and Erosion; the reader who has not already read these articles may find it useful to go back and do so. Causes of desertsThe immediate cause of a desert is, by definition, lack of rainfall. This itself can have a number of causes, which are not mutually exclusive: an area can be a desert for more than one reason.* High-pressure deserts. In zones of high atmospheric pressure, the ability of air to contain moisture is increased, resulting in little rainfall. Examples include the Sahara, Arabian, Thar, and Kalahari deserts, and the desert regions within the Arctic and Antarctic circles.* Mid-continent deserts. Areas in the middle of a continent can receive little rainfall simply because rain, originating from evaporation of seawater, will tend to fall before it can reach the middle of a large continent. Modern examples are the Turkmenistan, Gobi, and Great Australian deserts (the Great Australian Desert is also in a region of high pressure).* Rain-shadow deserts. Rain will tend to break over mountains, so the presence of mountains can prevent rain from the sea from coming inland. Examples of rain-shadow deserts include the Mojave desert in the rain-shadow of the Sierra Nevada, the Patagonian desert in the rain-shadow of the Andes, and the Iranian desert in the rain-shadow of the Zagros mountains.* Upwelling deserts. Finally, a desert may be by the coast and not in the rain-shadow of any mountains, but be adjacent to where a cold current of water rises to the ocean surface, reducing evaporation. Examples include the Atacama desert, the Western Sahara, and the Namib desert; these are all also in high-pressure zones. Deserts and waterAlthough rainfall is rare in deserts, its effects are important in understanding the geology and ecology of deserts. It is characteristic of deserts that they exhibit internal drainage: that is, the water input to the desert in the form of rain or runoff from mountains does not flow out of the desert in the form of a river, but rather evaporates within the desert: this produces some highly typical depositional features. Alluvial fans and bajadas. Deserts are often found associated with mountains: indeed, as we have seen, in some cases the mountains are the indirect cause of the desert. Runoff from rain on the mountains erodes them and transports the sediment downslope: when the sediment-rich water loses energy at the foot of the mountains, it forms a braided stream which rapidly deposits all but the lightest sediment in an alluvial fan (alluvial because it is formed by running water, and "fan" because this is the shape in which it is deposited). Two or more adjacent alluvial fans can merge to form what is known as a bajada.The photograph below shows an alluvial fan in Death Valley. The braided fanlike patterns made by the streams that deposited it are clearly visible. Playa lakes. These are temporary lakes, which, being shallow and fed only by intermittent rainfall, dry up in the arid climate of the desert, leaving behind a flat, heat-cracked bed of clay, known as a playa.If the water contains any dissolved minerals these will be deposited on top of the clay layer as the water evaporates, leaving an evaporite bed of such minerals as halite, calcite, gypsum, borax, or trona, depending on the rocks that were their source. Repetition of this process builds up alternating layers of detrital and evaporite sediments. The photograph below shows the surface of a playa. Oases. Oases are small lakes in the desert. These have a number of causes: they can be fed by springs; they can be deflation lakes, where erosion has caused a hollow the bottom of which lies below the water table; or they can be intermittent lakes filled by occasional rainfall and runoff. Deserts and windDeserts are the only places where wind is a major factor in erosion and deposition: soil bound by moisture and vegetation is harder for the wind to budge than loose, dry particles.Wind has a number of effects. First, it can remove the sand and other light particles from the surface. leaving bare rock or the stoney mosaic surface known as a desert pavement as shown in the photograph below. Second, it can erode rocks by abrading them with the particles it carries: the photograph below shows the effects of wind erosion in a desert environment. Thirdly, it can pile sand up into dunes, giving us our stereotypical image of a desert. A sandy desert is known as an erg. The dynamics and shapes of dunes vary depending on how variable the regional winds are. The animation below may however be considered a cross-section of a typical dune, lying with its long axis at right-angles to the direction of the wind. The wind transports the sand up the windward (stoss) side of the dune, to build up just over the crest of the downwind (lee) side of the dune. It accumulates until it reaches a critical angle at which it must avalanche, at which point the accumulated sand slides down the lee face of the dune to form a grainflow lamina. The result of the transport of grains from stoss to lee is that the dune will move downwind. However, it will not always erode completely on the stoss side, instead leaving behind it a set of cross-beds composed of the bottom of the grainflow laminae, as we have shown in our diagram. Then the next sand dune to pass that way can deposit another discontinuous set of cross-beds on top of that set (and on top of any sand that may have been deposited by the wind between the passing of the two dunes). By this process, set after set of cross-beds build up.The cross-bedding is informative in a number of ways. First, it shows us that the sand was deposited in dunes: we shall discuss in the next section how we can determine that these dunes were caused by currents of wind and not water.Secondly, you will note from our diagram that the gradient of a grainflow lamina tends to become shallower at the bottom: this feature can be seen in the cross-beds in the photograph. Now, tectonic events are quite capable of turning rocks sideways or upside-down. The curves of the cross-beds are one of a number of features, known as way-up structures, that allow geologists to determine which way was originally up: the curves are concave in the up direction, convex in the down direction. In the case shown in the photograph, we can see that the sandstone is still the right way up.Thirdly, the crossbeds show us the prevalent direction of the winds, as we shall discuss later on in the articles on paleoclimatology. Lithified deserts: how do we know?There are a number of features that we can use to identify sedimentary rocks that are the lithified remains of ancient deserts. In this section we shall review some of them.Desert sand is quartz, and the resulting sandstone will be quartz sandstone. This is not a definitive criterion for recognizing aeolian sandstone, since quartz sandstone can be formed under other conditions. We may, however, safely say that arkose sandstone or graywacke is not aeolian sandstone.Desert sands tend to be well-rounded, as a result of long abrasion, and well-sorted by size, and this is what we find in the grains of aeolian sandstone.Much, though not all, desert sand is colored, the grains being stained on the outside with iron oxides such as haematite, goethite and limonite, giving it a range of colors through red and orange to yellow. However, some desert sands are not stained in this way, so we cannot definitively say that sandstone lacking this feature is not desert sand.Cross-bedded features reveal transport by a current of wind or water. We can distinguish between the two cases by various features: for example, wind-formed dunes result in much bigger sets of cross-beds; also, aeolian sandstone uniquely exhibits pinstripe laminae, very narrow stripes only a few grains thick consisting of finer grains than are found in grainfall laminae.Playa deposits are extremely distinctive, and leave geologists no doubt that what they are looking at was once desert: they can only be formed by repeated episodes of deposition in and evaporation of a playa lake.Finally, we can look at the fossils, which, in a desert should be of land plants and animals, and of freshwater organisms in those rare places where the geology indicates a former oasis. No marine fossils should be found. Case study: the Navajo sandstoneAs an example, consider the Navajo Sandstone of the southwestern United States, a formation which extends over 400,000 square kilometers of northern Arizona, northwest Colorado, Nevada, and Utah, and which is 600 meters thick at its thickest point. You can read about it here and here.* The sand, which is cemented together by calcite, silica, and haematite, is very pure quartz sandstone without clays or feldspars.* The sandgrains are well-sorted, consisting mainly of medium-sized grains; and well-rounded. The photograph below shows a microscopic view of grains in the Navajo sandstone that have been crumbled gently by hand from the surrounding matrix. On the left are medium-sized grains from a grainflow lamina; on the right are predominantly finer grains from a pinstripe lamina. * In most of the Navajo sandstone, the grains are richly stained with iron oxides; this again can be seen in the photograph above.* The sandstone shows large cross-bedded features, of a type only observed to form in desert winds. The photograph showing crossbedded sandstone in the section on deserts and winds, above, is from the Navajo formation.* The sandstone exhibits pinstripe laminae, uniquely characteristic of aeolian sanddunes.* The sandstone contains layers of evaporites characteristic of playa lakes.* The sandstone contains no marine fossils, but instead contains fossils of land reptiles including Protosuchus, Ammosaurus, and Segisaurus. It contains fossilized footprints of land animals. It contains fossilized land plants including conifers, horsetails and ferns; root casts and stumps of trees can be found.* The land plants and animals and their footprints are found in association with rare thin interdune deposits consisting of lenses of limestone containing freshwater shells, indicating the presence of temporary freshwater oases (temporary, that is, in terms of a geological timescale); this is discussed here.In short, the Navajo sandstone was once a large sandy desert.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.

Volcanic AshIntroductionVolcanic ash is the name given to the fine particles of igneous rock produced by a volcano. As such, the name "volcanic ash" is a misnomer, since "ash" means a residue after incineration: the name is left over from times when people thought that volcanic ash really was ash.In this article we shall discuss the origin, transport, and lithification of volcanic ash, and as usual we shall consider how we can recognize volcanic ash, and its lithified counterpart (tuff) in the geological record. It will be helpful for the reader to have already read the main article on igneous rocks before reading this article. Origin and dispersal of volcanic ashA volcanic eruption can throw off a fine spray of lava into the atmosphere as well as or instead of producing a cohesive lava flow. Because it originates as a fine spray, the droplets of lava often cool too rapidly to form a crystal structure; instead, they form an amorphous solid (a glass), although sometimes cooling may be slow enough for minerals to form. Sometimes the ejecta will also include particles of material torn by the eruption from the solid rock forming the volcano. Collectively, these fine airborne particles are the constituents of volcanic ash.These particles can be distributed in two different ways. If the volcano is sufficiently forceful, it can throw the volcanic ash high into the air, where it will be distributed by wind. Particles transported in the way can travel great distances: for example, ash from the eruption of Krakatoa was carried all the way to the island of Sumatra.Alternatively, the ejected material can be transported in a pyroclastic flow. In this case, the ash-filled air is denser than the surrounding air because of its load of ash, and for this reason mixes poorly with the surrounding air, and so can travel over great distances (200 kilometers or more) before finally dissipating. The current can travel particularly far if it finds its way into a pre-existing channel such as a river valley. TuffLithified volcanic ash is known as tuff. This comes in two varieties.Welded tuff has its origins when the beds of volcanic ash, when they form, are hot enough for the clasts to compact and weld together.Cemented tuff is produced by the same cementing mechanism that cements together the clasts of coarse-grained sedimentary rocks such as sandstone. Volcanic ash: the igneous sedimentVolcanic ash is, obviously, an igneous rock: it comes out of volcanoes, and you can hardly get more igneous than that. However, in many ways it behaves like a sediment. As geologists say, it's "igneous when it goes up, and sedimentary when it comes down."For one thing, volcanic ash consists of small clasts, just like other sediments; also, as we have seen, it can undergo cementation just like the process that cements together sand grains to make sandstone. The image below is a scanning electron microscope image of a single clast from the eruption of Mt. St. Helens. Another feature that makes volcanic ash sedimentary in nature is that it can form graded beds like a sediment, with smaller particles on top. If the ash is shot high into the air, then the larger particles, the mass of which is better able to overcome air resistance, will settle first. If, on the other hand, the ash is transported by pyroclastic flow, then again the ash will still grade upwards from coarse to fine: as the current disperses and fails, the size of the particle that it can transport decreases, and so at any particular point in its path the particles that it deposits will decrease in size. (However occasionally, depending on the behavior of the volcano, the flow can also increase in speed, leading to inverse grading.)Finally, ash beds resemble other sediments because they are a good place to find fossils. Inhaling volcanic ash tends to shred the lungs of animals, and if the ash bed is reasonably thick the bodies will undergo rapid burial, leaving them well-preserved.However, volcanic ash is also igneous, and for this reason displays a number of features that make it of special interest to geologists. A particular volcanic eruption is a single event at a particular point in time: when we see a horizontally continuous layer of volcanic ash, we know that it was all laid down at the same time, and whatever is immediately beneath it was a snapshot of the landscape as it was just before the eruption. Such a geological feature is known as an event horizon.This is by no means true of sedimentary rocks in general. Just because you have a continuous horizontal stratum of, for example, sandstone, this does not necessarily mean that the beds at one end were laid down at the same time as the beds at the other end, as we shall discuss in a later article.Furthermore, each volcanic eruption has its own chemical signature: the exact mix of chemicals in one volcanic eruption will be different from another. This means that even if a layer of volcanic ash is not exposed everywhere, so that we cannot directly observe that it is a single layer, we can still check whether two exposed areas of volcanic ash do or don't belong to the same eruptive event.Finally, unlike sedimentary rocks, volcanic ash and the rocks that form from it can be dated directly. Volcanic ash and tuff: how do we know?It takes no subtle chain of reasoning to know that volcanic ash or tuff found in the geological record is indeed volcanic ash or tuff. For when a volcano erupts, we can watch volcanic ash beds being formed: and the beds in the geological record look like that.This appearance is even more evident under the microscope. In the photograph below we show a photomicrograph of welded tuff, where the clasts have been squashed and welded together (the larger, white fragments are broken crystals included in the ash). Note the distinctive shape of the clasts: no-one looking at welded tuff under a microscope could mistake it for anything but welded tuff. Finally, we may note that since ash and tuff are igneous in origin they have a chemical composition that lies on the usual felsic-ultramafic spectrum of igneous rocks: so we may speak of "rhyolite ash" or "basaltic tuff". So chemically there can be no doubt that volcanic ash and tuff are igneous in origin: but the fact that they are clastic means that we can't confuse them for any other sort of igneous rock.These considerations mean that geologists can be confident of identifying volcanic ash and tuff when they find them in the geological record.Edited by Dr Adequate, : No reason given.Edited by Dr Adequate, : No reason given.