کانی های سیلیسی-Silica minerals

Silica minerals

Silica (SiO2) occurs naturally in at least nine different varieties (polymorphs), which include tridymite (high-, middle-, and low-temperature forms), cristobalite (high- and low-temperature forms), coesite, and stishovite, in addition to high (β) and low (α) quartz. These forms have distinctive crystallography, optical characteristics, physical properties, pressure-temperature stability ranges, and occurrences (see table).

 

 

Fig. 1  Portion of an idealized sheet of tetrahedrally coordinated silicon atoms similar to that found in tridymite and cristobalite. Sharing of apical oxygens (which point in alternate directions) between silicons in adjacent sheets generates a continuous framework. (After J. J. Papike and M. Cameron, Crystal chemistry of silicate minerals of geophysical interest, Rev. Geophys. Space Phys., 14:37–80, 1976)

 

 

 

 

 

 

The transformation between the various forms are of two types. Displacive transformations, such as inversions between high-temperature (β) and low-temperature (α) forms, result in a displacement or change in bond direction but involve no breakage of existing bonds between silicon and oxygen atoms. These transformations take place rapidly over a small temperature interval and are reversible. Reconstructive transformations, in contrast, involve disruption of existing bonds and subsequent formation of new ones. These changes are sluggish, thereby permitting a species to exist metastably outside its defined pressure-temperature stability field. Two examples of reconstructive trans­formations are tridymite ⇌ quartz and quartz ⇌ stishovite.

 

 

Crystal structure

 

The crystal structures of all silica polymorphs except stishovite contain silicon atoms surrounded by four oxygens, thus producing tetrahedral coordination polyhedra. Each oxygen is bonded to two silicons, creating an electrically neutral framework. Stishovite differs from the other silica minerals in having silicon atoms surrounded by six oxygens (octahedral coordination). See also: Crystal structure

Ideal high tridymite is composed of sheets of SiO4 tetrahedra oriented perpendicular to the c crystallographic axis (Fig. 1) with adjacent tetrahedra in these sheets pointing in opposite directions. The apical oxygens of the tetrahedra are bonded to silicons in neighboring sheets, thus generating a continuous framework with hexagonal (P63/mmc) symmetry. Naturally occurring meteoritic high tridymite deviates somewhat from the ideal structure because adjacent sheets are slightly offset, producing orthorhombic (C2221) symmetry (Fig. 2). Low tridymite has a similar structure, but displacive transformations alter the geometry of the sheets, producing a lower symmetry. Terrestrial low tridymite contains three types of sheets in a complex stacking sequence resulting in triclinic (pseudoorthorhombic) symmetry. Meteoritic low tridymite has a monoclinic (Cc) cell caused by a distortion of the hexagonal rings shown in Fig. 1. Little is known about the structure of middle tridymite, but it is believed to be hexagonal.

 

 

Fig. 2  Portion of the orthorhombic high tridymite structure viewed down the c axis showing the slight displacement between adjacent sheets. In ideal hexagonal high tridymite, neighboring sheets would be superimposed. (After J. J. Papike and M. Cameron, Crystal chemistry of silicate minerals of geophysical interest, Rev. Geophys. Space Phys., 14:37–80, 1976)

 

 

 

 

 

 

High cristobalite, like tridymite, is composed of parallel sheets of SiO4 tetrahedra with neighboring tetrahedra pointing in opposite directions. However, the hexagonal rings are distorted and adjacent sheets are rotated 60° with respect to one another, resulting in the geometry shown in Fig. 3. The structure is cubic (Fd3m), with the layers of tetrahedra parallel to (111). Further distortion of these sheets at low temperature causes an inversion to tetragonal (P41212 or P43212) low crystobalite.

 

 

Fig. 3  Portion of cubic high cristobalite illustrating the distortion of tetrahedral sheets, which are oriented parallel to (111), and the 607 rotation of adjacent sheets. (After J. J. Papike and M. Cameron, Crystal chemistry of silicate minerals of geophysical interest, Rev. Geophys. Space Phys., 14:37–80, 1976)

 

 

 

 

 

 

Coesite also contains silicon atoms tetrahedrally coordinated by oxygen. These polyhedra share corners to form chains composed of four-membered rings. Bonding between chains creates a continuous framework of monoclinic symmetry (C2/c) in which each oxygen is shared between two silicons. The structure, part of which is shown in Fig. 4, is significantly more dense than tridymite and cristobalite.

Silicon in stishovite is octahedrally coordinated by oxygen (Fig. 5). These coordination polyhedra share edges and corners to form chains of octahedra parallel to the c crystallographic axis. The resultant structure is tetragonal (P42/mnm) and is similar to the structure of rutile (TiO2).

 

Properties

 

The physical, optical, and chemical characteristics of the silica polymorphs vary in relation to their crystal structures. Stishovite, which has the most dense packing of constituent atoms, has the highest specific gravity (G = 4.35), whereas tridymite and cristobalite, which have relatively open structures, are the silica polymorphs with the lowest specific gravities (GTr = 2.26, GCr = 2.32). In all instances the β forms of the polymorphs have less densely packed structures, hence lower specific gravities, than α counterparts. The refractive indices of the polymorphs are also related to crystal structure, and the mean refractive index is proportional to specific gravity, ranging from 1.47 for tridymite to 1.81 for stishovite. Specific optical and physical properties for the various silica minerals are given in the table. Twinning is common in most of these minerals and, as in the case of quartz, frequently results from the inversion of a high-temperature, high-symmetry form to a low-temperature, low-symmetry form.

Chemically, all silica polymorphs are ideally 100% SiO2. However, unlike quartz which commonly contains few impurities, the compositions of tridymite and cristobalite generally deviate significantly from pure silica. This usually occurs because of a coupled substitution in which a trivalent ion such as Al3+ or Fe3+ substitutes for Si4+, with electrical neutrality being maintained by monovalent or divalent cations occupying interstices in the relatively open structures of these two minerals. Such substitutions may decrease the SiO2 content to as little as 95 wt %.

 

 

Fig. 4  Four-membered tetrahedral rings in coesite as seen in the plane perpendicular to the twofold axis of symmetry. Bonding between rings in this plane and adjacent planes (as indicated by the single ring at an elevated level) produces the dense coesite structure. (After J. J. Papike and M. Cameron, Crystal chemistry of silicate minerals of geophysical interest, Rev. Geophys. Space Phys., 14:37–80, 1976)

 

 

 

 

 

 

 

 

Fig. 5  Octahedral oxygen coordination polyhedra surrounding central silicons in the stishovite structure. The numbers (0, 0.5) indicate the relative height of the silicon atoms on the a crystallographic axis. The chains of octahedra formed by sharing polyhedral edges run parallel to the c axis. Chains at different levels are interconnected by silicon-oxygen bonds. (After J. J. Papike and M. Cameron, Crystal chemistry of silicate minerals of geophysical interest, Rev. Geophys. Space Phys., 14:37–80, 1976)

 

 

 

 

 

 

 

Phase relationships and occurrences

 

The pressure-temperature stability fields for the silica minerals are shown in Fig. 6. Low quartz is the stable form at low temperatures and pressures. At 1 atm (102 kilopascals) high quartz is stable from 573 to 870°C (1063 to 1600°F), at which temperature tridymite becomes stable. The stability field for cristobalite ranges from 1470 to 1720°C (2680 to 3130°F), at which temperature SiO2 melts. Coesite and stishovite are stable only at exceedingly high pressures. However, because of the sluggishness with which reconstructive changes take place, various polymorphs can exist and occasionally crystallize metastably outside their defined limits of stability. Thus high tridymite (β2) can persist to 163°C (325°F), at which temperature it transforms displacively to middle tridymite (β1), which exists metastably to 117°C (243°F), at which low (α) tridymite forms. Impurities or disordering of the stacking sequence can cause these transformation temperatures to vary and possibly merge into a single high/low transformation. High cristobalite can exist metastably below 1450 to 268°C (or 2640 to 514°F; or lower), at which temperature it transforms displacively to low cristobalite. Variation in the temperature of inversion is caused by compositional variability and stacking disorder.

 

 

Fig. 6  Stability field of the silica polymorphs. °F = (°C × 1.8) + 32.1 kilobar = 102 MPa. (After C. Klein and C. S. Hurlbut, Jr., Manual of Mineralogy, 21st ed., John Wiley and Sons, 1993)

 

 

 

 

 

 

Tridymite and cristobalite crystallize as primary minerals in siliceous volcanic rocks such as rhyolites, trachytes, and andesites in which both minerals may occur in vesicles (cavities) or in the groundmass. Less frequently they are found in basaltic igneous rocks. They have been identified in siliceous sedimentary rocks, such as sandstones and arkoses, subjected to high temperatures during contact thermal metamorphism. Both minerals also occur in the silicate portions of meteorites and in some lunar igneous rocks. Cristobalite, and to a much lesser extent tridymite, is a common intermediate product in the transformation of amorphous biogenic silica (opal A) in marine sediments to quartz during the process of diagenesis (the lithification and alteration of unconsolidated sediments at low temperatures and pressures). The diagenetic sequence is as follows: opal A ⇂ opal CT (= porcellanite = cristobalite containing nonessential water in which the stacking sequence of silica sheets is disordered) ⇂ chalcedony (microcrystalline quartz) ⇂ quartz.

Naturally occurring coesite and stishovite were both first discovered in the Coconino sandstone adjacent to Meteor Crater, Arizona. These minerals are considered mineralogical indicators of the high pressures associated with the shock wave that resulted from the crater-forming meteorite impact. Since their discovery at Meteor Crater, they have both been found in the highly shocked rocks surrounding Ries Crater in Germany. Coesite has also been identified in rocks associated with several other impact or suspected impact craters, in tektites, in material ejected from human-caused explosion craters, and as inclusions in diamonds from kimberlite originating at great depth (hence at great pressure) within the Earth. Stishovite and coesite, therefore, are of great geological significance as indicators of very high pressure due either to depth of formation within the Earth or to shock processes such as those occurring during meteorite impact. See also: Coesite; Quartz; Stishovite

John C. Drake

 

Bibliography

 

 

• W. H. Blackburn and W. H. Dennen, Principles of Mineralogy, 2d ed., 1993
• W. A. Deer, R. A. Howie, and J. Zussman, Rock Forming Minerals, vol. 4: Framework Silicates, 1963
• C. Frondel, Dana's The System of Mineralogy, 7th ed., vol. 3: Silica Minerals, 1962
• K. Frye (ed.), The Encyclopedia of Mineralogy, 1982
• J. J. Papike and M. Cameron, Crystal chemistry of silicate minerals of geophysical interest, Rev. Geophys. Space Phys., 14:37–80, 1976
• D. Stöffler, Coesite and stishovite in shocked crystalline rocks, J. Geophys. Res., 76:5474–5488, 1971  Alifazeli=egeology.blogfa.com

 

 

 

 

تسونومی و رسوبات-Tsunamiites and seismites

Tsunamiites and seismites

Records of some of nature's most catastrophically powerful, short-lived phenomena are preserved in sediments and ancient sedimentary rocks called tsunamiites and seismites, which are important not only for reconstructing ancient events but also for evaluating future hazards. Tsunamiites (also spelled tsunamites) are composed of predominantly marine or lake sediments that were transported by huge water waves called tsunami, which are set into motion by earthquakes, massive landslides, volcanic explosions, and asteroid impacts. Tsunamiites also include continental sediments that were transported when tsunamis suddenly dislodged offshore sediment and mixed it with water to form turbidity currents—dense, turbid masses that flow down the slopes of large lakes and oceans and then commonly travel for hundreds of kilometers over the deep sea floor before being deposited. A seismite is a sediment or sedimentary rock that contains textures and structures produced by earthquake shaking; unlike a tsunamiite, the sediments in seismites were already accumulating when the disturbance occurred. By comparing their pre- and postearthquake features, seismites can be used to calculate the timing, frequency, and local intensities of historic or prehistoric earthquakes, the orientation of their stresses, and even the basin slopes on which the sediments were accumulating. Various seismite features may even provide estimates of the magnitude of the earthquake that produced them, if the distance to the epicenter and the depth of focus are reasonably well known.

 

Tsunamiites

 

Adjacent to beaches where high wave energies have concentrated abundant coarse sediment, tsunamiites brought onshore can consist mainly of sand or gravel and even boulders (Fig. 1a). A tsunamiite may be indistinguishable from the deposits of a powerful storm; however, the seaward backwash of a storm surge is much more gradual and less violent than that of a tsunami. In addition, a tsunami may be propagated as more than one wave; thus the sediment fabric of a tsunamiite may document multiple reversals of flow (Fig. 1b,c).

 

 

Fig. 1  Tsunamiites. (a) Cobbly gravel deposited on the backshore of the Hirahama coast in Japan by a tsunami in 1993. (b) Schematic diagram showing tsunamiites deposited by an incoming wave and its backwash, which also may dislodge offshore material to generate turbidity current. (c) Stratigraphy left by two tsunami waves and their return flows during the Hirahama tsunami of 1993. (Parts a and c modified with permission from F. Nanayama et al., Sedimentary differences between the 1993 Hokkaido-nansei-oki tsunami and the 1959 Miyakojima typhoon at Taisei, southwestern Hokkaido, northern Japan, Sed. Geol., 135:255–264, 2000)

 

 

 

بزرگنمایی تصویر

 

 

 

Tsunamiites may be remarkably thin compared to their hazard significance. Near Seattle, Washington, widespread marshes more than 2 km (1.2 mi) inland are draped with fine sand containing marine fossils only 5–15 cm (2–6 in.) thick, documenting a major earthquake and tsunami about 1000 years ago. Modern tsunamiites record how high above sea level the tsunami rose, from which the magnitude of the earthquake can be estimated. The Seattle tsunamiite is draped over terrain with a vertical relief of 2 m, indicating an earthquake of magnitude 7.5 or greater.

Volcanic explosions can set huge tsunamis into motion (Fig. 2). When the Santorini volcano in the Mediterranean violently collapsed 3500 years ago, tsunamis deposited raised beaches of pumice as far away as Israel, 1000 km (620 mi) to the east. Concurrently, 600 km (370 mi) west of the volcano, 65 km3 (15.6 mi3) of homogeneous mud was suddenly deposited in waters 4000 m (13,000 ft) deep, covering an area of about 5000 km2 (1930 mi2) and up to 24 m (79 ft) thick. The sediment, containing shallow-water carbonate minerals and fossils from the Gulf of Sirte off northern Africa, is interpreted as a tsunamiite or megaturbidite left by enormous turbidity currents that tsunamis triggered 800 km (500 mi) away from Santorini.

 

 

Fig. 2  Tsunamiites left in the deep Mediterranean by the Santorini explosion of 3500 years ago. (Data from M. B. Cita and G. Aloisi, Deep-sea tsunami deposits triggered by the explosion of Santorini (3500 y BP), eastern Mediterranean, Sed. Geol., 135:181–203, 2000; and W. Heicke and F. Werner, The Augias megaturbidite in the central Ionian Sea (central Mediterranean) and its relation to the Holocene Santorini event, Sed. Geol., 135:205–218, 2000)

 

 

 

 

 

 

Large as the Santorini tsunami must have been, it is dwarfed by the waves generated when asteroids hit the ocean at typical velocities of over 70,000 km/h (43,500 mi/h). The impact of an asteroid 10–14 km (6–9 mi) in size at the end of the Cretaceous Period 65 million years ago is blamed for killing off half of Earth's species, including the dinosaurs. It is believed to have made the Chicxulub crater, 180 km (112 mi) in diameter, which is centered on the Yucatan coast of Mexico (Fig. 3a). Large tsunamis left deposits at least 500 km (310 mi) inland, and as far from Chicxulub as northeastern Brazil. At numerous sites around the Gulf of Mexico and the Caribbean, sedimentary rocks 65 million years old with tsunamiite features are believed to be related to the end-Cretaceous asteroid because they contain microscopic glass droplets formed when rocks were melted by the heat of impact; quartz grains with their normal, invisible crystallographic structure cut across by microscopic “shock lamellae” from the enormous pressures of impact; and iridium, an element very rare on Earth but much more abundant in meteorites.

 

 

Fig. 3  Tsunamiites from the asteroid impact 65 million years ago. (a) Location of the Chicxulub impact site showing sites around and in the Gulf of Mexico and Caribbean Sea where tsunamiites from the impact have been recognized. Oceanic sites were cored by the Deep-Sea Drilling Project and Ocean Drilling Project. (b) Stratigraphy of the Peñalver Formation. (Modified from H. Takayama et al., Origin of the Peñalver Formation in northwest Cuba and its relation to K/T boundary impact event, Sed. Geol., 135:295–320, 2000)

 

 

 

 

 

 

One particularly well-studied end-Cretaceous tsunamiite is the Peñalver Formation of northwestern Cuba, a limestone 180 m (590 ft) thick (Fig. 3b). Its basal 25 m (62 ft) is a coarse conglomerate of angular fragments, including boulders a meter in diameter, that were ripped out of the underlying carbonate layer while it was still soft. Despite the coarseness of the basal conglomerate, its fabric—the fact that it is mixed randomly with smaller-grained sediment—documents that the debris was quickly mixed with water and flowed as a liquid. Above this base, the formation becomes finer upward, from carbonate sand interbedded with 14 thin layers of conglomerates probably brought in by successive tsunamis to calcareous mudstone. Abundant water-escape structures in the middle member of the formation suggest that the material settled rapidly from high-density suspensions, much like the Santorini tsunamiite. Remarkably, although most carbonate sediments are produced by organisms, the limestone was undisturbed by burrowing animals, and its fossils are mainly forms reworked from the underlying rocks. This suggests that the limestones were violently resuspended and then quickly redeposited during a period of major biological catastrophe.

 

Seismites

 

The term “seismites” has been widely used since it was first proposed in 1969 to describe fault-graded beds (graded beds cut by faults) of mud that were accumulating in the quiet water of deep marine basins and were still relatively unconsolidated when earthquakes deformed them. Seismites centimeters to a meter (inches to several feet) thick also have been found in marsh and lake deposits. The term “seismite” refers only to sedimentary masses (such as fault-graded beds) that were deformed without being moved any significant lateral distance. Thus, seismoturbidites, the deposits of turbidity currents that were triggered by earthquakes, are not seismites; they contain no original and fault-induced structures that could be used for evaluating seismic stresses and strains.

Homogeneous mud gradually changes in consistency downward as its own accumulating weight squeezes out water, making it more compact. Thus, from top to bottom, it responds to seismic shaking in liquid, plastic, and brittle fashion. Depending on the intensity of shaking, fault-graded or mixed beds may have a surface zone in which liquefaction and shaking has obliterated all depositional structures. Beneath it is a zone of somewhat compacted sediment fragments encased in a soupy matrix that may also have been intruded from below by plumose structures (Fig. 4). Next is a zone of more compacted mud, commonly broken by miniature step faults that may indicate how the basin floor originally sloped. A basal, relatively undisturbed zone of yet more compacted sediment may be cut by larger faults spaced a few meters apart. Several such paleoseismograms may be preserved in a single stratigraphic sequence. The ages of the earthquakes and their recurrence periods can be dated using radioactive isotopes, fossils, and archeological artifacts in the seismites and in the undisturbed deposits between them.

 

 

Fig. 4  Seismite features. (Modified extensively from M. A. Rodriguez-Pascua et al., Soft-sediment sediment deformation structures in lacustrine sediments of the Prebetic Zone, SE Spain, and their potential use as indicators of earthquake magnitudes during the Late Miocene, Sed. Geol., 135:117–135, 2000)

 

 

 

 

 

 

Seismite structures produced by weak shocks may simply be minor disruptions of laminae, or loop bedding (that is, laminate bundles pulled apart by tensile stresses). Among the most common seismite structures are convolute folds, typically about a meter long and a few meters high, encased in undeformed beds. Clearly, they were produced while still soft and close to the sea or lake floor.

Additional seismite features form in interbedded muds and sands. Shocks from earthquakes with magnitudes greater than 5 commonly liquefy watersaturated sands by breaking their grain-to-grain contacts. For the minute or so that the shaking lasts, the mixture of sand and water behaves as a slurry without strength. Driven by the weight of overlying deposits, the slurry rises up through vertical fissures in overlying sequences and is emplaced as sand dikes. The dikes may remain connected to their source beds, or may themselves serve as sources for subordinate, orthogonally oriented dikes that have no direct contact with the mother bed. Liquefaction also may cause water-saturated gravels to intrude overlying beds as irregular-shaped dikes, but gravels are less susceptible to liquefaction than sands, and such dikes may be fractured and faulted by continuing shocks after they are emplaced. Rising silts and sands may also escape entirely from their sources to form isolated, concave-upward pillow structures in overlying layers. Other masses of disturbed sediment can rise or sink, depending on their densities, and be emplaced as isolated pseudonodules.

Along fault scarps, alluvial fans accumulate gravels that are faulted, commonly more than once (Fig. 5). For example, a fault along the front of the Helan Mountains in north-central China experienced an earthquake of magnitude 8 in 1739 that caused 88 km (55 mi) of surface rupture and displaced the famous Great Wall of the Ming dynasty almost a meter vertically and 1.5 m (3.3 ft) laterally. Trenches dug across the fault scarp revealed faulted gravels and colluvial deposits produced not only by the 1739 event but by three older earthquakes as well that recurred at intervals of 2300–3000 years, as determined by radiocarbon dating.

 

 

Fig. 5  Alluvial-fan seismite in China. (Modified from Q. Deng and Y. Liao, Paleoseismology along the range-front fault of Helan Mountains, north central China, J. Geophys. Res., 101(B3):5873–5893, 1996)

 

 

 

 

 

 

 See also: Earthquake; Fault and fault structures; Sedimentology; Stratigraphy; Tsunami; Turbidity current; Volcano

Kelvin S. Rodolfo

 

Bibliography

 

 

• B. F. Atwater and A. L. Moore, A tsunami about 1000 years ago in Puget Sound, Washington, Science, 258:114–116, 1992
• M. B. Cita and G. Aloisi, Deep-sea tsunami deposits triggered by the explosion of Santorini (3500 y BP), eastern Mediterranean, Sed. Geol., 135:181–203
• P. Claeys et al., Distribution of the Chicxulub ejecta at the Cretaceous-Tertiary boundary, Geol. Soc. Amer. Spec. Pap., 359:55–69, 2002
• Q. Deng and Y. Liao, Paleoseismology along the range-front fault of Helan Mountains, north central China, J. Geophys. Res., 101(B3):5873–5893, 1996
• W. Heicke and F. Werner, The Augias megaturbidite in the central Ionian Sea (central Mediterranean) and its relation to the Holocene Santorini event, Sed. Geol., 135:205–218, 2000
• A. Seilacher, Fault-graded beds interpreted as seismites, Sedimentology, 13:155–159, 1969
• T. Shiki and T. Yamazaki, Tsunami-induced conglomerates in Miocene upper bathyal deposits, Chita Peninsula, central Japan, Sed. Geol., 104:175–188, 1996

 

Additional Readings

 

 

• F. N. Ettensohn et al. (eds.), Ancient seismites, Geol. Soc. Amer. Spec. Pap., vol. 359, 2002
• G. Ryder et al. (eds.), The Cretaceous-Tertiary event and other catastrophes in Earth history, Geol. Soc. Amer. Spec. Pap., vol. 307, 1996
• T. Shiki et al. (eds.), Sed. Geol., vol. 104, Spec. Iss.: Marine Sedimentary Events and Their Records, 1996
• T. Shiki et al. (eds.), Sed. Geol., vol. 135, Spec. Iss.: Seismoturbidites, Seismites and Tsunamiites, 2000
• General Reference for Tsunamiites and Seismites
• Description of Grand Banks Earthquake and Tsunamiite
• Photographs of Ordovician Seismites in Kentucky
• K/T Boundary in Badlands National Park
• Cretaceous-Tertiary Boundary Stratigraphy Including Possible Seismites and Tsunamiites in Badlands National Park, South Dakota
• U.S. Geological Survey Study of the Sedimentary Deposits of the 1998 Papua New Guinea Tsunami
•  Alifazeli=egeology.blogfa.com

 

رسوبات عمیق دریایی-Deep-marine sediments

Deep-marine sediments

 

The term “deep marine” refers to bathyal sedimentary environments occurring in water deeper than 200 m (650 ft), seaward of the continental shelf break, on the continental slope and the basin (Fig. 1). The continental rise, which represents that part of the continental margin between continental slope and abyssal plain, is included under the broad term “basin.” An example of well-developed shelf, slope, and basin settings can be seen in the modern Gulf of Mexico (Fig. 2). On the slope and basin environments, sediment-gravity processes (slides, slumps, debris flows, and turbidity currents) and bottom currents are the dominant depositional mechanisms, although pelagic and hemipelagic deposition is also important.  See also: Basin; Continental margin; Gulf of Mexico; Marine sediments

 

 

Fig. 1  Slope and basinal deep-marine sedimentary environments occurring at water depths greater than 200 m (650 ft). Slides, slumps, debris flows, turbidity currents, and various bottom currents are important processes in transporting and depositing sediment in the deep sea. Note the complex distribution of deep-marine deposits.

 

 

بزرگنمایی تصویر

 

 

Fig. 2  Bathymetric display of the northern Gulf of Mexico based on high-resolution bathymetric swath data. The data are displayed to give a 3-D perspective of present-day shelf, slope, and basin settings. Note the highly irregular sea-floor topography and intraslope basins controlled by salt tectonism. The intraslope basins commonly range in width from 10 to 20 km (6.2 to 12.4 mi). GB = Garden Bank area; GC = Green Canyon area; MT = mass transport (slide/slump/debris flow); MB = minibasin; LD = linear depression; MC = Mississippi Canyon; KC = Keathley Canyon.

 

 

 

 

 

 

In 1960s and 1970s, most deep-marine sedimentary systems were thought to be dominated by submarine-fan deposits formed by turbidity currents regardless of whether they were fan-shaped in morphology or whether they were turbidites in origin. A critical evaluation of deep-marine sedimentary systems revealed that, in fact, they are quite complex and that simple fan models are not realistic. Furthermore, the basic tenets of understanding deep-marine processes and their deposits have recently been challenged. In light of these developments, this article focuses on types of deep-marine processes and their deposits; and discusses aspects of deep-marine sedimentary environments in the categories of (1) submarine slopes, (2) submarine canyons and gullies, (3) submarine channels, (4) submarine fans, and (5) submarine basin plains.  See also: Submarine canyon; Turbidite; Turbidity current

 

Types of processes

 

The mechanics of deep-marine processes is critical in understanding the nature of transport and deposition of sand and mud in the deep sea. In deep-marine environments, gravity plays the most important role in transporting and depositing sediments. Sediment failure under gravity near the shelf edge commonly initiates gravity-driven deep-marine processes, such as slides, slumps, debris flows, and turbidity currents (Fig. 1). Sedimentary deposits reflect only depositional mechanisms, not transportational mechanisms. Therefore, general characteristics of each process listed below should be used only to interpret depositional mechanisms.  See also: Gravity; Sedimentology

 

Slides

 

A slide is a coherent mass of sediment that moves along a planar glide plane and shows no internal deformation (Fig. 1). Slides represent translational movement. Submarine slides can travel hundreds of kilometers. For example, the runout distance of Nuuanu Slide in offshore Hawaii is 230 km (143 mi). Long runout distances of 50–100 km (31–62 mi) of slides are common.

General characteristics of submarine slide deposits are:

• Gravel-to-mud lithofacies.
• Large dimensions, up to hundreds of kilometers long.
• Glide plane scar on the main shear surface, and basal zone of shearing.
• Upslope areas contain tensional faults.
• Downslope edges may show compressional folding or thrusting (that is, toe thrusts).
• Common in areas of tectonics, earthquakes, salt movements, and rapid sedimentation.
• Sheetlike geometry with irregular thickness and shape.
• Occur on slopes as gentle as 1–4°.
• Travel hundreds of kilometers down the gentle slopes (Fig. 3).

 

 

 

Fig. 3  Map showing long-distance transport of modern slides of about 300 km (186 mi) downslope from the shelf edge (100-fathom, 183-m, or 600-ft contour line) off northwestern Africa. Thickness of slide = 25 m (82 ft). (From R. D. Jacobi, Sediment slides on the northwestern continental margin of Africa, Mar. Geol., 22:157–173, 1976)

 

 

 

 

 

 

Examples of submarine slides reveal that modern ones are an order of magnitude larger than their known ancient counterparts. The modern Storegga Slide on the Norwegian continental margin has a width of 290 km (180 mi), a thickness of 430 m (1411 ft), and a width-to-thickness ratio of 675:1 (see table). A seismic-scale slide sheet was described, 6 km (3.7 mi) wide and 440 m (1444 ft) thick with a width-to-thickness ratio of 14:1, from the Ablation Point Formation (Kimmeridgian) exposed in Alexander Island of Antarctica. Individual sandy slide sheets reach a thickness of 50 m (164 ft) and a length of 1000 m (3281 ft).

 

Slumps

 

A slump is a coherent mass of sediment that moves on a concave-up glide plane and undergoes rotational movements causing internal deformation (Fig. 1).

General characteristics of submarine slump deposits are:

• Gravel-to-mud lithofacies.
• Basal zone of shearing.
• Slump folds.
• Contorted bedding.
• Steeply dipping and truncated layers.
• Undeformed units above deformed units.
• Lenticular-to-sheetlike geometry with irregular thickness.

 

Fig. 4  Conceptual model showing interaction between channelized (axial) turbidity currents, overbank turbidity currents, and bottom currents (contour currents) in deep-sea environments. (From G. Shanmugam, T. D. Spalding, and D. H. Rofheart, Process sedimentology and reservoir quality of deep-marine bottom-current reworked sands (sandy contourites): An example from the Gulf of Mexico, Amer. Ass. Petrol. Geol. Bull., 77:1241–1259, 1993)

 

 

 

 

 

 

 

Fig. 5  Modern United States continental slope showing variability in slope values among different areas: (a) New Jersey-Maryland: (2.5^°); (b) Louisiana: (0.5^°); and (c) Oregon: (2.0^°). (From L. F. Pratson and W. F. Haxby, What is the slope of the U.S. continental slope?, Geology, 24:3–6, 1996)

 

For example, one study reported a thickness of slumps varying from 10 to 374 m (33 to 1227 ft). Lateral dimensions (width/length) reached up to 64 km (40 mi). Width-to-thickness ratio varied from 10:1 to 3600:1 (see table). In the Cretaceous of the Norwegian North Sea, abrupt lateral changes in slump facies thickness occur across a distance of 29 m (95 ft) between two adjacent wells. Such abrupt terminations of slump deposits have been observed in outcrops in Antarctica. Slump and slide deposits show large vertical dimensions.  See also: Antarctica; Cretaceous

 

Debris flows

 

A downslope increase in mass disaggregation results in the transformation of slumps into debris flows (Fig. 1). Sediment is now transported as an incoherent viscous mass, as opposed to a coherent mass in slides and slumps. A debris flow is a sediment-gravity flow with plastic rheology (that is, fluids with yield strength) and laminar state. Deposition from debris flows occurs through freezing. The term “debris flow” is used here for both the process and the deposit of that process. The terms “debris flow” and “mass flow” are used interchangeably because each exhibits plastic flow behavior with shear stress distributed throughout the mass. Although only muddy debris flows (debris flows with mud matrix) received attention in the past, recent experimental and field studies show that sandy debris (debris flows with sand matrix) flows are equally important. Rheology is more important than grain-size distribution in controlling sandy debris flows, and the flows can develop in slurries of any grain size (very fine sand to gravel), any sorting (poor to well), any clay content (low to high), and any modality (unimodal and bimodal).  See also: Rheology

General characteristics of debris-flow deposits (both muddy and sandy types) are:

• Gravel-to-mud lithofacies.
• Floating or rafted mudstone clasts near the tops of beds.
• Inverse grading of floating mudstone clasts.
• Inverse grading, normal grading, inverse-to-normal grading, and no grading of matrix.
• Floating quartz pebbles and inverse grading of granules.
• Pockets of gravels.
• Planar clast fabric.
• Preservation of fragile shale clasts.
• Irregular, sharp upper contacts and lateral pinch-out geometries.
• Side-by-side occurrence of garnet granules (density 3.5–4.3) and quartz granules (density 2.65).

Lenticular-to-sheetlike geometry.

 

Using outcrop measurements, it can be established that debris-flow deposits are likely to have a width-to-thickness ratio of 30–50:1. In the examples used in this study, thicknesses of debris-flow deposits commonly range from 1 to 60 m (3 to 197 ft), but unusually thick deposits may occur in association with large slides. An example is the Storegga Slide in the Norwegian continental margin, 430 m (1411 ft) thick and 290 km (180 mi) wide, in which debris flow is a dominant facies. Debris-flow facies have width-to-thickness ratios of up to 500:1 or more. This is generally the result of amalgamation of depositional units.

 

 

Turbidity currents

 

With increasing fluid content, plastic debris flows tend to become newtonian turbidity currents (Fig. 1). Although turbidity currents can constitute a distal end member occurring in basinal areas, they can occur in any part of the system (proximal and distal). Turbidity currents can also occur above debris flows due to flow transformation in density-stratified flows. A turbidity current is a sediment-gravity flow with newtonian rheology (that is, fluids without yield strength) and turbulent state. Deposition from turbidity currents occurs through suspension settling. Deposits of turbidity currents are called turbidites. Although turbidity currents have received a lot of emphasis in the past, other processes are equally important in the deep sea (Fig. 1). In terms of transporting coarse-grained sediment into the deep sea, sandy debris flows and other mass flows appear to play a greater role than turbidity currents.

General characteristics of turbidites are:

• Fine-grained sand–to–mud lithofacies.
• Sharp or erosional basal contact.
• Gradational upper contact.
• Normal grading without complications (that is, no floating clasts or granules).
• Sheetlike geometry in basinal settings. Lenticular geometry may develop in channel settings.

 

In the geologic record, units showing normal grading commonly occur directly above units with inverse grading. Such inverse–to–normally graded beds have traditionally been interpreted as high-density turbidites. However, the concept of high-density turbidity currents and their deposits have come under severe criticism because there are at least five different ways that the concept of high-density turbidity currents has been defined: (1) flow density, (2) grain size, (3) driving force, (4) velocity, and (5) rate of deposition. These five concepts are not in agreement in terms of fluid dynamics. As a result, there are no standard and meaningful ways of interpreting deep-water rocks as high-density turbidites. In terms of fluid rheology and flow state, high-density turbidity currents are nothing but sandy debris flows.  See also: Fluid mechanics; Geologic time scale

Submarine fan lobes (lobate deposits that accumulate at the mouths of canyons and channels) show width-to-thickness ratios of 165:1 to 1200:1. Lateral dimensions of turbidites deposited on modern abyssal plains can be unusually large. For example, the “Black Shell turbidite” on the Hatteras Abyssal Plain in the western North Atlantic Ocean is 4 m (13 ft) thick and extends for 500 km (311 mi), with a width-to-thickness ratio of 125,000:1. It covers an area of 44,000 km2 (16,984 mi2). Classic examples of basin-plain turbidites, such as those exposed along the foreshore at Zumaya, Spain, are usually in the range of 10 cm to 1 m (4 in. to 3 ft) in thickness, but they can be traced for several kilometers. Extensive sheetlike turbidites deposited in the open oceans of Atlantic-type margins are seldom preserved in the geologic record.

 

Bottom currents

 

In large modern ocean basins, such as the Atlantic, thermohaline-induced geostrophic bottom currents within the deep and bottom water masses commonly flow approximately parallel to bathymetric contours (that is, along the slope; Fig. 1). They are generally referred to as contour currents. However, because not all bottom currents follow regional bathymetric contours, it is preferred that the term “contour current” be applied only to currents flowing parallel to bathymetric contours, and other currents be termed bottom currents. For example, wind-driven surface currents may flow in a circular motion (Fig. 1) and form eddies that reach the deep-sea floor, such as the Loop Current in the Gulf of Mexico, and the Gulf Stream in the North Atlantic. Local bottom currents that move up- and downslope can be generated by tides and internal waves, especially in submarine canyons. These currents are quite capable of erosion, transportation, and redeposition of fine-to-coarse sand in the deep sea.  See also: Gulf Stream; Ocean circulation

Bottom currents (1) generally persist for long time intervals and can develop equilibrium conditions; (2) transport sand primarily by traction (bedload movement-sliding), rolling, and saltation; (3) are sometimes free of sediment, and are termed clear-water currents; (4) entrain and transport passive sediment particles; (5) are driven by thermohaline, wind, wave, or tidal forces; and (6) commonly flow parallel to the strike of the regional slope (Fig. 4) but can also flow in circular motions (gyres) unrelated to the slope. These characteristics clearly discriminate deep-sea bottom currents from turbidity currents. Bottom currents operate parallel to the strike of the slope in most deep-marine settings independently of downslope turbidity currents, debris flows, and slumps. As a result, bottom currents can constantly rework sands introduced into the basin episodically by downslope gravity processes.  See also: Tidal power

Deposits of contour currents (bottom currents) have been termed contourites. However, the general term “bottom-current-reworked sands” is preferred for all types of reworking in the deep sea.

General characteristics of bottom-current-reworked sands are:

• Fine-grained sand and silt lithofacies.
• Thin-bedded to laminated sand (usually less than 5 cm or 2 in.) in deep-marine mud.
• Rhythmic occurrence of sand and mud layers.
• Sharp (nonerosional) upper contacts and sharp-to-gradational bottom contacts.
• Internal erosional surfaces.
• Well-sorted sand and low depositional matrix (clean sand).
• Inverse size grading (coarsening upward) at various scales.
• Horizontal lamination and low-angle cross lamination.
• Cross bedding.
• Lenticular bedding/starved ripples.
• Current ripples and wave ripples.
• Mud offshoots in ripples.
• Mud-draped ripples.
• Alternating traction and suspension structures.
• Absence of associated thin units with normal grading.
• Double mud layers (tidal).
• Sigmoidal cross bedding (tidal).
• Flaser bedding.
• Occurrence of sand layers with traction structures in discrete units, but not as part of a vertical sequence of structures, such as the Bouma Sequence with basal graded division.
• Lenticular-to-sheetlike geometry.

 

 

Pelagic and hemipelagic settling

 

Pelagic and hemipelagic processes generally refer to settling of mud fractions derived from the continents and shells of microfauna down through the water column throughout the entire deep-ocean floor (Fig. 1). Hemipelagites are deposits of hemipelagic settling of deep-sea mud in which more than 25% of the fraction coarser than 5 micrometers is of terrigenous, volcanogenic, or neritic origin. Although pelagic mud and hemipelagic mud accumulate throughout the entire deep-ocean floor, they are better preserved in parts of abyssal plains (Fig. 1). Rates of sedmentation vary from millimeters to greater than 50 cm (20 in.) per 1000 years, with the highest rates on the upper continental margin. 

General characteristics of pelagites and hemipelagites are:

• Mud lithofacies.
• Parallel lamination.
• Faint normal grading.
• Bioturbation.
• Deep-marine body fossils and trace fossils.
• Sheetlike geometry conformable to underlying sea-floor topography (drape).

 

 

 

Submarine slope environments

 

Submarine slopes are considered to be of the sea floor between the shelf-slope break and the basin floor (Fig. 1). Modern continental slopes around the world average 4°, but slopes range from less than 1° to greater than 40°. Slopes of active margins (for example, California and Oregon, about 2°) are relatively steeper than those of passive margins (for example, Louisiana, about 0.5°; Fig. 5). On constructive continental margins with high sediment input, gravity tectonics involving salt and shale mobility and diapirism forms intraslope basins of various sizes and shapes (for example, Gulf of Mexico; Fig. 2). Erosional features, such as canyons and gullies, characterize intraslope basins. Deposition of sand and mud occurs in intraslope basins. Slope morphology plays a major role in controlling deep-marine deposition through (1) steep versus gentle gradients, (2) presence or absence of canyons and gullies, (3) presence or absence of intraslope basins, and (4) influence of salt tectonics.  See also: Diapir; Erosion

 

Submarine canyon and gully environments

 

Submarine canyons and gullies are erosional features that tend to occur on the slope. Although canyons are larger than gullies, there are no standardized size criteria to distinguish between them. Submarine canyons are steep-sided valleys that incise the continental slope and shelf. They serve as major conduits for transporting sediment from land and the continental shelf to the basin floor. Modern canyons are relatively narrow, deeply incised, steeply walled, often sinuous valleys with predominantly V-shaped cross sections. Most canyons originate near the continental shelf break and generally extend to the base of the continental slope. Canyons commonly occur off the mouths of large rivers such as the Hudson and Mississippi, although many others, such as the Bering Canyon in the southern Bering Sea, have developed along structural trends.  See also: Bering Sea

Modern submarine canyons vary considerably in their dimensions. Their average length of canyons has been estimated to be about 55 km (34 mi), although the Bering Canyon, the world's longest, is nearly 1100 km (684 mi). The shortest canyons are those off the Hawaiian Islands, with average lengths of about 10 km (6.2 mi).

The water depths in which heads of canyons begin vary from a few meters (near the shoreline) to several hundred meters (shelf break and upper slope). Canyons on the Pacific margin develop at shallower depths than canyons on the Atlantic margin. For example, heads of California canyons begin at an average depth of about 35 m (115 ft), whereas canyons of the east coast of the United States begin at depths greater than 100 m (328 ft). The average depth of canyon termination has been estimated to be 2000 m (6562 ft). The average relief of canyon walls is over 900 m (559 ft). The submarine canyon with the greatest relief is the Great Bahama Canyon, where wall relief is up to 4285 m (14,058 ft).

Physical and biological processes common in submarine canyons are mass wasting, turbidity currents, bottom currents, and bioerosion. Mass wasting is a general term used for the failure, dislodgment, and downslope movement of sediment under the influence of gravity. Common examples of mass wasting are slides, slumps, and debris flows. Major slumping events can lead to formation of submarine canyons. The Mississippi Canyon in the Gulf of Mexico, for example, is believed to have been formed by retrogressive slumping during the Late Pleistocene fall in sea level. During the Holocene rise in sea level, it was partially filled.  See also: Holocene; Mass wasting; Pleistocene

Francis Shepard made significant contributions to understanding the processes that operate and modify modern submarine canyons. His current-meter records from various submarine canyons around the world revealed that bottom currents (rarely exceeding 50 cm/s or 25 in./s) flow almost continuously up or down canyon axes. According to Shepard, tidal and internal wave forces are the major causes for the semidiurnal up- and downcanyon flow reversals observed in submarine canyons. Such currents can move sediment up and down canyons.

Sea-level changes and tectonic settings are important controlling factors in the development of submarine canyons. Geologic evidence suggests that downcutting of submarine canyons by sediment-gravity flows took place primarily during periods of glacially lowered sea levels.  See also: Glaciology

 

Submarine channel environments

 

According to E. Mutti and W. R. Normark, “A channel is the expression of negative relief produced by confined turbidity current flow, and represents a major, long-term pathway for sediment transport.” In other words, all submarine channels are of turbidity current origin, and they all must be long-lived in order to qualify as channels. However, it is not always possible to establish whether a channel in the rock record was cut by a turbidity current or by some other process because processes that cut channels are not necessarily the same processes that fill channels. For example, many channels are filled with deposits of debris flows; however, debris flows are generally not capable of forming erosional channels. Another problem is determining whether an ancient channel acted as a pathway for a long or a short period of time. There are no standardized definitions of long and short periods of time in deep-water sedimentation. Dimensions of submarine channels vary highly in their widths and depths.

Submarine channels can be erosional, aggradational, or both. Erosional channels commonly develop on the slope, whereas aggradational channels with levees tend to develop on basin floors having gentle gradients. Again, recognition of the type of submarine channel is not easy. Also, there are no objective criteria to distinguish erosional canyons and gullies from erosional channels. Many channels that began as an erosional feature became an aggradational feature. Overbank sediments are characteristic of the aggradational type, and are fine-grained, thin-bedded, and current-laminated. Graded mudstones are dominant in overbank deposits. Persistent overbanking can result in the development of positive relief in the form of levees.

Submarine channels on mature passive-margin fans tend to be relatively long, bifurcating, low-gradient, and largely sinuous. By contrast, channels on active-margin fans are short, steep, and low-sinuosity. The relatively fine-grained (mud-rich) character of the transported sediment associated with channels on mature passive-margin fans, such as the Amazon and Mississippi, gives rise to excellent bank stability and favors the development of a single, largely sinuous commonly meandering channel. Channel shifting in such a system is probably mainly by periodic changes in course (avulsion). However, the origin of sinuous and meandering channels in submarine setting is still poorly understood because, unlike prolonged fluvial currents that develop sinuous channel morphology on land, turbidity currents are episodic events. Therefore, it is difficult to envision how these events could develop the highly meandering channel plan forms observed in the deep sea. Some sinuous channels have been attributed to structural origin (that is, they are controlled by fault motions).  See also: Geomorphology

The sand-rich character and steep gradient of active-margin settings should, by contrast, favor development of a braided channel system. A long-range side-scan sonar (GLORIA) survey of the Orinoco deep-sea fan (mixed setting), which abuts against the deformation front of the Barbados Outer Ridge, also reveals a braided distributary system.  See also: Sonar

 

High-amplitude reflection packets

 

In the Amazon Fan in the equatorial Atlantic, channel bifurcation through avulsion is thought to initially lead to deposition of unchannelized sandy flows in the interchannel area (Fig. 6a). Subsequent channel and levee development and progradation over these sandy deposits (Fig. 6b) produces a sheetlike geometry at the base of the new channel-levee system that returns high-amplitude reflection packets (HARPs) on seismic data. These sheetlike HARPs overlain by a channel-levee system (gull-wing geometry) are in many ways identical in appearance to a basin-floor fan overlain by a slope fan in a sequence-stratigraphic framework. However, there is a major difference between a basin-floor fan and a HARP. For example, a basin-floor fan is formed by progradation primarily during lowstands of sea level (allocyclic process), whereas HARPs are the results of channel bifurcation (autocyclic process). More importantly, the basin-floor fan and the slope fan are not contemporaneous, whereas a HARP and its overlying channel-levee system are essentially contemporaneous features. This again illustrates that caution must be exercised in interpreting seismic geometries in terms of processes.  See also: Seismic stratigraphy; Sequence stratigraphy

 

 

Fig. 6  Amazon Fan. (a) Channel bifurcation through avulsion on the deep-sea Amazon Fan results in unchannelized sandy flows by breaching their confining levee through a crevasse and spreading out initially as unchannelized flows into a lower interchannel areas. (b) Reestablishment of a new channel above these sandy deposits can result in a sheetlike geometry that returns high-amplitude reflections (HARPs) on seismic data. (From R. D. Flood et al., Seismic facies and Late Quaternary growth of Amazon submarine fan, in P. Weimer and M. H. Link, eds., Seismic Facies and Sedimentary Processes of Submarine Fans and Turbidite Systems, pp. 415–433, Springer-Verlag, New York, 1991)

 

 

 

 

 

 

 

Submarine-fan environments

 

Submarine fans represent fan-shaped or lobate deposits located in front of submarine canyons or channels. Submarine fans are considered to be formed primarily by turbidity currents coming from a point source. However, not all submarine canyons or channels develop submarine fans. Submarine-fan models, based on turbidite concepts, have been the most influential sedimentologic tools in the petroleum industry for interpreting deep-marine environments.

 

Sedimentologic fan models

 

Normark presented the first widely used sedimentologic model for modern submarine fans based on studies of small, sand-rich fans such as the San Lucas and Navy fans off California. He introduced the term “suprafan” to describe the lobe-shaped bulge on modern fans found immediately downfan of the termination of the major feeder channel. This morphologic feature was presumably formed by rapid deposition of coarse sediment by turbidity currents at the termination of the upper-fan valley. The suprafan lobe was thought to exhibit an overall mounded, hummocky morphology in bathymetric data.

Mutti and Ricci Lucchi proposed submarine-fan models based on outcrop studies in Italy and Spain that popularized the concept of submarine fans with channels in the middle-fan setting and depositional lobes (lobate deposits) in the lower-fan setting. The general characteristics of the depositional lobes of ancient submarine fans include the following: (1) they are considered to develop at or near the mouths of submarine-fan channels analogous to distributary mouth bars in deltaic systems; (2) they show an absence of basal channeling; (3) they usually display thickening-upward depositional cycles composed of classic turbidites; (4) their common thickness range is 3–15 m (9.8–49 ft); and (5) they exhibit sheetlike geometry.

Roger Walker combined the major elements of Normark's model for modern fans with facies concepts of ancient submarine fans of Mutti and Lucchi, and advocated a general fan model with a single feeder channel in the upper-fan area and suprafan lobes in the middle/lower fan areas. Subsequently, this general fan model became influential in hydrocarbon exploration and production because of its predictive capabilities.

 

Sequence-stratigraphic fan models

 

Based partly on the suprafan-lobe model of Normark, seismic facies and geometries are used to classify deep-marine systems into basin-floor fans and slope fans in a sequence-stratigraphic framework, and in turn these models are used to predict deposits of specific depositional processes (for example, turbidity currents). However, “turbidity current” has precise meanings in terms of rheology (newtonian) and flow state (turbulence). Evidence for newtonian rheology and flow turbulence cannot be established directly from seismic-reflection profiles or wireline-log motifs; rather, these properties can be ascertained only from actual sediment facies in cores or outcrops. This is because depositional features in cores and outcrops are only centimeters in scale, and these features are too small to be resolved in seismic data. Furthermore, the interpretation of specific seismic facies and geometries (for example, sheet, mounded, continuous, hummocky) as to sediment processes may vary from one worker to the next depending on one's experience. More importantly, a single depositional facies (sandy debris flows) can generate a variety of external seismic geometries (for example, mounded, sheet, and lateral pinchout) and internal seismic reflection patterns (for example, bidirectional downlap, hummocky/chaotic, and parallel/continuous).

 

Abandonment of submarine-fan models

 

By the early 1980s, so many deep-marine sedimentary systems were studied that the simple submarine models discussed earlier were proved to be obsolete. Because the morphologic characteristics of modern suprafan lobes are either not preserved in the rock record or cannot be planimetrically mapped in outcrops, Normark abandoned his “suprafan lobe” concept altogether. Walker also recently abandoned his general fan model as “a submarine fan model of the channel-depositional lobe type, influential in its time, but now obsolete because it ignored external controls, especially sea level fluctuations.”

The conceptual basin-floor fan model, characterized by mounded seismic facies, predicts sheetlike turbidite sands. However, one study provides a detailed description and interpretation of about 3600 m (12,000 ft) of core through a number of mounded seismic forms of “basin-floor fans” in the North Sea and Norwegian Sea. In these mounded features, turbidites are extremely rare (<1%). Mass-transport deposits, especially slumps, slides, and debris flows, are predominant in the core (50–100%) taken from mounded seismic facies. Recent data also suggest that some of these sands are laterally discontinuous. While features identified as basin-floor fans may occur at specific and predictable stratigraphic positions within a depositional sequence and produce characteristic seismic facies and reflection patterns on seismic data, this core study indicates that basin-floor fans do not represent specific depositional facies (for example, turbidites) and geometries (that is, sheetlike) as the model predicts. Accordingly, seismic mound models (for example, basin-floor fan) should be abandoned because they are based on a model (suprafan-lobe) that is no longer used.

 

 

Submarine basin plain environments

 

The term “abyssal plain” refers to a flat region of the modern deep ocean floor (Fig. 7), where the gradient is less than 1:1000. The geologic community, in referring to ancient examples, commonly uses the more general term “basin plain.” Basin plains form in response to filling and leveling sea-floor topography by ponding of turbidity currents and by other processes. Because of their flat nature, basin plains favor sheetlike geometries for deposits of turbidity currents and other suspension settling processes (pelagic and hemipelagic). Basin plains are also potential areas for sheetlike slides and debris flows. The Enderby abyssal plain occupies a vast area covering nearly 3.7 × 106 km2 (1.4 × 106 mi2; Fig. 7).

 

 

Fig. 7  Sizes of modern abyssal plains. (From P. P. E. Weaver, J. Thomson, and P. M. Hunter, Introduction, in P. P. E. Weaver and J. Thomson (eds.), Geology and Geochemistry of Abyssal Plains, Geol. Soc. London Spec. Publ., no. 31, pp. vii–xii, 1987)

 

 

 

 

 

 

 

Paradigm shift

 

Although the turbidite paradigm began in the 1950s, understanding of depositional processes in deep-marine environments is still in its infancy. No single facies model can adequately represent all deep-marine sedimentary systems. Submarine fan studies are presently in a state of flux. The days of interpreting complex deep-marine sequences as channels and lobes using fan models are over. Sedimentologic and sequence-stratigraphic lineages of fan concepts dominated by lobes show that their popularity escalated in the 1970s and 1980s but declined in the 1990s to the point of abandonment. The turbidite paradigm reached full circle in the 1990s, completing a remarkable scientific journey. Things are much improved in terms of available marine geological data, core, and outcrop studies, theoretical considerations, and flume experiments. Deep-marine systems are quite complex in terms of sea-floor topography, depositional processes, geometries, and stacking patterns. As a result, no single facies model can possibly explain all variations in the complex deep-sea environments.

Mass transport processes (slides, slumps, sand flows, and debris flows) have been observed in modern oceans; however, convincing direct observations of turbidity currents in modern oceans are lacking. It is ironic that there are numerous deep-marine facies models for deposits of turbidity currents that have never been observed, but there are no facies models for deposits of mass flows that are observed. This is perhaps due to the simplicity of turbidity current concepts and submarine-fan models, and the historical association between turbidites and sheet geometries. It is true that basinal turbidites are sheetlike in geometry; however, these turbidite sands are commonly thin-bedded and fine-grained and contain large amounts of mud. In comparison to turbidites, slope sands of debris flow origin are thicker-bedded and coarser-grained and contain lower amounts of mud. The current trend in the petroleum industry is to routinely apply submarine-fan models, developed for base-of-slope settings with smooth sea floors, to intraslope settings with highly irregular sea floors such as in the Gulf of Mexico. However, there is a need to develop separate models for slopes emphasizing slope processes and products. The conventional wisdom that slopes are areas of “bypassing” of sand is not always valid. Slopes are important future target areas where major petroleum reservoirs are waiting to be found. Some slopes undergo extensive gravity tectonic deformation, which leads to development of diapirs and intraslope basins with erosional features, such as canyons and gullies (for example, northern Gulf of Mexico, Niger Delta). Deposition of thick sand bodies can occur in such intraslope basins. Future models should take into account the great wealth of information available on modern and ancient slope processes and products such as ancient sandy slides, muddy slides, slumps, and debris flows. Debris flows can travel hundreds of kilometers on gentle gradients. On continental margins, debris flows tend to develop fingerlike patterns (Fig. 8). These details are seldom included in popular deep-marine models.  See also: Petroleum engineering

 

 

Fig. 8  GLORIA long-range side-scan sonar imagery showing fingerlike patterns of debris flows in the Bear Island Fan on the continental margin of the modern Norwegian-Barents Sea. These fans are composed of cohesive debris flows. Properties of the Bear Island Fan are: runout distance 100–200 km (62–124 mi); width 5–25 km (3–16 mi); thickness 10–50 m (33–164 ft); volume 10–30 km3 (2.4–7.2 mi3); slope angle—midslope 0.5°, lower slope 0.2°. (From A. Elverhoi et al., On the origin and flow behavior of submarine slides on deep-sea fans along the Norwegian-Barents Sea continental margin, Geo-Mar. Lett., 17:119–125, 1997)

 

 

 

 

 

 

A new paradigm for deep-marine systems will emerge that will be more inclusive in terms of debris flows and bottom currents than just turbidity currents and fan models. This paradigm shift will emphasize that deep-marine systems are extremely complex and each case is unique.

G. Shanmugam

 

Bibliography

 

 

• J. D. Clark and K. T. Pickering, Submarine Channels: Processes and Architecture, Vallis Press, London, 1996
• A. Elverhoi et al., On the origin and flow behavior of submarine slides on deep-sea fans along the Norwegian-Barents Sea continental margin, Geo-Mar. Lett., 17:119–125, 1997
• R. D. Flood et al., Seismic facies and Late Quaternary growth of Amazon submarine fan, in P. Weimer and M. H. Link (eds.), Seismic Facies and Sedimentary Processes of Submarine Fans and Turbidite Systems, pp. 415–433, Springer-Verlag, New York, 1991
• D. I. M. Macdonald, A. C. M. Moncrieff, and P. J. Butterworth, Giant slide deposits from a Mesozoic fore-arc basin, Alexander Island, Antarctica, Geology, 21:1047–1050, 1993
• L. F. Pratson and W. F. Haxby, What is the slope of the U.S. continental slope?, Geology, 24:3–6, 1996
• G. Shanmugam, The Bouma Sequence and the turbidite mind set, Earth-Sci. Rev., 42:201–229, 1997
• G. Shanmugam, Fifty years of turbidite paradigm (1950s–1990s): Deep-water processes and facies models—A critical perspective, Mar. Petrol. Geol., 17:285–342, 2000
• G. Shanmugam, High-density turbidity currents: Are they sandy debris flows?, J. Sed. Res., 66:2–10, 1996
• Alifazeli=egeology.blogfa.com

 

 

Additional Readings

 

 

• A. H. Bouma, W. R. Normark, and N. E. Barnes (eds.), Submarine Fans and Related Turbidite Systems, Springer-Verlag, New York, 1985
• J. E. Damuth, The Western Equatorial Atlantic: Morphology, Quaternary Sediments, and Climatic Cycles, unpublished Ph.D. dissertation, Columbia University, 1973
• C. D. Hollister, Sediment Distribution and Deep Circulation in the Western North Atlantic, unpublished Ph.D. dissertation, Columbia University, 1967
• R. D. Jacobi, Sediment slides on the northwestern continental margin of Africa, Mar. Geol., 22:157–173, 1976
• E. Mutti and F. Ricci Lucchi, Turbidites of the northern Apennines: Introduction to facies analysis (transl. by T. H. Nilsen, 1978), Int. Geol. Rev., 20:125–166, 1972
• E. Mutti and W. R. Normark, Comparing examples of modern and ancient turbidite systems: Problems and concepts, in J. R. Leggett and G. G. Zuffa (eds.), Marine Clastic Sedimentology: Concepts and Case Studies, pp. 1–37, Graham and Trotman, London, 1987
• G. Shanmugam et al., Basin-floor fans in the North Sea: Sequence stratigraphic models vs. sedimentary facies, Amer. Ass. Petrol. Geol. Bull., 79:477–512, 1995
• G. Shanmugam and R. J. Moiola, Submarine fans: Characteristics, models, classification, and reservoir potential, Earth Sci. Rev., 24:383–428, 1988
• G. Shanmugam, T. D. Spalding, and D. H. Rofheart, Process sedimentology and reservoir quality of deep-marine bottom-current reworked sands (sandy contourites): An example from the Gulf of Mexico, Amer. Ass. Petrol. Geol. Bull., 77:1241–1259, 1993
• P. R. Vail et al., The stratigraphic signatures of tectonics, eustacy and sedimentology—An overview, in G. Einsele, W. Ricken, and A. Seilacher (eds.), Cycles and Events in Stratigraphy, pp. 618–659, Springer-Verlag, Berlin, 1997
• P. P. E. Weaver, J. Thomson, and P. M. Hunter, Introduction, in P. P. E. Weaver and J. Thomson (eds.), Geology and Geochemistry of Abyssal Plains, Geol. Soc. London Spec. Publ., no. 31, pp. vii–xii, 1987
• Alifazeli=egeology.blogfa.com

 

ENCYCLOPEDIA ARTICLE: Deep-marine sediments

Dimensions of deep-marine facies
Example	Width:thickness ratio
Slide
Lower Carboniferous, England	7:1 (100 m wide/long, 15 m thick)
Cambrian-Ordovician, Nevada	30:1 (30 m wide/long, 1 m thick)
Jurassic, Antarctica	45:1 (20 km wide/long, 440 m thick)
Modern, U.S. Atlantic margin	40–80:1 (2–4 km wide/long, 50 m thick)
Modern, Gulf of Alaska	130:1 (15 km wide/long, 115 m thick)
Middle Pliocene, Gulf of Mexico	250:1 (150 km wide/long, 600 m thick)
Slide/slump/debris flow/turbidite
5000–8000 years B.P., Norwegian continental margin	675:1 (290 km wide/long, 430 m thick)
Slump
Cambrian-Ordovician, Nevada	10:1 (100 m wide/long, 10 m thick)
Aptian-Albian, Antarctica	10:1 (3.5 km wide/long, 350 m thick)
Slump/slide/debris flow
Lower Eocene, Gryphon Field, U.K.	20:1 (2.6 km wide/long, 120 m thick)
Paleocene, Faeroe Basin, north of Shetland Islands	28:1 (7 km wide/long, 245 m thick)
Slump
Modern, southeast Africa	170:1 (64 km wide/long, 374 m thick)
Carboniferous, England	500:1 (5 km wide/long, 10 m thick)
Lower Eocene, Spain	900–3600:1 (18 km wide/long, 5–20 m thick)
Debris flow
Modern, British Columbia	12:1 (50 m wide/long, 4 m thick)
Cambrian-Ordovician, Nevada	30:1 (300 m wide/long, 10 m thick)
Modern, U.S. Atlantic margin	500–5000:1 (10–100 km wide/long, 20 m thick)
Quaternary, Baffin Bay	1250:1 (75 km wide/long, 60 m thick)
Turbidites (depositional lobes)
Cretaceous, California	165:1 (10 km wide/long, 60 m thick)
Lower Pliocene, Italy	1200:1 (30 km wide/long, 25 m thick)
Turbidites (basin plain)
Miocene, Italy	11,400:1 (57 km wide/long, 5 m thick)
16,000 years B.P., Hatteras Abyssal Plain	125,000:1 (500 km wide/long, 4 m thick)

جزر و مد-Tidalites

Tidalites

Sediments deposited by tidal processes. Until recently, “tidalites” referred to sediments deposited by tidal processes in both the intertidal zone (between normal low- and high-tide levels) and shallow, subtidal (permanently submerged), tide-dominated environments less than 200 m (660 ft) deep. Tidalites are now known also to occur within supratidal environments (above normal high tide and flooded only during storms or very high spring tides) and submarine canyons at depths much greater than 200 m. Common usage has drifted toward describing tidalites as ripple- and dune-scale features rather than more composite deposits such as large linear sand ridges of tidal origin present on continental shelves or point bars associated with migrating tidal channels. Both of these larger-scale features, however, would be composed of tidalites.

 

Recognition criteria

 

By identifying tidalites in either the modern or the ancient geological record, geologists are implying that they know that the sediments were deposited by tidal processes rather than by storms or waves. Tidalites are not always easy to identify with certainty, especially in the rock record. In order to do so, it is necessary to understand the basic tidal cycles that can influence sedimentation.

 

Tidal theory

 

Tides are generated by the combined gravitational forces of the Moon and Sun on the Earth's oceans. Some sources are misleading in suggesting that the tidal forces from the Moon and Sun, in combination with centrifugal forces associated with the spin of the Earth, produce oceanic bulges on opposite sides of the Earth. While it is true that the combined gravitational forces of the Moon and, to a lesser extent, the Sun produce tides on the Earth, the Earth does not spin through two oceanic bulges that form on opposite sides of the Earth. This conceptual model has little bearing on real-world tides. Rather, water within each of the Earth's ocean basins is forced to rotate as discrete waves about a series of fixed (amphidromic) points (Fig. 1). For a fixed point along an ocean coastline, a tidal system is referred to as diurnal if it experiences the passing of the resultant tidal wave once every 24 h 50 min. The tidal system is semidiurnal if the resultant tidal wave passes the fixed point twice during the same time.

 

 

Fig. 1  North Sea amphidromic tidal system. Corange lines indicate equal tidal range. Cotidal lines show times of high water. Arrows show rotation directions of the tidal waves. (Modified from R. W. Dalrymple, Tidal Depositional Systems, in R. G. Walker and N. P. James, eds., Facies Models Response to Sea Level Changes, pp. 195–218, Geological Association of Canada, 1992)

 

 

 

 

 

 

In the open ocean, the motion of a tidal wave is largely expressed as a vertical movement of water masses. In shallow basins along the coast, water movements are more horizontal, with tides moving in and out of estuaries and embayments, resulting in a change in water level as the tidal wave passes. The daily or semidaily rise in tides is called the flood tide, and the fall is referred to as the ebb tide (Fig. 2a). Tidal currents are maximized between flood and ebb tides and minimized at highest flood or lowest ebb tides (Fig. 2b). The difference between the high tide and the low tide is called the tidal range.

 

 

Fig. 2  Flood–ebb cycle. Idealized (a) semidiurnal tidal cycle and (b) time–current velocity curve. (Modified from S. D. Nio and C. S. Yang, Recognition of tidally influenced facies and environments, Short Course Note Ser. 1, International Geoservices BV, Leiderdorp, Netherlands, 1989)

 

 

 

 

 

 

The intensity or height of the daily or twice-daily tides can vary in a number of ways. Cyclic semimonthly changes in daily tidal heights associated with neap–spring tidal cycles are the most pronounced of these. Spring tides occur twice a month when the tidal range is greatest, and neap tides occur twice a month when tidal range is least. Neap–spring cycles can be generated in two ways. The familiar neap–spring cycle is related to the phase changes of the Moon. Spring tides occur every 14.76 days when the Earth, Moon, and Sun are nearly aligned at new or full moon (Fig. 3a). Neap tides occur when the Sun and Moon are aligned at right angles from the Earth at first- and third-quarter phases of the Moon. The result is that spring tides are higher than neap tides (Fig. 3b). The time from new moon to new moon is called the synodic month, which has a modern period of 29.53 days. This type of neap–spring cycle is referred to as synodically driven, and it dominates the coastlines of western Europe and the eastern coastline of North America.

 

 

Fig. 3  Idealized models of origin of neap–spring tidal cycles: (a) Synodic month. (b) A segment of the 1991 predicted high tides from Kwajalein Atoll, Pacific. (c) Tropical month. (d) A segment of the 1994 predicted high tides from Barito River estuary, Borneo. (Modified from E. P. Kvale, K. H. Sowder, and B. T. Hill, Modern and ancient tides: Poster and explanatory notes, Society for Sedimentary Geology, Tulsa, OK, and Indiana Geological Survey, Bloomington, IN, 1998)

 

 

 

 

 بزرگنمایی تصویر

 

A second type of neap–spring cycle is less well known but no less common, and is related to the orbit of the Moon around the Earth. The Moon's orbital plane is inclined relative to the Earth's equatorial plane. The period of the variation in lunar declination relative to the Earth's Equator is called the tropical month, and is the time the Moon takes to complete one orbit, moving from its maximum northerly declination to its maximum southerly declination and return (Fig. 3c). In this type of neap–spring cycle, the tidal force depends on the position of the Moon relative to the Earth's Equator. The tide-raising force at a given location is greater when the Moon is at its maximum declination every 13.66 days. These periods correspond to the generation of spring tides (Fig. 3d). The neap tides occur when the Moon is over the Earth's Equator. The current length of the tropical month is 27.32 days, and neap–spring cycles in phase with the tropical month are referred to as tropically driven. These types of neap-spring cycles dominate coastlines in the Gulf of Mexico and many areas in the Pacific.

Besides generating neap–spring cycles in many parts of the world, the changing position of the Moon relative to the Earth's Equator through the tropical month causes the diurnal inequality of the tides in semidiurnal tidal systems. In tidal systems that experience two high tides and two low tides per day, the tropical monthly cycle results in the morning high tide being greater or lesser than the evening high tide. The diurnal inequality is reduced to zero when the Moon is over the Equator, resulting in the morning tide and the evening tide being of equal magnitude (Fig. 3b and d).

Other tidal cycles besides those mentioned above can influence sedimentation and have been documented in the geologic record. These include monthly, semiannual, and multiyear tidal cycles.  See also: Earth rotation and orbital motion; Moon; Tide

 

Examples of tidalites

 

How the various tidal cycles manifest themselves in the geologic record and how geologists can identify their influence on sedimentation has been studied for nearly 75 years. To recognize tidalites in the geologic record, geologists look for evidence of one or more of the following:

1. Sediment deposited by reversing currents (that is, flood–ebb cycles).

2. A stacked sequence of sediments that show a recurring change from sediments transported (and deposited) by currents at maximum current velocity to sediments deposited from suspension at minimum current velocity (Fig. 2b).

3. Stacked packages of sediments in which each package shows evidence of subaerial exposure superimposed on sediments deposited in subaqueous settings (sediments transported and deposited during flood tides and exposed during low ebb tide).

4. A sequence of sediment packages in which the thickness or accretion of successive packages of sediments varies in a systematic way, suggesting diurnal, semidiurnal, and/or neap–spring tidal cycles.

An example of a small-scale tidalite can be found in the Mansfield Formation (Pennsylvanian Period) in Orange County, Indiana (Fig. 4). The sample shown is from a rock core taken through this interval. The lighter-colored layers are siltstone, and the thin dark layers are finer-grained claystone. The regular and repeating change in deposition from siltstone to claystone indicates systematic current velocity fluctuations related to the tidal cycle over a 12-h period (see item 2 above). The thick–thin pairing of the lighter bands of siltstone suggests the influence of the semi-diurnal inequality over a 24-h period (see item 4). In addition, the regular and systematic overall thickening and thinning of the siltstone layers, as shown in the bar chart next to the core in Fig. 4, suggests that neap–spring tidal cycles controlled the thicknesses of the silt layers. The higher spring tides resulted in thicker accumulations of silt than the lower neap tides.  See also: Pennsylvanian; Sedimentary rocks

 

 

Fig. 4  The core shows small-scale tidalites from the Hindostan Whetstone beds, Mansfield Formation, Indiana. The chart shows thicknesses of layers as measured between dark clay-rich bands. The interval shows approximately one synodic month of deposition.

 

 

 

 

 

 

An example of a large-scale tidalite can be found in the Jurassic Sundance Formation of northern Wyoming (Fig. 5). This tidalite is the remnant of a migrating subtidal dune or sandwave. The preserved inclined beds of the avalanche face (I) indicate the migration direction of the dune from right to left (Fig. 5a). The evidence for tidal influence, however, lies within the inclined, less resistant, and more recessed lighter-colored bands (examples marked by arrows in Fig. 5a). In this interval (Fig. 5b), one sees evidence of (1) cessation of dune migration and a reversal of current direction from flood tide to ebb tide with small ripples migrating up the avalanche face (II); (2) a mud drape (III) resulting from fine-grained materials settling out of suspension when the current velocities reached zero as the tide reversed (Fig. 2); and (3) a reactivation of the migrating dune above the mud drape (IV) as current velocity increased during the next flood tide. The right-to-left migration of the dune was also controlled by the neap–spring cycle, with greater migration (interval between lighter-colored bands) occurring during spring tides and lesser migration during neap tides (Fig. 5a). In the example shown, the neap tide deposits are centered on line N.  See also: Jurassic

 

 

Fig. 5  Photographs from Sundance Formation of northern Wyoming. (a) Example of large-scale tidalites. (b) Closeup of inclined light-colored band showing evidence of current reversals.

 

 

 

 

 

 

 

 

Geologic record

 

Deposits of tidalites are known from every geologic period from the modern back into the Precambrian and from depositional environments with water chemistries ranging from fresh to hypersaline. Studies of tidalites are important because geologists have used these features not only to interpret the original depositional settings of the deposits but also to calculate ancient Earth–Moon distances, interpret paleoclimates existent during deposition, and calculate sedimentation rates.  See also: Depositional systems and environments; Geologic time scale; Marine sediments; Sedimentology

Erik P. Kvale

 

Bibliography

 

 

• C. Alexander, R. Davis, and V. Henry (eds.), Tidalites: Processes and Products, Geological Society Publishing, 1998
• D. E. Cartwright, Tides: A Scientific History, Cambridge University Press, 1998
• G. deV. Klein, A sedimentary model for determining paleotidal range, Geol. Soc. Amer. Bull., 82:2585–2592, 1971
• G. deV. Klein, Determination of paleotidal range in clastic sedimentary rocks, XXIV International Geological Congress, 6:397–405, 1972
• D. T. Pugh, Tides, Surges and Mean Sea Level, Wiley, 1987
• H. G. Reading (ed.), Sedimentary Environments: Processes, Facies and Stratigraphy, 3d ed., Blackwell Science, Cambridge, MA, 1996

 

Additional Readings

  

• J. R. L. Allen, Lower Cretaceous tides revealed by cross-bedding with mud drapes, Nature, 289:579–581, 1981
• J. R. Boersma and J. H. J. Terwindt, Neap–spring tide sequence of intertidal shoal deposits in a mesotidal estuary, Sedimentology, 28:151–170, 1981
• R. W. Dalrymple and Y. Makino, Description and genesis of tidal bedding in Cobequid Bay–Salmon River estuary, Bay of Fundy, Canada, in A. Taira and F. Masuda (eds.), Sedimentary Facies of the Active Plate Margin, pp. 151–177, Terra Publishing, Tokyo, 1989
• H. R. Feldman et al., Stratigraphic architecture of the Tonganoxie paleovalley fill (Lower Virgilian) in northeastern Kansas, Amer. Ass. Petrol. Geol. Bulle., 79:1019–1043, 1995
• E. P. Kvale et al., Calculating lunar retreat rates using tidal rhythmites, J. Sediment. Res., 69:1154–1168, 1999
• E. P. Kvale et al., Evidence of seasonal precipitation in Pennsylvanian sediments of the Illinois Basin, Geology, 22:331–334, 1994
• G. Shanmugam, Deep-marine tidal bottom currents and their reworked sands in modern and ancient submarine canyons, Mar. Petrol. Geol., 20:471–491, 2003
• B. Tessier, Upper intertidal rhythmites in the Mont-Saint-Michel Bay (NW France): Perspectives for paleoreconstruction, Mar. Geol., 110:355–367, 1993
• NOAA Tides and Currents
•  Alifazeli=egeology.blogfa.com

 

رسوبات رودخانه ای-Fluvial sediments

Fluvial sediments

Deposits formed by rivers. An alluvial river is one which flows within its own fluvial sediments, as distinct from one that has incised into the underlying bedrock. A river accumulates deposits because its capacity to carry sediment has been exceeded, and some of the sediment load is deposited. Rivers tend toward a state of dynamic equilibrium, in which they adjust their slope in response to changes in discharge and sediment load. The result is a channel profile that is steep in its source areas but flattens out downstream, and is graded to a slope of zero where the river discharges into a lake or the sea. Fluvial sedimentary accumulations range from temporary bars deposited on the insides of meander bends as a result of a loss of transport energy within a local eddy, to deposits tens to hundreds of meters thick formed within major valleys or on coastal plains as a result of the response of rivers to a long-term rise in base level or to the uplift of sediment source areas relative to the alluvial plain. Both these processes perturb the graded profile so that it tends to rise, creating space, or accommodation, for sediment. The same processes control the style of rivers and the range of deposits that are formed, so that a study of the deposits may enable the geologist to reconstruct the changes in controlling factors during the accumulation of the deposits.  See also: Depositional systems and environments; River; Stream transport and deposition

Coarse debris generated by mechanical weathering, including boulders, pebbles, and sand, is rolled or bounced along the river bed and is called bedload. The larger particles may be moved only infrequently during major floods. Finer material, of silt and clay grade, is transported as a suspended load, and there may also be a dissolved load generated by chemical weathering. Mass movement of large volumes of sediment by sediment gravity flows, typically debris flows, may occur when rare flash floods mobilize debris that may have been accumulating in source areas for some time. Whereas the volume of sediment tends to increase downstream within a drainage system, as tributaries run together, the grain size generally decreases as a result of abrasion and selective transport. This downstream grain-size decrease may assist in the reconstruction of transport directions in ancient deposits where other evidence of paleogeography has been obscured by erosion or tectonic change.  See also: Mass wasting

 

Types of river

 

River type may be described by two main variables, sinuosity and channel multiplicity. These variables combine to form four end-member styles, discussed below, although there are many examples of rivers showing various styles intermediate between these end members.

 

Braided rivers

 

These typically occur in areas of high sediment load and variable discharge. They consist of several or many branching, unstable channels of low sinuosity, and are characterized by abundant coarse bedload, forming bars, islands, and channel-floor deposits. The channel complex typically occupies most of the valley floor, leaving little room for a floodplain. Glacial outwash streams and ephemeral streams draining mountainous areas in arid regions are normally braided, and may form broad sheets of sand or gravel crossed by networks of shallow, shifting channels.

 

Meandering rivers

 

These are single-channel streams of high sinuosity, in which islands and midchannel bars are rare. Sediment in these rivers range from very coarse to very fine. A significant proportion of the bedload typically is deposited on the insides of meander bends, forming point bars. The channel, with its coarse deposits, may be confined to a narrow belt within an alluvial valley, flanked by a broad floodplain, upon which deposition of fine-grained sediment takes place only during flood events—seasonally or at longer intervals.

 

Anastomosed rivers

 

These develop in stable, low-energy environments or in areas undergoing rapid aggradation. They consist of a network of relatively stable, low- to high-sinuosity channels bounded by well-developed floodplains. Channels are characteristically narrow and accumulate narrow, ribbonlike sandstone bodies.

 

Straight channels

 

These are rare, occurring mainly as distributaries in some deltas.

Rivers which emerge from a mountainous catchment area into a low plain drop their sediment load rapidly. The channel may bifurcate, becoming braided in character. A distinctive landform, an alluvial fan, results.

 

 

Sedimentary facies

 

River deposits of sediment occur as four main types.

 

Channel-floor sediments

 

The coarsest bedload is transported at the base of the channel, commonly resulting in deposits of gravel (Fig. 1a), waterlogged vegetation, or fragments of caved bank material. In sand-bed rivers the channel floor commonly is covered by fields of large, sinuous-crested dunes or ripples (with amplitudes of 2 in. to 10 ft, or 5 cm to 3 m), which impart a trough–cross-bedded structure to the sand.

 

 

Fig. 1  Typical fluvial deposits. (a) Gravel and sand channel-fill and bar deposits exposed in a gravel quarry face about 40 ft (12 m) high, fluvioglacial outwash, Alberta, Canada. (b) Point bar, 13 ft (4 m) thick, Carboniferous, Alabama. (c) Typical floodplain deposits, Triassic, Arizona; outcrop is about 33 ft (10 m) high.

 

 

 

 

 

 

 

Bar sediments

 

Accumulations of gravel, sand, or silt occur along river banks and are deposited within channels, forming bars that may be of temporary duration, or may last for many years, eventually becoming vegetated and semipermanent. Bars attached to one of the channel banks are termed side or lateral bars. Those occurring on the insides of meander bends are termed point bars. They develop by lateral accretion as the meander widens or shifts in position by erosion on the outer bank of the bend (Fig. 1b and Fig. 2). Bars occurring within channels accrete by the addition of sediment on all sides, but most commonly preferentially on one side representing the inside of a bend in the adjacent channel, or at the downstream end of the bar. Such bars commonly have complex internal structures, reflecting many seasons of growth and intervals of erosion.

 

 

Fig. 2  Development of a fining-upward succession by lateral accretion of a point bar, such as that in Fig. 1b.

 

 

 

 

 

 

 

Channel-top and bar-top sediments

 

These are typically composed of fine-grained sand and silt, and are formed in the shallow-water regions on top of bars, in the shallows at the edges of channels, and in abandoned channels. Small-scale ripples, with amplitudes of less than 2 in. (5 cm), are typical sedimentary structures, together with roots and bioturbation structures.

 

Floodplain deposits

 

These are formed when the water level rises above the confines of the channel and overflows the banks (Fig. 1c). Much of the coarser sediment is deposited close to the channel, in the form of levees. Breaks in the channel bank, termed crevasses, permit the transportation of additional coarse sediment onto the floodplain, where it forms small deltalike bodies spreading out into the floodplain, termed crevasse splays. Much silt and mud may be carried considerable distances from the channel, forming blanketlike deposits. In swampy areas, floodplains may be the site of thick vegetation, which in time may be transformed into lignite and eventually into coal. Soils develop in response to weathering activity and plant growth, and may form distinctive brightly colored layers, termed paleosols. Nodular beds of calcium carbonate are a common component of paleosols, especially in relatively arid areas, where they form as a result of the evaporation of ground waters.  See also: Paleosol

 

 

Facies associations and sedimentary cycles

 

Fluvial sediments may be dominantly conglomeratic, sandy, or silty, depending on the nature of the sediment load of the river. This characteristic is mainly a function of slope and the proximity to sources, but is also a reflection of sediment availability. Certain source materials, such as fine-grained sediments or limestones, may yield coarse debris on erosion, but it is likely to be broken down into fine material or dissolved on prolonged transportation. Humid climates favor chemical and biochemical weathering processes, which yield a large suspended or dissolved sediment load. Coarse detritus is more typically the product of drier climates, in which mechanical weathering processes (such as frost shattering) are dominant.

The sedimentary facies described above were listed in approximate vertical spatial order, from channel floor to floodplain. This order is one of decreasing grain size upward, a feature which may commonly be observed in ancient fluvial deposits (Fig. 2). Such deposits may consist of a series of fining-upward successions, or cycles, each a few meters to few tens of meters in thickness.

There are a variety of causes of such cycles. The first that was recognized is the mechanism of lateral accretion, whereby point bars enlarge themselves in a horizontal direction as the meander bounding them migrates by undercutting the bank on the outside of the bend (Fig. 2). The depositional surface of the point bar may be preserved as a form of large-scale, low-angle cross-bedding within the deposit, its amplitude corresponding approximately to the depth of the channel. Similar cycles may be caused by the nucleation and growth of large compound bars or sand flats within braided channel systems. The accretion surfaces, in such cases, may dip in across- or down-channel directions. Individual flood events, especially on the sand flats of ephemeral stream systems, may form sheetlike flood cycle deposits up to a meter or so thick, the upward fining corresponding to decreasing energy levels as the flood waned. The gradual choking of a channel with sediment, and the progressive abandonment of the channel, will also generate a fining-upward cycle.

 

Regional controls

 

Tectonic activity in a fluvial catchment area may cause the generation of sedimentary cycles, which either fine or coarsen upward, depending on whether relief and slope are decreased or increased, respectively. Such cycles tend to be tens to hundreds of meters thick and to extend for several to many kilometers. They may have smaller cycles formed by channel fill and migration processes nested within them.

River systems are also affected by changes in base level, that is, by a rise or fall in the level of the lake or sea into which the river drains. A fall in base level may lead to widespread incision of channels along a coastal plain as they adjust to a lower river mouth. Between channels, sedimentation may cease, with the formation of widespread, well-developed paleosols. The same effect is brought about by peneplanation, that is, long-continued subaerial erosion in the absence of tectonic rejuvenation of the river system.

A rise in base level may flood the mouths of rivers, forming estuaries. However, if the sediment supply is adequate, sedimentation may be able to keep pace with such a base-level rise, with the river changing in style in response to changes in the balance between the rate of sediment input and the rate of base-level change (Fig. 3). These sedimentary responses to external forcing are part of a larger story concerning the regional and global controls of sedimentation. The regional erosion surface formed at a time of falling to low base level or deep peneplanation is termed a sequence boundary. It may be cut by incised channels that are, in turn, typically filled by coarse channel sediments (in sequence stratigraphic terminology, these deposits are classified as the lowstand systems tract). At the coast a rise in base level is commonly recorded by transgression (the transgressive systems tract). The coastal rivers at this time may be of anastomosed style. As base-level rise reaches its highest level, the rate of rise slows, and the rivers typically evolve into a meandering type (highstand systems tract). These changes may be reflected in the resulting sediments by changes in the geometry and spacing of channel sandstone and conglomerate bodies. Lowstand deposits commonly consist of coarse, laterally amalgamated channel deposits. A rapid rise in base level (the transgressive systems tract) may be marked by isolated channel sands spaced out within thick floodplain units. The spacing of such channel bodies becomes closer, and more units are amalgamated, in the highstand deposits, above.  See also: Floodplain; Sequence stratigraphy

 

 

Fig. 3  Response of river systems and their deposits to a cycle of fall and rise of base level. SB = sequence boundary. (Modified from K. W. Shanley and P. J. McCabe, Perspectives on the Sequence Stratigraphy of Continental Strata, American Association of Petroleum Geologists, vol. 78:560, 1994)

 

 

 

 

 

 

Far inland, base-level changes may not markedly affect the rivers unless they persist for very long periods of time. Changes in discharge and sediment yield in response to climate change are commonly more important in controlling river style and the resulting sediment types. For example, during the Pleistocene glaciation, rapid deposition of coarse sediments occurred at the margins of continental ice caps. Channel erosion, forming widespread surfaces of incision, occurred at times of change, between glacial and interglacial periods, because at these times river discharge tended to increase whereas sediment yield did not. Climatically driven erosion and deposition inland were therefore out of phase with the cycle of change generated by base-level change at the coast.

More than one forcing function, including climate change, base-level change, and tectonism, may be operating at any one time, resulting in complex patterns of cyclicity. The resulting sequences may be widespread. Reconstruction of this sequence stratigraphy may provide an essential mapping tool for those engaged in basinal exploration.

 

Tectonic setting of fluvial deposits

 

The thickest (up to 6 mi or 10 km) and most extensive fluvial deposits occur in convergent plate-tectonic settings, including regions of plate collision, because this is where the highest surface relief and consequently the most energetic rivers and most abundant debris are present. Some of the most important accumulations occur in foreland basins, which are formed where the continental margin is depressed by the mass of thickened crust formed by convergent tectonism. Examples include the modern Himalayan foreland basin of the Indus and Ganges valleys, the Devonian foreland basin west of the Appalachian Mountains, and the late Cenozoic foreland basin of France and Germany, north of the alpine mountain chain.  See also: Basin

Thick fluvial deposits also occur in rift basins, where continents are undergoing stretching and separation. The famous hominid-bearing sediments of Olduvai Gorge and Lake Rudolf are fluvial and lacustrine deposits formed in the East Africa Rift System. Triassic fault-bounded basins along the North American Atlantic coast and through western Europe are an older but comparable example. Fluvial deposits are also common in wrench-fault basins, such as those in California.

 

Economic importance

 

Significant volumes of oil and gas are trapped in fluvial sandstones. Major reservoirs include those of Triassic-Jurassic age in the North Sea Basin; Triassic sandstones of the Paris Basin; Permian-Triassic sandstones of Prudhoe Bay on the Alaskan North Slope; the Lower Cretaceous reservoirs of the giant Daqing field of the Songliao Basin, China; Jurassic sandstones of interior Australia; the heavy-oil sands of the Cretaceous Athabasca and related deposits in Alberta; and numerous large to small fields in mature areas such as the Alberta Basin (Cretaceous), the southern midcontinent (Pennsylvanian), and the Gulf Coast (Cretaceous and Oligocene fields).

Placer gold, uranium, and diamond deposits of considerable economic importance occur in the ancient rock record in South Africa and Ontario, Canada, and in Quaternary deposits in California and Yukon Territory. Economically significant roll-front uranium deposits occur in the Mesozoic deposits of the American Western Interior and elsewhere, primarily in fluvial facies.

Fluvial deposits are also essential aquifers, especially the postglacial valley-fill complexes of urban Europe and North America. Much work needs to be done to investigate the internal geometry of these deposits in order to resolve problems of domestic and industrial pollution that now interfere with the use of the ground water from these sources.

Andrew D. Miall

 

Bibliography

 

 

• P. A. Carling and M. R. Dawson (eds.), Advances in Fluvial Dynamics and Stratigraphy, Wiley, 1996
• C. R. Fielding (ed.), Current research in fluvial sedimentology, Sed. Geol., special issue, vol. 85, 1993
• A. D. Miall, The Geology of Fluvial Deposits, Springer-Verlag, 1996
•  Alifazeli=egeology.blogfa.com

 

گسل و ساختارهای گسلی-Fault and fault structures

Fault and fault structures

Products of fracturing and differential movements along fractures in continental and oceanic crustal rocks. Faults range in length and magnitude of displacement from small structures visible in hand specimens, displaying offsets of a centimeter (1 cm = 0.4 in.) or less, to long, continuous crustal breaks, extending hundreds of kilometers (1 km = 0.6 mi) in length and accommodating displacements of tens or hundreds of kilometers. Faults exist in deformed rocks at the microscopic scale, but these are generally ignored or go unrecognized in most geological studies. Alternatively, where microfaults systematically pervade rock bodies as sets of very closely spaced subparallel, planar fractures, they are recognized and interpreted as a type of cleavage which permitted flow of the rock body. Fractures along which there is no visible displacement are known as joints. These include shear joints, formed by fracturing and imperceptible movement of the walls of the fractures parallel to fracture surfaces, and tension joints, formed by negligible or barely visible displacement of the walls of the fractures perpendicular to the fracture surfaces. Large fractures which have accommodated major dilational openings (a meter or more) perpendicular to the fracture surfaces are known as fissures. Formation of fissures is restricted to near-surface conditions, for example, in areas of crustal stretching of subsidence. When faulting takes place under conditions of high temperature or pressure, zones of penetrative shear flow may develop which are best described as ductile fault zones.

 

Locating faults

 

The recognition of faults in continental regions of moderate to excellent rock exposure is generally straightforward. Classically, the process of systematic geological mapping has proved to be a powerful method for locating faults. In essence, faults can be identified and tracked by recognizing in mapped patterns the truncation and offset of one or several bedrock units. Depending on the nature of the exposure and movement on the fault, truncation and offset might produce a simple horizontal shifting of the dominant mapped pattern of bedrock units; alternatively, the faulting might lead to a pattern of repetition or omission of specific rock formations within the geologic column.

Valuable physical signatures which reveal the presence of faults include abnormally straight top­ographic lineaments or fault-line scarps; aligned springs issuing from fractured and favorably displaced bedrock; intensely fractured rocks, perhaps with zones of angular chunks of brecciated, rotated materials; fracture surfaces naturally polished through the movement process and etched with striations or grooves; dragging (folding) of rock layers out of their normal orientation; loosely consolidated, ground-up rock flour or paste, commonly referred to as gouge; radically crushed, cataclastically deformed rocks known as mylonites; and alteration, silicification, or mineralization brought about by circulation of hot fluids through shattered bedrock. Faults are simple to locate in areas of active or very recent mountain building, especially where faulting has broken the ground surface and produced scarps. See also: Mylonite

 

 

Fig. 1  Transform faults linking oceanic ridge segments. (After W. J. Morgan, Rises, trenches, great faults, and crustal blocks, J. Geophys. Res., 73:1959–1982, 1968)

 

 

 

بزرگنمایی تصویر

 

 

 

In continental areas of very poor rock exposure, in the subsurface, and in ocean basins, faults are much more difficult and costly to locate. However, major faults are routinely discovered through application of geophysical methods, especially seismic, gravity, and magnetic surveying. Abrupt contrasts in the geophysical signatures of rocks at depth signal the sharp truncation of bedrock by faults and allow the pattern of faults to be mapped. The geophysical exploration of fault and fault structures in the ocean floor has completely changed the way in which geologists view the Earth and earth dynamics. It is known that major fracture zones, unlike any recognized in continental regions, exist in the ocean floor. These fractures, hundreds to thousands of kilometers in length and spaced at tens of kilometers, pervade the Mid-Oceanic Ridge system and are usually oriented perpendicular to the crest of the ridge segment which they occupy. Interpreting these to be enormous faults which accommodate the movement of newborn oceanic crust as it spreads bilaterally from the Mid-Oceanic Ridge system, J. T. Wilson named them transform faults. See also: Geophysical exploration; Mid-Oceanic Ridge; Marine geology

 

Transform faults

 

Three fundamental types of transform faults exist: the first connects one ridge segment to another; the second connects a ridge segment (where new oceanic crust forms) to a trench (where oceanic crust is consumed through subduction); the third connects two trenches. Ridge-ridge transform faults, perhaps the easiest to visualize (Fig. 1), link parallel but offset ridge segments, and are characterized by shallow earthquake activity restricted to the part of the transform between the two ocean-ridge segments. The sense of movement along each ridge-ridge transform fault, as deduced from fault-plane solutions, is opposite to that which would explain the offset of the oceanic ridge. In essence, the movement along the transform fault does not produce the offset, but rather records the differential sliding-past (shearing) of the new ocean floor moving in opposite directions from the two ocean-ridge segments which are connected by the transform. Thus the transform faults serve, in this example, as part of the divergent plate boundary which accommodates sea-floor spreading and the movement of the two plates can be described by an imaginary pole of rotation located on the Earth's surface at a position defined by the intersection of great circles drawn perpendicular to points on the array of transform faults. The rate of relative movement along the ridge-ridge transform is a function of the rate of generation of new oceanic crust along the ridge and is generally on the order of several centimeters per year. See also: Mountain systems; Plate tectonics; Transform fault

Transform fault patterns are quite complicated, as is their evolution through time. Boundary zones between three adjacent tectonic plates (or triple junctions) are particularly difficult to evaluate with respect to geometry and kinematics (motions). Nonetheless, plate-tectonic analysis, including detailed assessment of transform faults, has had a revolutionary impact on understanding of the Cenozoic tectonic evolution of the Earth's crust. For example, scientists generally accept Tanya Atwater's hypothesis that the infamous San Andreas fault system of California is a transform fault boundary separating two enormous crustal plates, the North American and the Pacific. Relative movement between these plates is horizontal and right-handed, and in magnitude amounts to hundreds of kilometers.

 

Movements

 

This interpretation of the San Andreas Fault clearly demonstrates that some major fault systems in continental regions can be better understood in the context of the “new rules” afforded by study of transform faulting and plate tectonics. However, most fault systems in continental regions cannot be clearly and quantitatively linked to specific plate-tectonic movements or configurations, present or past, and they are analyzed and understood in an entirely different way; the guidelines for analysis are well established. In addition to describing the physical and geometric nature of faults and interpreting time of formation, it has been found to be especially important to determine the orientations of minor fault structures (such as striae and drag folds) which record the sense of relative movement.

 

 

Fig. 2  Slip on faults. (a) Block before faulting; (b) normal-slip; (c) reverse -slip; (d) strike-slip; (e) oblique-slip. (After F. Press and R. Siever, Earth, 2d ed., W. H. Freeman, 1978)

 

 

 

 

 

 

Evaluating the movement of faulting can be difficult, for the apparent relative movement (separation) of fault blocks as seen in map or outcrop may bear little or no relation to the actual relative movement (slip). The slip of the fault is the actual relative movement between two points or two markers in the rock that were coincident before faulting (Fig. 2). Strike-slip faults have resulted in horizontal movements between adjacent blocks; dip-slip faults are marked by translations directly up or down the dip of the fault surface; in oblique-slip faults the path of actual relative movement is inclined somewhere between horizontal and dip slip. Strike-slip faults are described as left- or right-handed, depending on the sense of actual relative movement; dip-slip faults are described as normal-slip, thrust-slip, or reverse-slip, depending on the sense of actual relative movement and on the dip and dip direction of the fault surface. Listric normal- slip faults are a type of normal-slip fault in which the inclination of the fault decreases. Movement on such a curved fault surface produces a profound backward rotation of the upper block.

Recognizing even the simplest translational fault movements in nature is often enormously difficult because of complicated and deceptive patterns created by the interference of structure and topography, and by the absence of specific fault structures which define the slip path (Fig. 3). While mapping, the geologist mainly documents apparent relative movement (separation) along a fault, based on what is observed in plan-view or cross-sectional exposures. In the separation sense, left- and right-lateral faults are those displaying apparent horizontal shifts of rock in map view. Such shifts are said to be apparent because the left- or right-lateral offset might actually have been produced by dip-slip, not strike-slip, movements. Normal, reverse, and thrust are separation terms for faults, and again the usage of each is based on apparent offset in cross-sectional view and on the dip and dip direction of the fault.

 

 

Fig. 3  Slip versus separation. (a) AB is the dip-slip. (b) After erosion of top of footwall block. The block has undergone a right separation. (After M. P. Billings, Structural Geology, 3d ed., Prentice-Hall, 1972)

 

 

 

 

 

 

Movement and offset on large, regional fault systems must be evaluated on the basis of displaced geologic terrains and abnormal stratigraphic relations. For example, 435 to 500 mi (700 to 800 km) of left-slip fault movement has been postulated in northern Mexico during the time period 150,000,000 to 170,000,000 years ago. The basis of the interpretation is truncation and offset of terrains of Precambrian and Paleozoic rocks in California and Arizona. Low-angle thrust movements in western Utah produced in Cretaceous time a 44-mi (70-km) west-to-east transport of thick Precambrian through Paleozoic miogeoclinal strata onto thin shelf and platform strata of the same age.

 

Stress conditions

 

Theory on faulting is based on applied physics and engineering, and focuses on the stress conditions under which rocks break. The theory is almost exclusively concerned with the brittle behavior of crustal rocks; as such, it is most applicable to faulting at upper crustal layers in the Earth. Results of deformational experiments under controlled temperature, pressure, and strain-rate conditions bear importantly on modern understanding of the dynamics of faulting. The results of theoretical and experimental work provide insight into why faults can be conveniently separated into categories of normal-slip, thrust-slip, and strike-slip.

 

 

Fig. 4  Mohr circle diagram constructed from three sets of compressional tests of limestone. Angles (61°, 65°, and 71°) are 2θ values. 1 psi = 6.89 kPa.

 

 

 

 

 

 

Fig. 5  Cross section of listric normal faults in the Basin and Range Province. (After R. E. Anderson, Geologic Map of the Black Canyon 15-Minute Quadrangle, Mohave County, Arizona, and Clark County, Nevada, USGS Map GQ-1394, 1978)

 

 

 

 

 

 

Forces which act on a rock body may be resolved by vector analysis into components of force acting in specific directions. These, in turn, can be converted to magnitudes and directions of stresses which tend to deform the body. This is done by dividing the force component by the surface area (of the body) on which it acts. Two types of stresses are distinguished, normal stress (σ), which acts perpendicular to a given surface, and shear stress (τ), which acts parallel to the given surface. Stress analysis, in effect, evaluates the magnitude of shear and normal stresses acting in all directions throughout a body and predicts the orientations of the surfaces along which faulting should occur.

Evaluating the distribution of stresses in a body that is acted upon by forces discloses that there are three unique directions within the stress field, called principal stress directions. These stress directions are mutually perpendicular, and the value of shear stress equals zero only along these three directions. Furthermore, one of the principal stress directions is characterized by the maximum value of normal stress (σ1) within the system, and another (σ3) is characterized by the minimum normal stress. Maximum shear stress values (τmax) occur along lines oriented 45° to the principal stress directions.

Two-dimensional mathematical analysis demonstrates that the magnitude of normal stress (σ) on any plane (that is, on any potential surface of faulting) in the body is given by Eq. (1), where σ =

 

 

 (1) 

greatest principal normal stress, σ3 = least principal normal stress and θ = angle between the greatest principal stress axis and the direction of shearing stress (in the plane for which σ is being evaluated). The magnitude of shear stress (τ) on that same plane is given by Eq. (2). The distribution of

 

 

 (2) 

paired values of normal stress and shear stress as a function of θ is such that shear stress (τ) is zero at values of 0°, 90°, 180°, and 270°, but attains maximum values at 45°, 135°, 225°, and 310°. Normal stress reaches a maximum at 90° and 270°, but is minimal at 0° and 180°. These variations in normal and shear stress values may be portrayed on a Mohr circle diagram, a graphical representation of the above equations (Fig. 4). Points on the periphery of the circle have coordinates (σ,τ) which correspond in value to normal stress and shear stress on a plane which makes an angle of θ with σ1, the greatest principal stress direction.

Given such a stress distribution, and assuming that the differential stress conditions (σ1 − σ3) exceed the strength of the rock body, the orientation of faulting can be determined. The Mohr-Coulomb law of failure [Eq. (3)]

 

 

 (3) 

predicts that faulting should occur at a critical shear stress level (τc) where τ0 = cohesive strength of the rock, σ = normal stress on the fault plane, and φ = angle of internal friction. The coefficient of internal friction, tan φ, equals σ/τ at failure. For most rocks, the coefficient of internal friction has a value between 0.4 and 0.7; thus the angle of internal friction, φ, commonly varies from 20 to 35°. The value of θ for such internal friction typically ranges from 27 to 35°, and is often 30°. In practice, the failure points on Mohr circle diagrams, as generated through experimental deformation, do not conform to the ideal straight-line failure envelope predicted on the basis of the Mohr-Coulomb law. Rather, the failure envelopes are smoothly curved in a way that describes an increase in 2θ with an increase in confining pressure (σ3). Predictions as to when and at what angle faulting should occur are further complicated by variables of temperature, strain rate, pore-fluid pressure, and presence of fractures. In fact, fundamental questions have been raised regarding the extent to which theory and short-term experimental work can be applied to some natural geological systems.

What arises from theory and experiments is that fractures form in an orientation such that they contain the axis of intermediate stress (σ2) and make an angle of θ (commonly around 30°) with σ1 and 90 − θ with σ3. Since the Earth's surface has no shear stress, principal stress directions near the Earth's surface tend to be vertical and horizontal, and depending on the relative configuration of the principal stresses, will give rise to thrust-slip, normal-slip, or strike-slip faults. The direction of movement on these faults is such that the wedge receiving the greatest compressive stress moves inward, whereas the wedge receiving the least compressive stress moves outward.

 

Examples

 

There are many excellent natural examples of normal-slip, thrust-slip, and strike-slip faults. The Basin and Range Province of the western United States displays a unique physiographic basin/range style because of pervasive large-scale normal-slip faulting (Fig. 5), which resulted from regional crustal extension within the last 15,000,000 years. Normal-slip faults accommodate extension.

Thrust-slip faults are an integral part of the tectonic framework of the southern Appalachian Mountains, the Sevier orogenic belt of western Utah and western Wyoming/eastern Idaho, and the Canadian Rockies. In mountain belts throughout the world, thrusting has played a major role in accommodating crustal shortening during mountain building. Major mechanical questions have arisen as to how enormous masses of crustal rocks can be thrust tens or hundreds of kilometers, since the apparent force required should have crushed the rock mass before it moved. The paradox of regional overthrusting has led to theories of faulting that emphasize the importance of factors such as high pore-fluid pressure, gravitational sliding, viscous creep of ductile materials, and underthrusting.

The best-documented strike-slip faulting is concentrated near margins of lithospheric plates, but such faulting has occurred in foreland tectonic regions as well, including the Rocky Mountains of the American West and mainland China. A major theme that has emerged from the study of high-angle faults, especially strike-slip faults, is that ancient faults in basement rocks are commonly reactivated in post-Precambrian time, producing zones of concentrated, superposed strain. This interpretation has been used to explain the classic monoclinal uplifts of the Colorado Plateau. See also: Graben; Horst; Structural geology

George H. Davis

 

Bibliography

 

 

• M. P. Billings, Structural Geology, 4th ed., 1986
• J. G. Dennis, Structural Geology: An Introduction, 1987
• R. E. Powell and J. C. Matti (eds.), San Andreas Fault System: Displacement, Palinspastic Reconstruction, and Geologic Evolution, 1993
• F. Press and R. Siever, Understanding Earth, 3d ed., 2000
• R. J. Twiss and E. M. Moores, Structural Geology, 1992
• J. T. Wilson, A new class of faults and their bearing on continental drift, Nature, 207:343–347, 1965
• Alifazeli=egeology.blogfa.com

تحلیل گسل-Fault analysis

Fault analysis

The detection and diagnosis of malfunctions in technical systems. Such systems include production equipment (chemical plants, steel mills, paper mills, and power stations), transportation vehicles (ships, airplanes, automobiles), and household appliances (washing machines, air conditioners). In any of these systems, malfunctions of components may lead to damage of the equipment itself, degradation of its function or product, jeopardy of its mission, and hazard to human life. While the need to detect and diagnose malfunctions is not new, advanced fault detection has been made possible only by the proliferation of the computer. Fault detection and diagnosis actually means a scheme in which a computer monitors the technical equipment to signal any malfunction and determines the components responsible. The detection and diagnosis of the fault may be followed by automatic actions, enabling the fault to be corrected such that the system may operate successfully even under the particular faulty condition.

 

Diagnostic concepts

 

Fault detection and diagnosis applies to both the basic technical equipment and the actuators and sensors attached to it. In the case of a chemical plant, the former includes the reactors, distillation columns, heat exchangers, compressors, storage tanks, and piping. Typical faults are leaks, plugs, surface fouling, and broken moving parts. The actuators are mostly valves, together with their driving devices (electric motors and hydraulic or pneumatic drives). The sensors are devices measuring the different physical variables in the plant, such as thermocouples, pressure diaphragms, and flow meters. Actuator and sensor fault detection is very important because these devices are prone to faults.

The on-line or real-time detection and diagnosis of faults means that the equipment is constantly monitored during its regular operation by a permanently connected computer, and any discrepancy is signaled almost immediately. On-line monitoring is very important for the early detection of any component malfunction before it can lead to more substantial equipment failure. In contrast, off-line diagnosis involves monitoring the system by a special, temporarily attached device, under special conditions (for example, car diagnostics at a service station).

The diagnostic activity may be broken down into several logical stages. Fault detection is the indication of something going wrong in the system. Fault isolation is the determination of the fault location (the component which malfunctions), while fault identification is the estimation of its size. On-line systems usually contain the detection and isolation stage; in off-line systems, detection may be superfluous. Fault identification is usually less important than the two other stages.

Fault detection and isolation can never be performed with absolute certainty because of circumstances such as noise, disturbances, and model errors. There is always a trade-off between false alarms and missed detections, with the proper balance depending on the particular application. In professionally supervised large plants, false alarms are better tolerated and missed detections may be more critical, while in consumer equipment (including cars) the situation may be the opposite.

 

Approaches

 

A number of different approaches to fault detection and diagnosis may be used individually or in combination.

 

Limit checking

 

In this approach, which is the most widely used, system variables are monitored and compared to preset limits. This technique is simple and appealing, but it has several drawbacks. The monitored variables are system outputs that depend on the inputs. To make allowance for the variations of the inputs, the limits often need to be chosen conservatively. Furthermore, a single component fault may cause many variables to exceed their limits, so it may be extremely difficult to determine the source. Monitoring the trends of system variables may be more informative, but it also suffers from the same drawbacks as limit checking.

 

Special and multiple sensors

 

Special sensors may be applied to perform the limit-checking function (such as temperature or pressure limit sensors) or to monitor some fault-sensitive variable (such as vibration or sound). Such sensors are used mostly in noncomputerized systems. Multiple sensors may be applied to measure the same system variable, providing physical redundancy. If two sensors disagree, at least one of them is faulty. A third sensor is needed to isolate the faulty component (and select the accepted measurement value) by “majority vote.” Multiple sensors may be expensive, and they provide no information about actuator and plant faults.

 

Frequency analysis

 

This procedure, in which the Fourier transforms of system variables are determined, may supply useful information about fault conditions. The healthy plant usually has a characteristic spectrum, which will change when faults are present. Particular faults may have their own typical signature (peaks at specific frequencies) in the spectrum.  See also: Fourier series and transforms

 

Fault-tree analysis

 

Fault trees are the graphic representations of the cause-effect relations in the system. On the top of the tree, there is an undesirable or catastrophic system event (top event), with the possible causes underneath (intermediate events), down to component failures or other elementary events (basic events) that are the possible root causes of the top event. Thelogic relationships from bottom up are represented by AND and OR (or more complex) logic gates. Fault trees can be used in system design to evaluate the potential risks associated with various component failures under different design variants (bottom-up analysis). In a fault diagnosis framework, the tree is used top down; once the top event is observed, the potential causes are analyzed by following the logic paths backward.

 

Parameter estimation

 

This procedure uses a mathematical model of the monitored system. The parameters of the model are estimated from input and output measurements in a fault-free reference situation. Repeated new estimates are then obtained on-line in the normal course of system operation. Deviations from the reference parameters signify changes in the plant and a potential fault. The faulty component location may be isolated by computing the new physical plant parameters and comparing them with those from the model.  See also: Estimation theory; Model theory

 

Consistency checking

 

This is another way of using the mathematical-system model. The idea is to check if the observed plant outputs are consistent with the outputs predicted by the model (Fig. 1). Discrepancies indicate a deviation between the model and the plant (parametric faults) or the presence of unobserved variables (additive faults). This testing concept is also called analytical redundancy since the model equations are used in a similar way as multiple sensors.

 

 

Fig. 1  Two stages of model-based fault detection and isolation. (After R. Isermann and B. Freyermuth, eds., Proceedings of the IFAC SAFEPROCESS Symposium, Baden-Baden, Germany, September 10–13, 1991, Pergamon, 1992)

 

 

 

 

 

 

In preparation for fault monitoring by analytical redundancy methods, a mathematical model of the plant needs to be established. This may be done from “first principles,” relying on the theoretical understanding of the plant's operation, or by systems identification using experimental data from a fault-free plant.

The actual implementation of fault monitoring usually consists of two stages (Fig. 1). The first is residual generation, where residuals are mathematical quantities expressing the discrepancy between the actual plant behavior and the one expected based on the model. Residuals are nominally zero and become nonzero by the occurrence of faults. The second stage is residual evaluation and decision making, where the residuals are subjected to threshold tests and logic analysis. Disturbances and model errors may also cause the residuals to become nonzero, leading to false alarms.

Fault isolation requires specially manipulated sets of residuals. In the most frequently used approach, residuals are arranged so that each one is sensitive to a specific subset of faults (structured residuals). Then in response to a particular fault, only a fault-specific subset of residuals triggers its test, leading to binary fault codes.

 

Principal component analysis (PCA)

 

In this approach, empirical data (input and output measurements) are collected from the plant. The eigenstructure analysis of the data covariance matrix yields a statistical model of the system in which the eigenvectors point at the “principal directions” of the relationships in the data, while the eigenvalues indicate the data variance in the principal directions. This method is successfully used in the monitoring of large systems. By revealing linear relations among the variables, the dimensionality of the model is significantly reduced. Faults may be detected by relating plant observations to the normal spread of the data, and outliers indicate abnormal system situations. Residuals may also be generated from the principal component model, allowing the use of analytical redundancy methods in this framework.  See also: Eigenfunction

 

 

Example of fault-tree analysis

 

The schematic of a simple electrical circuit in which a light is operated by a pair of three-way switches is shown in Fig. 2. (Such circuits are used in long hallways.) Figure 3 shows the detailed fault tree of the circuit. The tree goes down to subcomponents (contacts of the switches) in order to illustrate more complex logic relations on this simple system. Note that nonfailure events (operating conditions) are also among the basic events because such conditions (the position of each switch) determine whether a particular failure event triggers the top event.

 

 

Fig. 2  Simple electrical circuit: a lamp operated by a pair of three-way switches.

 

 

 

 

 

 

 

Fig. 3  Detailed fault tree of the circuit shown in Fig. 2.

 

 

 

 

 

 

 

Example of consistency checking

 

Traditionally, a few fundamental variables, such as coolant temperature, oil pressure, and battery voltage, have been monitored in automobile engines by using limit sensors. With the introduction of onboard microcomputers, the scope and number of variables that can be considered have been extended. Active functional testing may be applied to at least one actuator, typically the exhaust-gas recirculation valve. Model-based schemes to cover the components affecting the vehicle's emission control system are gradually introduced by manufacturers. One approach (Fig. 4) uses analytical redundancy to monitor two groups of actuators (fuel injectors and exhaust gas recirculation) and four sensors (throttle position, manifold pressure, engine speed, and exhaust oxygen). By the appropriate selection of the model relations, the residuals are insensitive to the load torque and the vehicle's mass. The structured residual technique is used to support fault isolation. The critical issue is to find sufficiently general models so that a single scheme may function well across an entire automobile product line and under widely varying operating conditions.  See also: Automotive engine; Microcomputer; Microprocessor

Janos J. Gertler

 

 

Fig. 4  Car engine system with onboard fault detection and diagnosis. (After R. Isermann and B. Freyermuth, eds., Proceedings of the IFAC SAFEPROCESS Symposium, Baden-Baden, Germany, September 10–13, 1991, Pergamon, 1992)

 

 

 

 

 

 

 

Bibliography

 

 

• L. H. Chiang, R. D. Braatz, and E. Russel, Fault Detection and Diagnosis in Industrial Systems, Springer, 2001
• P. L. Clemens, Fault Tree Analysis, 4th ed., Jacobs Sverdrup, 2002
• J. Gertler, Fault Detection and Diagnosis in Engineering Systems, 1998
• J. Gertler, Survey of model-based failure detection and isolation in complex plants, IEEE Control Sys. Mag., 8(7):3–11, 1988
• R. Patton, P. Frank, and R. Clark (eds.), Fault Diagnosis in Dynamic Systems, 1989
• Alifazeli=egeology.blogfa.com

 

Additional Readings

 

 

• Fault Tree Analysis
• Alifazeli=egeology.blogfa.com

تریلوبیت-Trilobita

Trilobita

class of extinct Paleozoic arthropods, occurring in marine rocks of Early Cambrian through late Permian age. Their closest living relatives are the chelicerates, including spiders, mites, and horseshoe crabs (Xiphosura). About 3000 described genera make trilobites one of the most diverse and best-known fossil groups (Fig. 1). Species diversity peaked during the Late Cambrian and then declined more or less steadily until the Late Devonian mass extinction. Only four families survived to the Mississippian, and only one lasted until the group's Permian demise. Their dominance in most Cambrian marine settings is essential to biostratigraphic correlation of that system.  See also: Cambrian; Chelicerata; Devonian; Permian

Trilobites are typically represented in the fossil record by the mineralized portion of their exoskeleton, either as carcass or molt remains. The mineralized exoskeleton (Fig. 2) was confined mostly to the dorsal surface, curved under as a rimlike doublure (Figs. 1h and 2b); a single mineralized ventral plate, the hypostome, was suspended beneath the median region of the head (Fig. 2b). The mineralized exoskeleton was composed of low magnesian calcite and a minor component of organic material. Most of the ventral exoskeleton, including the appendages, was unmineralized.

 

 

Fig. 1  Trilobite diversity and preservation. (a) Olenellid, Lower Cambrian (British Columbia). External mold of exoskeleton in shale. (b) Ogygopsis, Middle Cambrian (British Columbia). Molt assemblage, with cranidium and pygidium aligned, but librigenae, hypostome, and rostral plate inverted and rotated backward. Thorax is missing. Internal mold. (c) Slab of Upper Cambrian limestone (Sweden) with abundant disarticulated trilobite sclerites, mostly cranidia of Olenus. (d) Triarthrus, Upper Ordovician (New York). Exoskeletons replaced by pyrite, preserving antennae. (e–h) Struszia, Silurian (Northwest Territories), dorsal and anterolateral views of a cephalon, and ventral views of a cranidium and partial cranidium with attached librigena. Exoskeletons replaced by quartz; silicified fossils freed by dissolving limestone in acid. (i) Phacops, Devonian (Ohio). Enrolled exoskeleton, showing large lenses of schizochroal eye. (j) Griffithides, Mississippian (Indiana). Dorsal view of exoskeleton.

 

بزرگنمایی تصویر

 

 

 

 

Fig. 2  Morphological features of Trilobita. (a) Calymene, Silurian (New York). Dorsal view of mineralized exoskeleton. (b) Phacops, Devonian (Ohio). Ventral (bottom) view of cephalon and anterior segments of thorax, with hypostome attached. (c) Modocia, Middle Cambrian (Utah). Dorsal view of cephalon and thorax preserved in shale; pygidium and last thoracic segment missing.

 

 

 

 

 

 

Morphology of exoskeleton

 

The term "trilobite" refers to the longitudinal division of the body into an axial lobe and two lateral pleural regions (Fig. 2c); axial furrows separate the three divisions. The head shield, consisting of up to six fused segments and an anterior presegmental region, is called the cephalon. Its median (axial) lobe contains the glabella, typically convex and indented by a transverse occipital furrow and several pairs of lateral glabellar furrows (Fig. 2a). In primitive trilobites, segments of the palpebro-ocular (eye) lobes can be traced across the anterior region of the glabella (Fig. 1a).

Trilobites preserve the oldest known visual system in the history of life. Most had rigid compound eyes analogous to those of a housefly. The eyes are situated on the pleural field (genae, or cheeks). Most trilobites had a large number of small eye lenses that shared a single corneal covering (holochroal eye) [Fig. 1e and f.] The suborder Phacopina, a major Ordovician-through-Devonian group, had large separated lenses (schizochroal eye) [Fig. 1i].

In most trilobites, a facial suture, used for molting the exoskeleton, is developed on the dorsal side of the cephalon; it passes from the ventral side usually in front of the glabella, separates the visual surface of the eye from the palpebral lobe, and exits the cephalon in front of, through, or behind the genal angle. These different configurations of the suture are termed proparian (Fig. 1e–h), gonatoparian (Fig. 2a), and opisthoparian (Figs. 1j and 2c), respectively. The area between the axial furrow and facial suture is the fixigena (fixed cheek); together with the axial region of the cephalon (including the glabella), this single skeletal part (or sclerite) is the cranidium (Fig. 1g). The librigenae (free cheeks) represent the pleural areas outside the facial suture. Most trilobites had the anterior branches of the facial suture separated on the doublure by a rostral plate (Fig. 1b and f), although some had a median suture and others lost the ventral sutures and fused the doublure medially. The hypostome was rigidly sutured to the roof of the doublure in some groups (the conterminant condition; Fig. 2b), but in others it was free and supported by soft tissue (the natant condition).

Anterior wings on the lateral part of the hypostome bear processes that permitted its attachment to a stalk in the cephalic axial furrow (Fig. 2b). The thorax is composed of from 2 to more than 60 articulated segments (although typically 6 to 16), each consisting of an axial ring and pleural band. Articulation of the thorax, via processes and sockets on adjacent pleurae, allowed flexibility for enrollment (Fig. 1i). The pygidium is a posterior sclerite composed of one or more fused segments. Primitively, it is much smaller than the cephalon, but is enlarged in many groups (Fig. 1b), and may bear spines along its margin. The cephalic doublure sometimes has notches or furrows that accommodated the pygidium and thoracic tips when the trilobite enrolled (Fig. 2b).

Appendages, preserved by pyrite or phosphate replacement or as films on shale, are well known for only a few trilobite species. A single pair of long, jointed antennae (Fig. 1d) projected forward from beneath the hypostome. Known Cambrian and Ordovician species have three pairs of postantennal cephalic appendages, while a Devonian example has four. In most cases, these show little structural differentiation from each other, or from postcephalic appendages on each segment along the length of the body (Fig. 3). The appendages are biramous, consisting of a jointed walking leg, or telopodite, and a filamentous exite, which attach toward the body axis to a spine-bearing coxa. Appendage-related musculature attached to the ventral exoskeleton at knoblike apodemes (Fig. 1e and h), just inward of the axial furrow. Enrollment and outstretching were achieved by flexor and extensor muscles; longitudinal, dorsoventral, and horizontal muscles have been observed, as well as a system of intersegmental bars. The exite (gill branch) functioned as a respiratory organ. The mouth opening was positioned above the rear margin of the hypostome and was directed posteriorly. The gut looped backward beneath the glabella, with the digestive tract extending along the axis to a posterior anus.

 

 

Fig. 3  Triarthrus eatoni, Upper Ordovician (New York). Reconstruction with dorsal exoskeleton removed on right side to show appendages. Antennae are incomplete (compare Fig. 1d). Exites of first nine postantennal appendages are removed to show structure of telopodite. The mouth was positioned above the posterior margin of the hypostome. (After H. B. Whittington and J. E. Almond, Appendages and habits of the Upper Ordovician trilobite Triarthrus eatoni, Phil. Trans. Roy. Soc. Lond., B317:28, 1987)

 

 

 

 

 

 

 

Development and molting

 

Embryonic development of trilobites is unknown. Phosphatized arthropod eggs, which may be those of trilobites, have been discovered in Cambrian rocks. The term "protaspis" is applied to the earliest calcified larval stages, in which the cephalon and protopygidium are fused as an unjointed dorsal shield (Fig. 4). Several molts may occur within the protaspid period. The meraspid period is defined by articulation of the cephalon and transitory pygidium as separate sclerites; successive degrees are marked by the release of segments from the anterior part of the transitory pygidium to form the thorax. The holaspid has the complete adult complement of thoracic segments; development in this period is marked by continued increase in size and by changes in shape, but without further addition of segments to the thorax. Adult size ranges from 1.5 mm to 70 cm (0.06 to 28 in.); 2–5 cm (0.8–2 in.) is typical.

 

 

Fig. 4  Flexicalymene senaria, Middle Ordovician (Virginia). Complete exoskeletons of protaspid larvae obtained from silicified residues. (a) Dorsal view and (b) ventral view of second of four protaspid instars for this species. (c) Dorsal view and (d) ventral view of fourth protaspid instar. Holaspides closely resemble the related genus Calymene (Fig. 2a). (After B. D. E. Chatterton et al., Larvae and relationships of the Calymenina (Trilobita), J. Paleontol., 64:259, 1990)

 

 

 

 

 

 

Trilobites show the typical arthropod solution to the problem of increasing size with a stiffened exoskeleton: they molted at regular intervals throughout the life cycle. In most species, this was effected by shedding the librigenae along the facial suture and shedding the hypostome. The soft-bodied animal emerged from the resulting gap. Several different molt strategies were employed by different trilobite groups, however, including shedding the entire cephalon, and inverting and rotating various skeletal elements (Fig. 1b). Molting results in the typical preservation of trilobite remains as disarticulated sclerites (Fig. 1c).

 

Ecology and macroevolution

 

Most trilobites were benthic deposit feeders or scavengers, living on the sediment-water interface or shallow-burrowing just beneath it. Some were evidently carnivores, equipped with sharp spines and processes projecting ventrally from their appendages. A few Cambrian and Ordovician groups acquired giant eyes coupled with narrow, streamlined bodies. The morphology and broad geographic and environmental ranges of these groups suggest they were active swimmers. Through their history, trilobites became adapted to all marine environments, from shallow high-energy shorefaces to deep-water, disaerobic habitats.

Trilobites are the most common marine fossils of the Cambrian Period, and their remains typically account for more than 90% of preserved Cambrian fossil assemblages. They were important through the Early Ordovician, but their numerical contribution to onshore communities was much reduced as a result of the Ordovician Radiation of marine life. This event saw filter-feeding organisms (for example, articulate brachiopods, bryozoans, crinoids) proliferate and rapidly evolve to dominate marine communities, a pattern that would last through the Paleozoic Era. Trilobites remained major components of deeper-water communities through the Silurian. Within-habitat, species diversity was generally constant in all environments from the Cambrian through the Silurian, despite their increasingly reduced relative abundance. This indicates that trilobites were largely unaffected by the major events of the Early Paleozoic and that their decline in importance was largely a function of increases in other groups.

Global trilobite diversity increased rapidly following the acquision of hard parts during the Cambrian Explosion, and it peaked during the Late Cambrian. Overall diversity gradually declined during the Ordovician, although a major subset of trilobite groups experienced an evolutionary burst during the Ordovician Radiation. The end-Ordovician mass extinction decimated the group, cutting their global diversity by about half. Surviving families were mainly those that had radiated during the Middle Ordovician. Global diversity continued to decline during the Silurian, although the most speciose trilobite faunas ever found occurred in this period. By the Devonian, trilobites were a relatively minor group, absent from many marine faunas, although still sometimes locally abundant. The Late Devonian mass extinction all but obliterated the trilobites, as only a handful of lineages survived. During the Late Paleozoic, trilobites were typically rare and confined to a limited number of facies. The last trilobites became extinct during the great end-Permian mass extinction.  See also: Cambrian; Ordovician; Paleozoic; Silurian

 

Classification

 

Trilobita is usually assigned the ranking of class within Arthropoda. Affinities with Chelicerata are expressed by their grouping as Arachnata. Older classifications recognized a phylum or subphylum Trilobitomorpha, grouping Trilobita with an unnatural assortment of trilobite-, chelicerate-, or crustacean-like taxa lumped as Trilobitoidea. The soft-bodied Early-Middle Cambrian order Nectaspida is the closest relative (sister group) of the calcified Trilobita.

The high-level classification of trilobites remains controversial. Post-Cambrian groups (for example, orders Phacopida, Odontopleurida, Lichida, Proetida, Aulacopleurida) are well understood and are grouped into orders or suborders based on distinctive adult and larval morphologies. Cambrian trilobites are generally less well known (despite their abundance as fossils) and have tended to be classified in a small number of large unnatural orders such as Ptychopariida. A particular problem is a lack of understanding of the origins of post-Cambrian orders among Cambrian taxa, a phenomenon termed cryptogenesis. The result is that relationships between named orders of trilobites are essentially unknown. Recent progress has resulted from study of silicified life histories (Fig. 4), but inferring the high-level phylogeny of trilobites remains the cardinal problem in the paleobiology of the group.

A group of blind marine arthropods, the Agnostida, has traditionally been recognized as an order of trilobites. Agnostids share a calcified dorsal exoskeleton with Trilobita, but otherwise lack most diagnostic trilobite features, including a calcified protaspid stage, facial sutures, articulating thoracic segments, and a true transitory pygidium. The appendages of agnostids are also fundamentally unlike those of trilobites. Their affinities are currently debated, with some workers defending their position as ingroup trilobites and others considering the agnostids to be stem group Crustacea.  See also: Arthropoda; Taxonomy

Gregory D. Edgecombe

Jonathan Adrain

 

Bibliography

 

 

• R. A. Fortey, Ontogeny, hypostome attachment and trilobite classification, Palaeontology, 33:529–576, 1990
• R. A. Fortey, Trilobite! Eyewitness to Evolution, 2000
• R. L. Kaesler (ed.), Treatise on Invertebrate Paleontology, pt. O (rev.), vol. 1, 1997
• H. B. Whittington, Trilobites, 1990
• Alifazeli=egeology.blogfa.com

 

Additional Readings

 

 

What are Trilobites?

Alifazeli=egeology.blogfa.com

 

 

گسلهای ترانسفورم-Transform fault

Transform fault

One of the three fundamental types of boundaries between the mobile lithospheric plates that cover the surface of the Earth. Whereas spreading centers mark sites where crust is created between diverging plates, and subduction zones are where crust is destroyed between convergent plates, transform faults separate plates that are sliding past each other with neither creation nor destruction of crust. The primary tectonic feature of all transform faults is a strike-slip fault zone, a generally vertical fracture parallel to the relative motion between the two plates that it separates. Strike-slip fault zones are described as right-lateral if the far side is moving right relative to the near side (for example, the Queen Charlotte zone; Fig. 1), left-lateral if it is moving to the left (for example, the North Caribbean zone; Fig. 1). Not all such fault zones are plate-bounding transform faults. Small-scale strike-slip faulting is a common secondary feature of many subduction zones, especially where plate convergence is oblique, and of some spreading centers, especially those with propagating rifts; it also occurs locally deep in plate interiors. The distinguishing characteristic of a transform fault is that both ends extend to a junction with another type of plate boundary. At these junctions the divergent or convergent motion along the other boundaries is transformed into purely lateral slip. See also: Earth crust; Plate tectonics; Subduction zones

 

 

Fig. 1  Various types of transform fault mapped in parts of North and Central America. Area inside circle is shown in detail in Fig. 3.

 

بزرگنمایی تصویر

 

Types

 

Transform faults are most readily classified by the types of plate boundary intersected at their ends, the variety of lithosphere (oceanic or continental) they separate, and by whether they are isolated or are part of a multifault system. The common oceanic type is the ridge-ridge transform, linking two literally offset axes of a spreading center (for example, Clipperton and Siqueiros transforms; Fig. 1). Also common are transform faults that link the end of a spreading center to a triple junction, the meeting place of three plates and three plate boundaries. For example, the Panama transform links a spreading axis on the mid-oceanic ridge between the diverging Cocos and Nazca plates to a Cocos-Nazca-Caribbean plate triple junction at the continental margin of Central America. See also: Lithosphere; Mid-Oceanic Ridge

 

 

Fig. 2  The effect of changing plate motion on transform faults and their fracture zones. (a) East-west plate motion; offset spreading axes with a left-lateral North Transform and a right-lateral South Transform. (b) Change in motion; a small rotation of the direction of plate motion adds components of plate convergence to North Transform and divergence to South Transform. (c) Adjustment to change; fault is segmented into two fault zones aligned parallel to the new plate motion, which are linked by a short new spreading axis.

 

 

 

 

 

 

Other types are long trench-trench transforms at the northern and southern margins of the Caribbean plate, and the combined San Andreas/Gulf of California transform, which separates the North American and Pacific plates for 1500 mi (2400 km) between triple junctions at Cape Mendocino (California) and the mouth of the Gulf of California (Fig. 1). Strike-slip faulting in the Gulf of California (and on the northern Caribbean plate boundary) occurs along several parallel (en echelon) zones linked by short spreading centers, and the overall structure is more properly called a transform fault system; similar fault patterns are found at many ridge-ridge transforms. The San Andreas part of this plate boundary exhibits another type of transform fault system, one with several simultaneously active zones that overlap, rather than replacing each other in stepwise fashion; this pattern may be characteristic of wholly continental transforms.

 

Geology of oceanic transforms

 

The structure of ridge-ridge transforms on mid-oceanic ridges varies to some extent with their length, ranging 6–600 mi (10–1000 km), and with the rate of slip of their strike-slip faults 0.8–8 in. (20–200 mm) per year, but the structure depends mainly on the geologic history of changing plate motions. In the absence of such changes, where transform faults separate plates that have maintained the same motion for millions of years, the characteristic structure is a transform valley parallel to the direction of relative plate motion, and thereby transverse to the mid-oceanic ridge. The valley floor is occupied by a strike-slip fault zone, a band of shattered rock only 300–600 ft (100–200 m) wide, that is often marked by a groove or rift in the sea floor. There is some correlation of valley depth and transform length, and most ridge-ridge transforms longer than 60 mi (100 km) have valleys deeper than the 3300–6600-ft-deep (1000–2000-m) axial rift valleys typical of the crests of slow-spreading ridges. Slow-slipping transform valleys are generally deepest at their ends, at their orthogonal intersections with axial rift valleys, whereas the deepest parts of fast-slipping valleys are usually in their midsections, their ends being partly filled with lava that spills over from the intersecting spreading axes. See also: Lava

Transform valleys are structural troughs opened by a small component of extension across ridge-ridge transform faults. Where the faults are strictly parallel to relative plate motion, the origin of this extensional stress is probably thermal contraction of the young lithosphere accreted at the intersecting spreading axes. Larger components of extension, creating deeper transform valleys, can result from small angular changes in the direction of plate motion; conversely, opposite changes can add a component of valley-closing compression. The motions of most oceanic plates are changing continuously, albeit slowly, affecting all the transform faults along their boundaries. A small rotation in the direction of plate motion has an opposite effect on the stresses at adjacent left-lateral and right-lateral transforms (Fig. 2). Such a motion has affected adjacent transform faults on part of the East Pacific Rise, where relative plate motion (spreading direction) has rotated anticlockwise by about 5° in the past 5 million years. The resulting convergence across the left-lateral Clipperton transform has closed its transform valley and thrust up a median ridge 2000 ft (600 m) high, with the strike-slip fault zone along its crest. The same rotation has caused extension at the right-lateral Siqueiros transform (Fig. 1). Some transforms react to a change of this sort by opening wider and deeper transform valleys, often accompanied by uplift of the valley margins by 0.6–1.8 mi (1–3 km) to form structures known as transverse ridges. In some cases, opening transform faults by adding a component of extension allows seawater to penetrate deep into the lithosphere, chemically altering the upper mantle to a low-density rock (serpentinite) that rises along the fault zone, forming median ridges similar in shape but quite different in origin to those at compressed transform faults. In a few examples, crustal divergence across the transform fault allows magma to leak out, building yet another type of median ridge. The Siqueiros transform shows a more common response to the change in plate motion. Instead of maintaining the same strike-slip fault zone, and adding an extensional component to its lateral motion, it developed a set of new fault zones, each parallel to the new plate motion, and slightly oblique to the overall trend of the transform fault system. Magma does leak out within the fault system, but only at the short new spreading axes which link the parallel fault zones. See also: Serpentinite

The geology of ridge-ridge transform faults is sensitive to the history of recent changes in the direction of plate motion, and the pattern of the fracture zones that they leave on the flanks of the mid-oceanic ridges provides a record of these changes. Fracture zones are bands of rough topography that extend down ridge flanks from the ends of transform faults. Their name is inherited from an early false interpretation that they are belts of strike-slip faulting across the flanks of mid-oceanic ridges. A foundation of the theories of sea-floor spreading and plate tectonics was the recognition that fracture zones are not active fault zones, merely the seams between crust that differs in age becuase it has spread different distances from laterally offset spreading axes. The lateral offsets occur at transform faults, so fracture zones are their inactive continuations. The azimuth of a fracture zone is parallel to the direction of plate motion at the time that the crust on its younger side spread off the risecrest, with fracture zone bends and kinks marking changes in direction. Mapping the trends of fracture zones is a principal method of investigating the past movements of lithospheric plates. See also: Marine geology

 

Geology of continental transforms

 

Transform faults within the continental lithosphere fracture crust that is much thicker and less homogeneous than oceanic crust. Perhaps as a result, the fault zones tend to be less straight, with many local deviations from azimuths parallel to relative plate motion. Bends in the fault zones add components of extension or compression to the dominantly strike-slip motion, resulting in along-strike alternations of collapsed extensional basins and uplifted compressional ridges. A well-known compressional bend is the Big Bend of the San Andreas fault zone north of Los Angeles (Fig. 3), where oblique convergence of the Pacific and North American plates is raising the San Bernandino and San Gabriel mountains. Some of the sediment-filled basins formed along continental transform faults are important petroleum reservoirs, and secondary deformation on the margins of the fault zone commonly folds the sediment layers to form trapping structures for oil fields. See also: Basin; Petroleum geology

 

 

Fig. 3  Map showing parts of the San Andreas and Gulf of California fault systems. Many minor but still earthquake-generating strike-slip fault zones have been omitted. The abandoned offshore fault zones, relics of a time when the plate boundary was closer to the continental margin, still have a low level of residual earthquake activity. The individual fault zones are San Andreas (SA), Hosgri (H), South San Andreas (SSA), San Jacinto (SJ), Elsinore (E), San Clemente (SC), San Benito (SB), and Guaymas (G).

 

 

 

 

 

 

Most continental transform fault systems have several belts of faulting, with complex spatial patterns of overlapping and splaying fault zones, and complex geologic histories, involving constant shifting of the share of the total interplate displacement among several zones, accompanied by the birth of new fault zones and the abandonment of others. In southern California, for example, motion on the San Andreas transform fault system is now concentrated on three narrow fault zones (Fig. 3) which differ in age and are accompanied by a multitude of less active subparallel zones, some of which may become dominant traces in the near geologic future. Very detailed geologic studies are needed to unravel the histories of continental transform fault systems, which do not leave fracture zone traces like their oceanic counterparts. The total lateral displacement between the two sides of a fault zone, commonly amounting to tens or hundreds of miles, can be estimated by recognizing the two displaced halves of preexisting geologic features that were split and separated by fault motion. On a much shorter time scale, recent displacements can be monitored by offsets in human-made features such as fence lines.

Some of the less active fault zones in the region of the San Andreas and Gulf of California systems are senescent rather than nascent transform faults. Until 6 million years ago, transform faulting was centered west of Baja California and west of the southern California coastline, and the inland shift of the Pacific–North America plate boundary has caused the almost complete cessation of faulting on the offshore San Benito and San Clemente fault zones.

 

Shearing continental margins

 

The now-inactive San Benito fault zone was representative of an important class of transform faults that extend along continental margins, at or near the boundary between oceanic and continental lithosphere. A still-active example is the Queen Charlotte fault zone off the British Columbia coast (Fig. 1). Continental margins shaped by lateral shearing (transform faulting) have very steep continental slopes, but often with steps on the slope called marginal plateaus, crustal blocks that have subsided between the shifting fault zones of a transform fault system. The shifting is commonly away from the oceanic/continental boundary into adjacent weaker continental lithosphere; indeed, the San Andreas fault system can be considered a marginal shear zone that has shifted unusually far inland.

Most shearing margins of western North America were formerly, with an earlier arrangement of a lithospheric plates, convergent (subduction zone) margins. Shearing margins with active transform faults also play a role during the birth of ocean basins by continental rifting. Initial rifting, as in the split of North America from Africa about 200 million years ago, is commonly on laterally displaced fractures that develop into spreading axes linked by transform faults. These transform faults become part of the oceanic/continental boundary once continental separation has proceeded far enough for sea-floor spreading to occur, and eventually become ridge-ridge transforms once a risecrest develops in the new ocean basin. Many of the ridge-ridge transform faults on the Mid-Atlantic Ridge are inherited from fault zones that once formed shearing parts of the continental margin. Shearing margins occur on the boundaries of the very small, young ocean basins that have opened by the splitting of Baja California from mainland Mexico (for example, Guaymas Basin, with the Guaymas transform fault on its northeast side; Fig. 3). See also: Continental margin

 

Earthquakes

 

Along a few strike-slip fault zones, lithospheric plates slide quietly and almost continuously past each other by the process called aseismic creep. Much more often, frictional resistance to the sliding in the brittle crust causes the accumulation of shear stresses that are episodically or periodically relieved by sudden shifts of crustal blocks, creating earthquakes. The largest lateral shifts (slips) of the ground surface along major continental transform faults have been associated with some of the largest earthquakes on record; in 1906 the Pacific plate alongside 270 mi (450 km) of the San Andreas Fault suddenly moved an average of 15 ft (4.5 m) northwest relative to the North American plate on the other side, and the resulting magnitude-8.2 earthquake destroyed much of San Francisco. The average slip in this single event was equivalent to about 150 - 250 years of Pacific–North American plate motion.

The maximum size of earthquake that a transform fault can generate is limited by the length of the fault, though generally, even in large earthquakes like the one in San Francisco in 1906, a fault does not fail along its entire length. The frequency of earthquakes is controlled by the average speed of relative plate motion across a transform fault plate boundary and, for fault systems with multiple overlapping fault zones, by the share of this motion that is carried by any individual fault. However, many local geologic and tectonic factors intrude to complicate estimates of how frequently a particular transform fault will produce earthquakes of any specified size or destructive power, and it is still more difficult to predict the exact timing of such an event. Extrapolation of the past record is probably the best method of estimating future magnitudes and frequencies, becuase many transform faults do seem to have a characteristic size of large earthquakes, and a consistent ratio of small to large events. In most cases, however, especially for remote oceanic transform faults, the record is too short and too incomplete to be of much practical use. See also: Earthquake; Fault and fault structures; Seismology

Peter Lonsdale

 

Bibliography

 

 

• W. G. Ernst (ed.), The Geotectonic Development of California, 1981
• P. J. Fox and D. G. Gallo, The geology of North Atlantic transform plate boundaries and their aseismic extensions, The Geology of North America, vol. M, pp. 157–172, 1986
• P. J. Fox and D. G. Gallo, Transforms of the eastern central Pacific, The Geology of North America, vol. N, pp. 111–124, 1989
• J. T. Wilson, A new class of faults and their bearing on continental drift, Nature, 207:343–347, 1965
• Alifazeli=egeology.blogfa.com 

 

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EarthView 3.7.1

یك كاغذ دیواری زنده و یك محافظ صفحه‌ی نمایش كه تصاویر زیبایی از سیاره‌ی مظلوم زمین را در طول شبانه روز به نمایش می‌گذارد. این نرم‌افزار تصاویری رنگی و پر‌وضوح را در كلیه‌ی مقاطع تولید كرده و نقشه‌ها و مناظر شهری، جلوه‌های جوی، نمایش وقت محلی و خیلی چیزهای دیگر را در چنته دارد. جزییاتی را كه از سیاره‌ی زمین در صفحه‌ی نمایش كامپیوتر خود خواهید دید، بهترین دلیل برای نصب این

نرم‌افزار است.

برای بازدید از سایت رسمی این نرم‌افزار این‌جا را كلیك كنید. برای دریافت بدنه‌ی سه‌ مگابایتی آن نیز می‌توانید تا ادامه‌ی متن چند قدم بردارید...

EarthView 3.7.1

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Adobe Reader 8.1.1 for Windows XP SP2/ Vista

اين هم جدیدترین روایت نرم‌افزاری است كه با ویژگی‌ها و امكانات واقعاً جدید و منحصر به فرد خود برای رویت اسناد و مدارك مبتنی بر ‌پی‌دی‌اف و با شكل و شمایل كاملا دگرگون شده اكنون در اختیارتان قرار دارد. این نرم‌افزار با رابط جدید، ابزارهای تازه و گزینه‌های بیش‌تر خود بهترین وسیله برای دیدن و خواندن اسناد و كتاب‌های پی‌دی‌اف است.

این نرم‌افزار كمكی در حدود 22 مگابایت حجم دارد و در ادامه‌ی متن منتظر آدم‌های كنجكاو نشسته است. برای بازدید از سایت رسمی این نرم‌افزار نیز می‌توانید از این محل بازدید كنید...

Adobe Reader 8.1.1 for Windows XP SP2/ Vista

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Lavasoft Ad-Aware 2007 PRO v7.0.2.3

این نرم‏افزار یك استودیوی ضبط صدای پیشرفته و كاملاً مجهز است كه با داشتن  تمام امكانات و جلوه‏های صوتی، از شما یك میكس‏كننده‏ی حرفه‏ای می‏سازد. این نرم‏افزار یك قوه‏ی محركه‏ی صوتی 32بیتی عالی دارد كه می‏تواند فایل‏های صوتی از نوع WAV،MP3 ،OGG ، و WMA را پشتیبانی كرده و به این ترتیب شما را قادر می‏سازد با تركیب چند قطعه‏ی موسیقی با یك‏دیگر، یك قطعه‏ی جدید خلق كرده و آن را به صورت یك سی‏دی درآورید و یا در اینترنت منتشر نمایید.  حقیقت حیرت انگیزی كه در مورد استودیوهای ضبط خانگی امروزی وجود دارد آن است كه شما فقط به یك كامپیوتر و یك نرم‏افزار ضبط چند كاناله مانند همین نرم‏افزار احتیاج دارید تا موسیقی‏های خارق‏العاده‏ی خود را خلق نمایید. طرز كار با این نرم‏افزار، ساده‏تر از آن است كه تصور می‏كنید. مزیت اصلی این نرم‏افزار، تعداد بی‏نهایت كانال‏های آن است كه به ظرفیت كلّی كامپیوترتان بستگی دارد. شما می‏توانید در هر كانال (یا شیار) یك قطعه‏ی موسیقی را قرار داده و پس از اعمال تغییرات و ویرایش‏های لازم، همه‏ی آن‏ها را با هم تركیب نموده و به صورت یك قطعه‏ی واحد درآورید. كارهایی كه می‏توانید با این نرم‏افزار انجام دهید، آن‏قدر متنوع و متعدد هستند كه اگر بخواهیم در مورد هر كدام از آن‏ها به بحث بپردازیم ، به صد من كاغذ نیاز خواهیم داشت. بنابراین ، صرفاً برای آن‏كه اوضاع دست‏تان بیاید، بهتر است به ادامه‌ی مطلب بروید و كالبد 10 مگابایتی آن را دریافت كرده و خودتان آن را از نزدیک امتحان کنید. برای بازدید از منزل این نرم‌افزار سرشناس نیز می‌توانید این‌جا را كلیك كنید...

Acoustica Mixcraft 3.1 Build 41

 

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Name: Acoustica

Serial: 07DDB-UZMPM-A4UBR-Z1AGW

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Name: Mixcraft

Serial: 09PUM-RIZEH-WAK4D-XQMVR

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Name: AcousticaMixcraft

Serial: 00MA4-UBRZ1-AGWZE-1B4I1

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Name: multi-trackaudio

Serial: 00AW4-XAL3S-HXWXG-C3YJK

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Name: multi-trackrecorder

Serial: 00AW4-XAL3S-HXWXG-R4OOM

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Acoustica Mixcraft 3.1 Build 41 Full

 

DU Meter v4.00.Build.R3009

یك نرم‌افزار برجسته كه دقیقا نشان می‌دهد در هر لحظه چه مقدار اطلاعات از اتصالات شبكه‌ی كامپیوتری شما در حال عبور و مرور است. این نمایش هم به صورت یك نمودار عددی و هم به صورت یك نمودار گرافیكی، شما را در جریان وضعیت داده‌هایی كه از كامپیوترتان خارج می‌شوند، و داده‌هایی كه وارد كامپیوترتان می‌گردند، قرار می‌دهد.

 جدیدترین روایت این نرم‌افزار به حجم تقریبی یك و نیم مگابایت، یكی از ابزار‌های ضروری برای همه‌ی آن‌هایی است كه به نوعی وارد شبكه‌ها می‌شوند- چه شبكه‌های موضعی و چه شبكه‌های اینترنتی- و می‌خواهند ببینند كه میزان تبادلات داده‌ای آن‌ها چقدر است. برای دریافت این وسیله‌ی كارآمد، به ادامه‌ی متن تشریف ببرید، اما برای بازدید از سایت رسمی آن كافی است این‌جا را كلیك كنید...

DU Meter v4.00.Build.R3009

Crack

NextUp TextAloud v2.261

یك نرم‌افزار فوق‌العاده برای تبدیل متون نوشتاری به كلام ناطق. در واقع این برنامه از صداهای مصنوعی استفاده می‌كند تا نوشته‌ها را به اصوات گفتاری تبدیل نماید. این برنامه متون نوشتاری را از ای‌میل، صفحات وب، گزارشات و چیزهایی از این قبیل می‌خواند و آن‌ها را به صورت اصوات شنیداری در می‌آورد. به علاوه این برنامه می‌تواند مطالب خواندنی را به صورت فایل‌های صوتی نیز تبدیل كند تا بتوانید آن‌ها را روی دستگاه‌های پخش صوت مانند ام‌پی‌تری پلیرها نیز پخش نمایید. اگر دوست دارید یك نفر مفت و مجانی بنشیند و برای‌تان كتاب و مجله و مقاله بخواند، دست به دامن این نرم‌افزار شوید.

این نرم‌افزار در حدود 3 مگابایت حجم دارد كه می‌توانید كالبد آن‌را در ادامه‌ی متن دریافت كنید. برای بازدید از سایت رسمی آن نیز می‌توانید این‌جا را كلیك كنید...

NextUp TextAloud v2.261

 

Keygen

SlySoft CloneCD 5.3.1.0

كامل‌ترین نرم‌افزار برای كپی‌گرفتن از سی‌دی‌های صوتی و داده‌ای خود ؛ آن هم كپی‌هایی كه با اصل خود مو نمی‌زنند- حتا سی‌دی‌هایی كه قفل خورده‌ باشند. این نرم‌افزار منحصر به فرد كه همه‌ی ایرانی‌ها ارادت خاصی به آن دارند، به شما امكان می‌دهد كه ظرف چند دقیقه هر نوع سی‌دی را كپی كنید.

روایت پنجم این نرم‌افزار علاوه بر امكان كپی گرفتن از سی‌دی‌ها و قابلیت كپی‌كردن تمام فرمت‌های دی‌وی‌دی ، خود را با ویندوز ویستا نیز تطبیق داده است. سرتان را درد نمی‌آوریم. خودتان خوب می‌دانید این گوهر نرم‌افزاری چقدر گران‌بهاست. اما اگر هنوز هم شك دارید، بهتر است سری به سایت رسمی آن بزنید تا با قابلیت‌های این نرم‌افزار از نزدیك آشنا شوید. اگر هم تصمیم‌تان را گرفته‌اید، می‌توانید به ادامه‌ی متن بروید و جثه‌ی دو و نیم مگابایتی آن را دریافت كنید...

SlySoft CloneCD 5.3.1.0

Mirror

Smart Pix Manager v9.01

همه‌ی شما روی كامپیوتر‌تان انبوهی از عكس، تصاویر، فایل‌های صوتی، اسناد و فیلم‌های ویدیویی دارید كه آن‌را مثل بازار شام كرده‌اند! اما برای اداره كردن این همه مواد چند رسانه‌ای و یا گشتن به دنبال یك عكس یا فایل گرافیكی همواره یك راه بهتر نیز وجود دارد. این نرم‌افزار یكی از آن راه‌های بهتر است كه می‌تواند شما را در پیدا كردن هر آن‌چه كه به دنبالش هستید، كمك كرده و با ویرایش و تبدیل كردن تصاویر به هم، و محافظت از آن‌ها به وسیله‌ی رمز عبور از مجموعه‌ی شخصی تان نگهداری كرده و شما را در مدیریت آن‌ها به خوبی همراهی نماید.

برای پی‌بردن به امكانات و ویژگی‌های این نرم‌افزار لازم است از سایت رسمی آن بازدید كنید. اما اگر تصمیم‌تان را گرفته‌اید و می‌خواهید كالبد سه و نیم مگابایتی آن‌را دریافت كنید، كافی است به ادامه‌ی متن تشریف ببرید...

Smart Pix Manager v9.01

Mirror

VSO PhotoDVD v2.5.0.4b

یكی از بهترین نرم‌افزار‌ها برای جمع‌آوری عكس‌ها و تبدیل آن‌ها به یك دی‌وی‌دی. این برنامه‌ می‌تواند یك نمایش عكس به عكس از تصاویر‌تان درست كرده و آن‌ها را به بهترین شكل ممكن، روی دی‌وی‌دی حك نماید. به این ترتیب می‌توانید عكس‌های خود را روی پخش‌كننده‌ی دی‌وی‌دی منزل‌تان به تماشا نشسته و یا آن‌ها را در اختیار دوستان و آشنایان‌تان قرار دهید. یك ساحره‌ی قدم‌به‌قدم نیز در این برنامه وجود دارد كه كار با آن را به صورت ساده‌تری در می‌آورد.

برای بازدید از سایت رسمی این نرم‌افزار، از منزل آن دیدن كنید، برای دریافت بدنه‌‌ی 9 مگابایتی‌اش نیز می‌توانید به ادامه‌ی متن تشریف ببرید...

VSO PhotoDVD v2.5.0.4b

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