Desert Landscapes and Life

Desert Landscapes and Life

kacimo samy   7:19 ص

Desert Landscapes and Life 

The popular media commonly portray deserts as endless vistas of sand, punctuated by the occasional palm-studded oasis. In reality, not all desert landscapes are buried by sand. Some deserts  are vast, rocky plains; others sport a stubble of cacti and other hardy desert plants; and still others display intricate rock formations that look like medieval castles. Explorers of the Sahara, for example, traditionally distinguished among hamada (barren, rocky highlands), reg (vast, stony plains), and erg (sand seas in which large dunes form). 

In this post, we’ll see how the erosional and  depositional processes described above lead to the formation of such contrasting landscapes.

Rocky Cliffs and Mesas 

In hilly desert regions, the lack of soil exposes rocky ridges and cliffs. As noted earlier, cliffs erode when rocks split away along joints. When this happens, the cliff face steps back into the land but retains roughly the same form. The process, commonly referred to as cliff retreat, or scarp retreat, occurs in fits and starts. A cliff may remain unchanged for decades or centuries, and then suddenly a block of rock falls off and crumbles into rubble at the foot of the cliff. Cliffs exposing alternating layers of strata with contrasting strength develop a step-like shape; strong layers (sandstone or limestone) become vertical cliffs, and weak layers (shale) become rubble-covered slopes. 

Mesas and buttes form in deserts as cliffs retreat over time.

With continued erosion and cliff retreat, a plateau of rock slowly evolves into a cluster of isolated hills, ridges, or columns. Flat-lying strata or flat-lying layers of volcanic rocks erode to make flat-topped hills, which go by different names depending on their size. Large examples (with a top surface area of several square km) are mesas, from the Spanish word for table. Medium-sized examples are buttes (figure above a, b), and small examples, whose height greatly exceeds their top surface area, are chimneys (figure above c). Natural arches form when erosion along joints leaves narrow walls of rock. When the lower part of the wall erodes while the upper part remains, an arch results. 

 The formation of cuestas and inselbergs, due to erosion and deposition in deserts.

In places where bedding dips at an angle, an asymmetric ridge called a cuesta develops. A joint-controlled cliff forms the steep front side of a cuesta, and the tilted top surface of a resistant bed forms the gradual slope on the backside (figure above a). If the bedding dip is steep to near vertical, a narrow symmetrical ridge, called a hogback, forms. 

After a long period of erosion, all that may remain of a once broad region of uplifted land is a relatively small island of rock, surrounded by alluvium-filled basins. Geologists refer to such islands of rock by the German word inselberg (island mountain; figure above b). Depending on the rock type or the orientation of stratification in the rock, and on rates of erosion, inselbergs may be sharp-crested, plateau-like, or loaf-shaped (steep sides and a rounded crest). Inselbergs with a loaf geometry, as exemplified by Uluru (Ayers Rock) in central Australia (figure above c), are also known as bornhardts.

Desert Pavement 

Desert pavement, and a hypothesis for how it forms by building up a soil from below.

In many locations, the desert surface resembles a tile mosaic in that it consists of separate stones that fit together tightly, forming a fairly smooth surface layer above a soil composed of silt and clay. Such natural mosaics constitute desert pavement (figure above a, b). Typically, desert varnish coats the top surfaces of the stones forming desert pavement. Recently, researchers have suggested that pavements form when windblown dust slowly sifts down onto the stones and then washes down between the stones. In this model, the pavement is “born at the surface,” meaning that the stones forming the pavement were never buried, but have been progressively lifted up as soil collects and builds up beneath (figure above c). Over time, the rocks at the surface crack, perhaps due to differential heating by the desert sun. Sheetwash, during downpours, may wash away fine sediment between fragments, and when soils dry and shrink between storms, the clasts settle together, locking into a stable, jigsaw-like arrangement.

Stony Plains and Pediments 

The coarse sediment eroded from desert mountains and ridges washes into adjacent lowlands and builds out to form gently sloping alluvial fans. The surfaces of these gravelly piles are strewn with pebbles, cobbles, and boulders, and are dissected by dry washes (wadis or arroyos). Portions of these stony plains may evolve into desert pavements. 

When travelers began trudging through the desert of the southwestern United States during the 19th century, they found that in many locations the wheels of their wagons were rolling over flat or gently sloping bedrock surfaces. These bedrock surfaces extended outward like ramps from the steep cliffs of a mountain range on one side, to the alluvium-filled valleys on the other (see 2nd figure). Geologists now refer to such surfaces as pediments. Pediments develop when sheetwash during floods carries sediment away from the mountain front, during mountain-front retreat. The moving sediment grinds away the bedrock that it tumbles over.

Seas of Sand: The Nature of Dunes 

The types of sand dunes and the cross beds within them.

A sand dune is a pile of sand deposited by a moving current. Dunes in deserts may start to form where sand becomes trapped on the windward side of an obstacle, such as a rock or a shrub. Gradually, the sand builds downwind into the lee of the obstacle. Dunes display a variety of shapes and sizes, depending on the character of the wind and the sand supply (figure above a). Where the sand is relatively scarce and the wind blows steadily in one direction, beautiful crescents called barchan dunes develop, with the tips of the crescents pointing downwind. If the wind shifts direction frequently, a group of crescents pointing in different directions overlap one another, creating a constantly changing star dune. Where enough sand accumulates to bury the ground surface completely, and only moderate winds blow, sand piles into simple, wave-like shapes called transverse dunes. The crests of transverse dunes lie perpendicular to the wind direction. Strong winds may break through transverse dunes and change them into parabolic dunes whose ends point in the upwind direction. Finally, if there is abundant sand and a strong, steady wind, the sand streams into longitudinal dunes, whose axis lies parallel to the wind direction.

In a sand dune, sand saltates up the windward side of the dune, blows over the crest of the dune, and then settles on the steeper, lee face of the dune. The slope of this face attains the angle of repose, the slope angle of a free-standing pile of sand. As sand collects on this surface, it eventually becomes unstable and slides down the slope, so geologists refer to the lee side of a dune as the slip face. As more and more sand accumulates on the slip face, the crest of the dune migrates downwind, and former slip faces become preserved inside the dune. In cross section, these slip faces appear as cross beds (figure above b, c).
Figures credited to Stephen Marshak.

Deposition in Deserts

Deposition in Deserts

kacimo samy   6:24 ص

Deposition in Deserts

We've seen that erosion relentlessly eats away at bedrock and sediment in deserts. Where does the debris go? Below, we examine the various desert settings in which sediment accumulates.

Talus Aprons 

Production and transportation of debris and sediment in deserts.

Over time, joint-bounded blocks of rock break off ledges and cliffs on the sides of hills. Under the influence of gravity, the resulting debris tumbles downslope and accumulates as talus, a pile of debris at the base of a hill. Talus can survive for a long time in desert climates, so we typically see aprons of talus  fringing the bases of cliffs in deserts (figure above a).

Alluvial Fans 

Flash floods can carry sediment downstream in a steep-walled canyon. In some cases, the water contains so much sediment that it is best called a debris flow. When the turbulent water flows out into a plain at the mouth of a canyon, it spreads out over a broader surface and slows. As a consequence, sediment in the water settles out. The channel that emerges from the mountains tends to subdivide into a number of subchannels or distributaries that diverge outward. The network of distributaries spreads the sediment, or alluvium, out into a broad alluvial fan, a wedge- or apron-shaped pile of sediment (figure above b). Alluvial fans emerging from adjacent valleys may merge and overlap along the front of a mountain range, producing an elongate wedge of sediment called a bajada. Over long periods, the sediment of bajadas can fill in adjacent valleys to depths of several kilometres.

Playas and Salt Lakes 

Playas form where a shallow, salty lake dries up.

Water from a flash flood occasionally reaches the centre of an alluvium-filled basin. But if the supply of water is relatively small, the water quickly sinks into the permeable alluvium without accumulating as a standing body. During a particularly large storm or an unusually wet spring, however, a temporary lake may develop over the low part of a basin. During drier times, such desert lakes evaporate entirely, leaving behind a dry, flat, exposed lake bed known as a playa (figure above a, b). Over time, a smooth crust of clay and various salts (halite, gypsum, borax, and other minerals) accumulates on the surface of playas.

Where sufficient water flows into a desert basin, it creates a permanent lake. If the basin is an interior basin, with no outlet to allow water to flow out, the lake becomes very salty, because although its water escapes by evaporation in the desert sun, its salt cannot. The Great Salt Lake, in Utah, exemplifies this process. Even though the streams feeding the lake are fresh enough to drink, their water contains trace amounts of dissolved salt ions. Because the lake has no outlet, these ions have become concentrated in the lake over time, making it even saltier than the ocean.

Deposition from the Wind 

As mentioned earlier, wind carries two kinds of sediment loads a suspended load of dust-sized particles and a surface load of sand. Much of the dust is carried out of the desert and accumulates elsewhere. Sand, however, cannot travel far, and accumulates within the desert in piles called dunes, ranging in size from less than a meter to over 300 m high. In favourable locations, dunes accumulate to form vast “sand seas.”
Figures credited to Stephen Marshak.

The Nature and Locations of Deserts

The Nature and Locations of Deserts

kacimo samy   10:26 ص

What Is a Desert? 

Formally defined, a desert is a region that is so arid (dry) that it supports vegetation on no more than 15% of its surface. In general, desert conditions exist where less than 25 cm of rain falls per year, on average. Because of the lack of water, deserts contain no permanent streams, except for those that bring water in from temperate regions elsewhere.

Note that the definition of a desert depends on a region’s aridity, not on its temperature. Geologists, therefore, distinguish between cold deserts, where temperatures generally stay below about 20C for the year, and hot deserts, where summer daytime temperatures exceed 35C. Cold deserts exist at high latitudes where the Sun’s rays strike the Earth obliquely and thus don’t provide much energy, at high elevations where the air is too thin to hold much heat, or in lands adjacent to cold oceans, where the cold water absorbs heat from the air above. Hot deserts develop at low latitudes where the Sun’s rays strike the desert at a high angle, at low elevations where dense air can hold a lot of heat, and in regions distant from the cooling effect of cold ocean currents. The hottest recorded temperatures on Earth occur in low-latitude, low-elevation  deserts 58C (136F) in Libya and 57C (133F) in Death Valley, California.

Types of Deserts 

Each desert on Earth has unique characteristics of landscape and vegetation that distinguish it from others. Geologists group deserts into five different classes, based on the environment in which the desert forms (figure below). 

Subtropical deserts form because the air that convectively flows downward in the subtropics warms and absorbs water as it sinks.

• Subtropical deserts: Subtropical deserts (such as the Sahara, Arabian, Kalahari, and Australian) form because of the regional  pattern of air circulation in the atmosphere. At the equator, the air becomes warm and humid, for sunlight is intense and water rapidly evaporates from the ocean. The hot, moisture-laden air rises to great heights above the equator. As this air rises, it expands and cools, and can no longer hold so much moisture. Water condenses and falls in downpours that feed the lushness of the equatorial rain forest. The now-dry air high in the troposphere spreads laterally north or south. When this air reaches latitudes of 20 to 30 C, a region called the subtropics, it has become cold and dense enough to sink. Because the air is dry, no clouds form, and intense solar radiation strikes the Earth’s surface. The sinking, dry air becomes denser and heats up, soaking up any moisture present. In the regions swept by this hot air on its journey back to the equator, evaporation rates greatly exceed rainfall rates, so the land becomes parched. 
• Deserts formed in rain shadows: As air flows over the sea toward a coastal mountain range, the air must rise (Fig. 17.3). As the air rises, it expands and cools. The water it contains condenses and falls as rain on the seaward flank of the mountains, nourishing a coastal rain forest. When the air finally reaches the inland side of the mountains, it has lost all its moisture and can no longer provide rain. As a consequence, a rain shadow forms, and the land beneath the rain shadow becomes a desert. A rain-shadow desert can be found east of the Cascade Mountains in the state of Washington.

 The formation of a rain-shadow desert. Moist air rises and drops rain on the coastal side of the range. By the time the air has crossed the mountains, it is dry. 

• Coastal deserts formed along cold ocean currents: Cold ocean water cools the overlying air by absorbing heat, thereby decreasing the capacity of the air to hold moisture. For  example, the cold Humboldt Current, which carries water northward from Antarctica to the western coast of South America, cools the air that blows east, over the coast. The air is so dry when it reaches the coast that rain rarely falls on the coastal areas of Chile and Peru. As a result, this region hosts a desert landscape, including one of the driest deserts in the world, the Atacama (figure below a, b). Portions of this narrow desert received no rain at all between 1570 and 1971. 
• Deserts formed in the interiors of continents: As air masses move across a continent, they lose moisture by dropping rain, even in the absence of a coastal mountain range. Thus, when an air mass reaches the interior of a broad continent, it has become so dry that the land beneath becomes arid. The largest present-day example of such a continental-interior desert, the Gobi, lies in central Asia. 
• Deserts of the polar regions: So little precipitation falls in Earth’s polar regions (north of the Arctic Circle and south of the Antarctic Circle) that these areas are, in fact, arid. Polar regions are dry, in part, for the same reason that the subtropics are dry (the global pattern of air circulation means that the air flowing over these regions is dry), and in part, for the same reason that coastal areas along cold currents are dry (cold air holds little moisture).

 The formation of a coastal desert.

Different regions of the land surface have become deserts at different times in the Earth’s history, because plate movements change the latitude of landmasses, the position of landmasses relative to the coast, and the proximity of landmasses to a mountain range. Because of plate tectonics, some regions that were deserts in the past are temperate or tropical regions now, and vice versa.

Figures credited to Stephen Marshak.

Coastal Problems and Solutions

Coastal Problems and Solutions

kacimo samy   10:38 ص

Coastal Problems and Solutions

Contemporary Sea-Level Changes 

 Future sea-level rise, due to melting of polar ice, would flood many coastal cities.

People tend to view a shoreline as a permanent entity. But in fact, shorelines are ephemeral geologic features. On a time scale of hundreds to thousands of years, a shoreline moves inland or seaward depending on whether relative sea level rises or falls or whether sediment supply increases or decreases. In places where sea level is rising today, shoreline towns will eventually be submerged. For example, the Persian Gulf now covers about twice the area that it did 4,000 years ago. And if present rates of sea-level rise along the East Coast of the United States continue, major coastal cities such as Washington, New York, Miami, and Philadelphia may be inundated within the next millennium (figure above).

Beach Destruction-Beach Protection?

Examples of beach erosion.

In a matter of hours, a storm especially a hurricane can radically alter a landscape that took centuries or millennia to form. The backwash of storm waves sweeps vast quantities of sand seaward, leaving the beach a skeleton of its former self. The surf submerges barrier islands and shifts them toward the lagoon. Waves and wind together rip out mangrove swamps and salt marshes and break up coral reefs, thereby destroying the organic buffer that can protect a coast, leaving it vulnerable to erosion for years to come. Of course, major storms also destroy human constructions: erosion undermines shore-side buildings, causing them to collapse into the sea; wave impacts smash buildings to bits; and the storm surge very high water levels created when storm winds push water toward the shore floats buildings off their foundations (figure above a, b).

But even less-dramatic events, such as the loss of river sediment, a gradual rise in sea level, a change in the shape of a shoreline, or the destruction of coastal vegetation, can alter the balance between sediment accumulation and sediment removal on a beach, leading to beach erosion. In some places, beaches retreat landward at rates of 1 to 2 m per year. 

Techniques used to preserve beaches.

In many parts of the world, beach front property has great value; but if a hotel loses its beach sand, it probably won’t stay in business. Similarly, a harbour can’t function if its mouth gets blocked by sediment. Thus property owners often construct artificial barriers to alter the natural movement of sand along the coast, sometimes with undesirable results. For example, beach-front property owners may build groins, concrete or stone walls protruding perpendicular to the shore, to prevent beach drift from removing sand (figure above a). Sand accumulates on the up-drift side of the groin, forming a long triangular wedge, but sand erodes away on the down-drift side. Needless to say, the property owner on the down-drift side doesn't appreciate this process. Harbour engineers may build a pair of walls called jetties to protect the entrance to a harbour (figure above b). But jetties erected at the mouth of a river channel effectively extend the river into deeper water and thus may lead to the deposition of an offshore sandbar. Engineers may also build an offshore wall called a breakwater, parallel or at an angle to the beach, to prevent the full force of waves from reaching a harbour. With time, however, sand builds up in the lee of the breakwater and the beach grows seaward, clogging the harbor (figure above c). To protect expensive shore side construction, people build seawalls out of riprap (large stone or concrete blocks) or reinforced concrete on the landward side of the beach (figure above d), but during a storm, these can be undermined. 

In some places, people have given up trying to decrease the rate of beach erosion and instead have worked to increase the rate of sediment supply. To do this, they pump sand from farther offshore, or truck in sand from elsewhere to replenish a beach. This procedure, called beach nourishment, can be hugely expensive and at best provides only a temporary fix, for the backwash and beach drift that removed the sand in the first place continue unabated as long as the wind blows and the waves break.

Destruction of Wetlands and Reefs 

Bad cases of beach pollution create headlines. Because of beach drift, garbage dumped in the sea in an urban area may drift along the shore and be deposited on a tourist beach far from its point of introduction. Oil spills, from ships that flush their bilges or from tankers that have run aground or foundered in stormy seas, or from offshore well leaks, have contaminated shorelines at several places around the world. 

The influx of nutrients, from sewage and agricultural run-off, into coastal waters can create dead zones along coasts. A dead zone is a region in which water contains so little oxygen that fish and other organisms within it die. Dead zones form when the concentration of nutrients rises enough to stimulate an algae bloom, for overnight respiration by algae depletes dissolved oxygen in the water, and the eventual death and decay of plankton depletes oxygen even more. One of the world’s largest dead zones occurs in the Gulf of Mexico, offshore of the Mississippi River’s mouth. 

Coastal wetlands and coral reefs are particularly susceptible to changes in the environment, and many of them have been destroyed in recent decades. Their loss both increases a coast’s vulnerability to erosion and, because they provide spawning grounds for marine organisms, disrupts the global food chain. Destruction of wetlands and reefs happens for many reasons. Wetlands have been filled or drained to be converted to farmland, housing developments, resorts, or garbage dumps. Reefs have been destroyed by boat anchors, dredging, the activities of divers, dynamite explosions intended to kill fish, and quarrying operations intended to obtain construction materials. Chemicals and particulates entering coastal water from urban, industrial, and agricultural areas can cause havoc in wetlands and reefs, for these materials cloud water and/or trigger algal blooms, killing filter-feeding organisms. Toxic chemicals in such run-off can also poison plankton and burrowing organisms and, therefore, other organisms progressively up the food chain. 

Global climate change also impacts the health of organic coasts. For example, transformation of once vegetated regions into deserts means that the amount of dust carried by winds from the land to the sea has increased. This dust can interfere with coral respiration and can bring dangerous viruses. A global increase in seawater temperature may be contributing to reef bleaching, the loss of coral colour due to the death of the algae that live in coral polyps. The statistics of wetland and reef destruction worldwide are frightening ecologists estimate that between 20% and 70% of wetlands have already been destroyed, and along some coasts, 90% of reefs have died.

Hurricanes-A Coastal Calamity 

Characteristics and paths of hurricanes in the western North Atlantic.

Global-scale convection of the atmosphere, influenced by the Coriolis effect, causes currents of warm air to flow steadily from east to west in tropical latitudes. As the air flows over the ocean, it absorbs moisture. Because air becomes less dense as it gets warmer, tropical air eventually begins to rise like a balloon. As the air rises, it cools, and the water vapour it contains condenses to form clouds (mists of very tiny water droplets). If the air contains sufficient moisture, the clouds grow into a cluster of large thunderstorms, which consolidate to form a single, very large storm. Because of the Coriolis effect, this large storm evolves into a rotating swirl called a tropical disturbance. If the disturbance remains over warm ocean water, as can happen in late summer and early fall, rising warm moist air continues to feed the storm, fostering more growth. Eventually a spiral of rapidly circulating clouds forms, and the tropical disturbance becomes a tropical depression. Additional nourishment causes the tropical depression to spin even faster and grow broader, until it becomes a tropical storm and receives a name. If a tropical storm becomes powerful enough, it becomes a tropical cyclone. Formally defined, a tropical cyclone is a huge rotating storm, which forms in tropical latitudes, and in which winds exceed 119 km per hour (74 mph). It resembles a giant counter-clockwise spiral of clouds 300 to 1,500 km (930 miles) wide when viewed from space (figure above a). Such a storm is called a hurricane in the Atlantic and eastern Pacific, a typhoon in the western Pacific, and simply a cyclone around Australia and in the Indian Ocean. 

Atlantic hurricanes generally form in the ocean to the east of the Caribbean Sea, though some form in the Caribbean itself. They first drift westward at speeds of up to 60 km per hour (37 mph). They may eventually turn north and head into the North Atlantic or into the interior of North America, where they die when they run out of a supply of warm water (figure above b). Weather researchers classify the strength of hurricanes using the Saffir-Simpson scale, which runs from 1 to 5; somewhat different scales are used for typhoons and cyclones. On the Saffir-Simpson scale, a Category 5 hurricane has sustained winds of >250 km/hr (>156 mph). The highest wind speed ever recorded during a hurricane was in excess of 300 km/hr. 

A typical hurricane (or typhoon or cyclone) consists of several spiral arms extending inward to a central zone of relative calm known as the hurricane’s eye (figure above c). A rotating vertical cylinder of clouds, the eye wall, surrounds the eye. Winds spiral toward the eye, so like an ice skater who spins faster when she brings her arms inward, the winds accelerate toward the interior of the storm and are fastest along the eye wall. Thus, hurricane-force winds affect a belt that is only 15% to 35% as wide as the whole storm (figure above d). On the side of the eye where winds blow in the same direction as the whole storm is moving, the ground speed of winds is greatest, because the storm’s overall speed adds to the rotational motion.

Hurricanes pose extreme danger in the open ocean, because their winds cause huge waves to build, and thus have led to the foundering of countless ships. They also cause havoc in coastal regions, and even inland, though they die out rapidly after moving onshore. The coastal damage happens for several reasons: 

• Wind: Winds of weaker hurricanes tear off branches and smash windows. Stronger hurricanes uproot trees, rip off roofs, and collapse walls. 
• Waves: Winds shearing across the sea surface during a hurricane generate huge waves. In the open ocean, these waves can 
• capsize ships. Near shore, waves batter and erode beaches, rip boats from moorings, and destroy coastal property. 5 Storm surge: Rising air in a hurricane causes a region of extremely low air pressure beneath. This decrease in pressure causes the surface of the sea to bulge upward over an area with a diameter of 60 to 80 km. Sustained winds blowing in an onshore direction build this bulge even higher. When the hurricane reaches the coast, the bulge of water, or storm surge, swamps the land. If the bulge hits the land at high tide, the sea surface will be especially high and will affect a broader area. 
• Rain, stream flooding, and landslides: Rain drenches the Earth’s surface beneath a hurricane. In places, half a meter or more of rain falls in a single day. Rain causes streams to flood, even far inland, and can trigger landslides. 
• Disruption of social structure: When the storm passes, the hazard is not over. By disrupting transportation and communication networks, breaking water mains, and washing away sewage-treatment plants, hurricane damage creates severe obstacles to search and rescue, and can lead to the spread of disease, fire, and looting. 

Nearly all hurricanes that reach the coast cause death and destruction, but some are truly catastrophic. Storm surge from a 1970 cyclone making landfall on the low-lying delta lands of Bangladesh led to an estimated 500,000 deaths. In 1992, Hurricane Andrew leveled extensive areas of southern Florida, causing over $30 billion in damage and leaving 250,000 people homeless. Hurricane Katrina, in 2005, stands as the most destructive hurricane to strike the United States. Let’s look at this storm’s history. 

Hurricane Katrina

The devastation of coastal areas by Hurricane Katrina.

Tropical Storm Katrina came into existence over the Bahamas and headed west. Just before landfall in southeastern Florida, winds strengthened and the storm became Hurricane Katrina. This hurricane sliced across the southern tip of Florida, causing several deaths and millions of dollars in damage. It then entered the Gulf of Mexico and passed directly over the Loop Current, an eddy of summer-heated water from the Caribbean that had entered the Gulf of Mexico. Water in the Loop Current reaches temperatures of 32C (90F), and thus stoked the storm, injecting it with a burst of energy sufficient for the storm to morph into a Category 5 monster whose swath of hurricane-force winds reached a width of  325 km (200 miles). When it entered the central Gulf of Mexico, Katrina turned north and began to bear down on the Louisiana-Mississippi coast. The eye of the storm passed just east of New Orleans, and then across the coast of Mississippi. Storm surges broke records, in places rising 7.5 m (25 feet) above sea level, and they washed coastal communities off the map along a broad swath of the Gulf Coast (figure above a, b). In addition to the devastating wind and surge damage, Katrina led to the drowning of New Orleans. 

To understand what happened to New Orleans, we must consider the city’s geologic history. New Orleans grew on the Mississippi Delta, between the banks of the Mississippi River on the south and Lake Pontchartrain (actually a bay of the Gulf of Mexico) on the north. The older parts of the town grew up on the relatively high land of the Mississippi’s natural levee. Younger parts of the city, however, spread out over the topographically lower delta plain. As decades passed, people modified the surrounding delta landscape by draining wetlands, by constructing artificial levees that confined the Mississippi River, and by extracting groundwater. Sediment beneath the delta compacted, and the delta’s surface has been starved of new sediment, so large areas of the delta sank below sea level. Today, most of New Orleans lies in a bowl-shaped depression as much as 2 m (7 feet) below sea level the hazard implicit in this situation had been recognized for years (figure above c). 

The winds of Hurricane Katrina ripped off roofs, toppled trees, smashed windows, and triggered the collapse of weaker buildings, but their direct consequences were not catastrophic. However, when the winds blew storm surge into Lake Pontchartrain, its water level rose beyond most expectations and pressed against the system of artificial levees and flood walls that had been built to protect New Orleans. Hours after the hurricane 

eye had passed, the high water of Lake Pontchartrain found a weakness along the floodwall bordering a drainage canal and pushed out a section. Breaks eventually formed in a few other locations as well. So, a day after the hurricane was over, New Orleans began to flood. As the water line climbed the walls of houses, brick by brick, residents fled first upstairs, then to their attics, and finally to their roofs. Water spread across the city until the bowl of New Orleans filled to the same level as Lake Pontchartrain, submerging 80% of the city (figure above d).

Floodwaters washed some houses away and filled others with debris (figure above e). The disaster took on national significance, as the trapped population sweltered without food, drinking water, or adequate shelter. With no communications, no hospitals, and few police, the city almost descended into anarchy. It took days for outside relief to reach the city, and by then, many had died and parts of New Orleans, a cultural landmark and major port, had become uninhabitable.
Figures credited to Stephen Marshak.

Causes of Coastal Variability

Causes of Coastal Variability

kacimo samy   9:01 ص

Causes of Coastal Variability 

Coastal variability depends upon following factors.

Plate Tectonic Setting 

The tectonic setting of a coast plays a role in determining whether the coast has steep-sided mountain slopes or a broad plain that borders the sea. Along an active margin, compression squeezes the crust and pushes it up, creating mountains like the Andes along the western coast of South America. Along a passive margin, the cooling and sinking of the lithosphere may create a broad coastal plain, a flatland that merges with the continental shelf, as exists along the Gulf Coast and south-eastern Atlantic coast of the United States. 

Not all passive margins have coastal plains. The coastal areas of some passive margins were uplifted during the rifting event that preceded establishment of the passive margin. For example, highlands formed during rifting border the Red Sea and portions of the Brazilian and southern African coasts. Highlands also rise along the east coast of Australia.

Relative Sea-Level Changes 

Because of sea-level drop during the ice age, there was more dry land.

Sea level, relative to the land surface, changes during geologic time. Some changes develop due to vertical movement of the land. These may reflect plate-tectonic processes or the addition or removal of a load (such as a glacier) on the crust. Local changes in sea level may reflect human activity when people pump out groundwater or oil, for example, the pores between grains in the sediment beneath the ground collapse, and the land surface sinks. Some relative sea-level changes, however, are due to a global rise or fall of the ocean surface. Such eustatic sea-level changes may reflect changes in the volume of mid-ocean ridges. An increase in the number or width of ridges, for example, displaces water and causes sea level to rise. Eustatic sea-level changes may also reflect changes in the volume of glaciers, for glaciers store water on land (figure above). As glaciers grow, sea level falls, and as glaciers shrink, sea level rises. 

Features of emergent coastlines (relative sea level is falling) and submergent coastlines (relative sea level is rising).

Geologists refer to coasts where the land is rising or rose relative to sea level as emergent coasts. At emergent coasts, steep slopes typically border the shore. A series of step-like terraces form along some emergent coasts (figure above a). These terraces reflect episodic changes in relative sea level and/or ground uplift. Those coasts at which the land sinks relative to sea level become submergent coasts (figure above b). At submergent coasts, landforms include estuaries and fjords that  developed when the rising sea flooded coastal valleys. 

Sediment Supply and Climate 

The quantity and character of sediment supplied to a shore affects its character. That is, coastlines where the sea washes sediment away faster than it can be supplied (erosional coasts) recede landward and may become rocky, whereas coastlines that receive more sediment than erodes away (accretionary coasts) grow seaward and develop broad beaches. 

Climate also affects the character of a coast. Shores that enjoy generally calm weather erode less rapidly than those constantly subjected to ravaging storms. A sediment supply large enough to generate an accretionary coast in a calm environment may be insufficient to prevent the development of an erosional coast in a stormy environment. The climate also affects biological activity along coasts. For example, in the warm water of tropical climates, mangrove swamps flourish along the shore, and coral reefs form offshore. The reefs may build into a broad carbonate platform such as appears in the Bahamas today. In cooler climates, salt marshes develop, whereas in arctic regions, the coast may be a stark environment of lichen-covered rock and barren sediment.

Source: Essentials of Geology; book by Stephen Marshak

Coastal Landforms

Coastal Landforms

kacimo samy   1:21 م

Where Land Meets Sea: Coastal Landforms

Tourists along the Amalfi coast of Italy thrill to the sound of waves crashing on rocky shores. But in the Virgin Islands sunbathers can find seemingly endless white sand beaches, and along the Mississippi delta, vast swamps border the sea. Large, dome-like mountains rise directly from the sea in Rio de Janeiro, Brazil, but a 100-m-high vertical cliff marks the boundary between the Nullarbor Plain of southern Australia and the Great Southern Ocean. As these examples illustrate, coasts, the belts of land bordering the sea, vary dramatically in terms of topography and associated landforms. 

Beaches and Tidal Flats 

Characteristics of beach, barrier islands and tidal flats.

For millions of vacationers, the ideal holiday includes a trip to a beach, a gently sloping fringe of sediment along the shore. Some beaches consist of pebbles or boulders, whereas others consist of sand grains (figure above a, b). This is no accident, for waves winnow out finer sediment like silt and clay and carry it to quieter water, where it settles. Storm waves, which can smash cobbles against one another with enough force to shatter them, have little effect on sand, for sand grains can’t collide with enough energy to crack. Thus, cobble beaches exist only where nearby cliffs supply large rock fragments. 

The composition of sand varies from beach to beach because different sands come from different sources. Sands derived from the weathering and erosion of silicic-to- intermediate rocks consist mainly of quartz; other minerals in  these rocks chemically weather to form clay, which washes away  in waves. Beaches made from the erosion of limestone, or of  coral reefs and shells, consist of carbonate sand, including masses of sand-sized chips of shells. And beaches derived from  the erosion of basalt boast black sand, made of tiny basalt grains.

A beach profile, a cross section drawn perpendicular to the shore, illustrates the shape of a beach (figure above c). Starting from the sea and moving landward, a beach consists of a foreshore zone, or intertidal zone, across which the tide rises and falls. The beach face, a steeper, concave-up part of the foreshore zone, forms where the swash of the waves actively scours the sand. The backshore zone extends from a small step, cut by high-tide swash to the front of the dunes or cliffs that lie farther inshore. The backshore zone includes one or more berms, horizontal to landward-sloping terraces that receive sediment only during a storm. 

Geologists commonly refer to beaches as “rivers of sand,” to emphasize that beach sand moves along the coast over time it is not a permanent substrate. Wave action at the shore moves an active sand layer on the sea floor on a daily basis. Inactive sand, buried below this layer, moves only during severe storms or not at all. Longshore drift, discussed earlier, can transport sand hundreds of kilometres along a coast in a matter of centuries. Where the coastline indents landward, beach drift stretches beaches out into open water to create a sand spit. Some sand spits grow across the opening of a bay, to form a baymouth bar (figure above d). 

The scouring action of waves sometimes piles sand up in a narrow ridge away from the shore called an offshore bar, which parallels the shoreline. In regions with an abundant sand supply, offshore bars rise above the mean high-water level and become barrier islands (figure above e), and the water between a barrier island and the mainland becomes a quiet-water lagoon, a body of shallow seawater separated from the open ocean. Though developers have covered some barrier islands with expensive resorts, in the time frame of centuries to millennia, barrier islands are temporary features and may wash away in a storm.

Tidal flats, regions of clay and silt exposed or nearly exposed at low tide but totally submerged at high tide, develop in regions protected from strong wave action (figure above f). They are typically found along the margins of lagoons or on shores protected by barrier islands. Here, sediments accumulate to form thick, sticky layers. 

Rocky Coasts 

Erosion landforms of rocky shorelines.

More than one ship has met its end, smashed and splintered in the spray and thunderous surf of a rocky coast, where bedrock cliffs rise directly from the sea. Lacking the protection of a beach, rocky coasts feel the full impact of ocean breakers. The water pressure generated during the impact of a breaker can pick up boulders and smash them together until they shatter, and it can squeeze air into cracks, creating enough force to push rocks apart. Further, because of its turbulence, the water hitting a cliff face carries suspended sand and thus can abrade the cliff. The combined effects of shattering, wedging, and abrading, together called wave erosion, gradually undercut a cliff face and make a wave-cut notch (figure above a). Undercutting continues until the overhang becomes unstable and breaks away at a joint, creating a pile of rubble at the base of the cliff that waves immediately attack and break up. In this process, wave erosion cuts away at a rocky coast, so that the cliff gradually migrates inland. Such cliff retreat may leave behind a wave-cut bench, or platform, that becomes visible at low tide (figure above b). 

Other processes besides wave erosion break up the rocks along coasts. For example, salt spray coats the cliff face above the waves and infiltrates into pores. When the water evaporates, salt crystals grow and push apart the grains, thereby weakening the rock. Biological processes also contribute to erosion, for plants and animals in the intertidal zone bore into the rocks and gradually break them up. 

Many rocky coasts are irregular with headlands protruding into the sea and embayments set back from the sea. Wave energy focuses on headlands and disperses in embayments, a result of wave refraction. The resulting erosion removes debris at headlands, and sediment accumulates in embayments (figure above c). In some cases, a headland erodes in stages (figure above d). Because of refraction, waves curve and attack the sides of a headland, slowly eating through it to create a sea arch connected to the mainland by a narrow bridge. Eventually the arch collapses, leaving isolated sea stacks just offshore (figure above d). Once formed, a sea stack protects the adjacent shore from waves. Therefore, sand may collect in the lee of the stack, slowly building a tombolo, a narrow ridge of sand that links the sea stack to the mainland.

Estuaries 

The Chesapeake Bay estuary formed when the sea flooded river valleys. The region is sinking relative to other coast areas because it overlies a buried meteor crater.

Along some coastlines, a relative rise in sea level causes the sea to flood river valleys that merge with the coast, resulting in estuaries, where seawater and river water mix. You can recognize an estuary on a map by the dendritic pattern of its river-carved coastline (figure above). Oceanic and fluvial waters interact in two ways within an estuary. In quiet estuaries, protected from wave action or river turbulence, the water becomes stratified, with denser oceanic salt water flowing upstream as a wedge beneath less-dense fluvial freshwater.  In turbulent estuaries, oceanic and fluvial water combine to create nutrient-rich brackish water with a salinity between that of oceans and rivers. Estuaries are complex ecosystems inhabited by unique species of shrimp, clams, oysters, worms, and fish that can tolerate large changes in salinity.

Fjords 

Fjord landscapes form where relative sea-level rise drowns glacially carved valleys.

During the last ice age, glaciers carved deep valleys in coastal mountain ranges. When the ice age came to a close, the glaciers melted away, leaving deep, U-shaped valleys. The water stored in the glaciers, along with the water within the vast ice sheets that covered continents during the ice age, flowed back into the sea and caused sea level to rise. The rising sea filled the deep valleys, creating fjords, or flooded glacial valleys. Coastal fjords are fingers of the sea surrounded by mountains; because of their deep-blue water and steep walls of polished rock, they are distinctively beautiful (figure above).

Coastal Wetlands 

Examples of coastal wetlands.

A flat-lying coastal area that floods during high tide and drains during low tide, but does not get pummeled by intense waves, can host salt-resistant plants and evolve into a coastal wetland. Wetland-dominated shorelines are sometimes called “organic coasts.” Researchers distinguish among different types of coastal wetlands based on the plants they host. Examples include swamps (dominated by trees), marshes (dominated by grasses; figure above a), and bogs (dominated by moss and shrubs). So many marine species spawn in wetlands that as a whole, wetlands account for 10% to 30% of marine organic productivity. 

In tropical or semitropical climates (between 30 north and 30 south of the equator), mangrove trees may become the dominant plant in swamps (figure above b). Some mangrove species form a broad network of roots above the water surface, making the plant look like an octopus standing on its tentacles, and some send up small protrusions from roots that rise above the water and allow the plant to breathe. Dense mangrove swamps counter the effects of stormy weather and thus prevent coastal erosion.

Coral Reefs 

The character and evolution of coral reefs.

Along the azure coasts of Hawaii, visitors swim through colorful growths of living coral. Some corals look like brains, others like elk antlers, still others like delicate fans (figure above a). Sea anemones, sponges, and clams grow on and around the coral. Though at first glance coral looks like a plant, it is actually a colony of tiny invertebrates related to jellyfish. An individual coral animal, or polyp, has a tubelike body with a head of tentacles. 

Coral polyps secrete calcite shells, which gradually build into a mound of solid limestone whose top surface lies from just below the low-tide level down to a depth of about 60 m. At any given time, only the surface of the mound lives the mound’s interior consists of shells from previous generations of coral. The realm of shallow water underlain by coral mounds, associated organisms, and debris comprises a coral reef. Reefs absorb wave energy and thus serve as a living buffer zone that protects coasts from erosion. Corals need clear, well-lit, warm (18–30C) water with normal oceanic salinity, so coral reefs grow only along clean coasts at latitudes of less than about 30 (figure above b). 

Marine geologists distinguish three different kinds of coral reef, on the basis of their geometry (figure above c). A fringing reef forms directly along the coast, a barrier reef develops offshore, and an atoll makes a circular ring surrounding a lagoon. As Charles Darwin first recognized back in 1859, coral reefs associated with islands in the Pacific start out as fringing reefs and then later become barrier reefs and finally atolls. This progression reflects the continued growth of the reef as the island around which it formed gradually sinks. Eventually, the reef itself sinks too far below sea level to remain alive and becomes the cap of a flat-topped seamount known as a guyot.

Recognizing Depositional Environments

Recognizing Depositional Environments

kacimo samy   4:44 ص

How Do We Recognize Depositional Environments? 

Geologists refer to the conditions in which sediment was deposited as the depositional environment. Examples include beach, glacial, and river environments. To identify depositional environments, geologists, like crime scene investigators, look for clues. Detectives may seek fingerprints and bloodstains to identify a culprit. Geologists examine grain size, composition, sorting, bed-surface marks, cross bedding, and fossils to identify a depositional environment. Geological clues can tell us if the sediment was deposited by ice, strong currents, waves, or quiet water, and in some cases can provide insight into the climate at the time of deposition. With experience, geologists can examine a succession of beds and determine if it accumulated on a river floodplain, along a beach, in shallow water just offshore, or on the deep ocean floor.

Let’s now explore some examples of different depositional environments and the sediments deposited in them, by imagining that we are taking a journey from the mountains to the sea, examining sediments as we go. We will see that geologists distinguish among three basic categories of depositional environments: terrestrial, coastal, and marine.

Terrestrial (Nonmarine) Sedimentary Environments 

We begin our exploration with terrestrial depositional environments, those that develop inland, far enough away from the shoreline that they are not affected by ocean tides and waves. The sediments settle on dry land, or under and adjacent to freshwater.  

In some settings, oxygen in surface water or groundwater reacts with iron to produce rust-like iron-oxide minerals in terrestrial sediments, which give the sediment an overall reddish hue. Strata with this hue are informally called redbeds.

Glacial environments 

High in the mountains, where it’s so cold that more snow collects in the winter than melts away,  glaciers rivers or sheets of ice develop and slowly flow. Because ice is a solid, it can move sediment of any size. So as a glacier moves down a valley in the mountains, it carries along all the sediment that falls on its surface from adjacent cliffs or gets plucked from the ground at its base or sides. At the end of the glacier, where the ice finally melts away, the sediment that had been in or on the ice accumulates as “glacial till” (a in figure above). Till is unsorted and unstratified it contains clasts ranging from clay size to boulder size all mixed together.

Mountain stream environments

As we walk down beyond the end of the glacier, we enter a realm where turbulent streams rush downslope in steep-sided valleys. This fast-moving water has the power to carry large clasts; in fact, during floods, boulders and cobbles can tumble down the stream bed. Between floods, when water flow slows, the largest clasts settle out to form gravel and boulder beds, while the stream carries finer sediments like sand and mud farther downstream (b in figure above). Sedimentary deposits of a mountain stream would, therefore, include breccia and  conglomerate.

Alluvial-fan environments

Our journey now takes us to the mountain front, where the fast-moving stream empties onto a plain. In arid regions, where there is not enough water for the stream to flow continuously, the stream deposits its load of sediment near the mountain front, producing a wedge-shaped apron of gravel and sand called an alluvial fan  (c in figure above). Deposition takes place here because when the stream pours from a canyon mouth and spreads out over a broader region, friction with the ground causes the water to slow down, and slow-moving water does not have the power to move coarse sediment. The sand here still contains feldspar grains, for these have not yet weathered into clay. Alluvial-fan sediments become arkose and conglomerate.

Sand-dune environments

If the climate is very dry, few plants can grow and the ground surface lies exposed. Strong winds can move dust and sand. The dust gets carried away, and the resulting well-sorted sand can accumulate in dunes. Thus, thick layers of well-sorted sandstone, in which we can find large cross beds, are relicts of desert sand-dune environments (d in figure above).

River (fluvial) environments

In climates where streams flow, we find several distinctive depositional environments. Rivers transport gravel, sand, silt, and mud. The coarser sediments tumble along the bed in the river’s channel and collect in cross-bedded, rippled layers while the finer sediments drift along, suspended in the water. This fine sediment settles out along the banks of the river, or on the floodplain, the flat land on either side of the river that is covered with water only during floods. On the floodplain, mud layers dry out between floods, leading to the formation of mud cracks. River sediments lithify to form sandstone, siltstone, and shale. Typically, the coarser sediments of channels are surrounded by layers of fine-grained floodplain deposits, so in cross section, the channel has a lens-like shape (e in figure above). Geologists commonly refer to river deposits as fluvial sediments, from the Latin word fluvius, for river.

Lake environments

In temperate climates, where water remains at the surface throughout the year, lakes form. In lakes, the relatively quiet water can’t move coarse sediment; any coarse sediment brought into the lake by a stream settles out at the stream’s outlet. Only fine clay makes it out into the centre of the lake, where it settles to form mud on the lake bed. Thus, lake sediments typically consist of finely  laminated shale (f in figure above). 

At the mouths of streams that empty into lakes, small deltas may form. A delta is a wedge of sediment that accumulates where moving water enters standing water. Deltas were so named because the map shape of some deltas resembles the Greek letter delta ($), as we discuss further in Chapter 14. In 1885, an American geologist named G. K. Gilbert showed that such deltas contain three components (figure above): topset beds composed of gravel, foreset beds of gravel and sand, and silty bottomset beds.

Coastal and Marine Environments 

Along the seashore, a variety of distinct coastal environments occur; the character of each reflects the nature of the sediment supply and the climate. Marine environments start at the high-tide line and extend offshore, to include the deep ocean floor. The type of sediment deposited at a location depends on the climate, water depth, and whether or not clastic grains are available.

Marine delta deposits

After following the river downstream for a long distance, we reach its mouth, where it empties into the sea. Here, the river builds a delta of sediment out into the sea. River water stops flowing when it enters the sea, so sediment settles out. Large deltas are much more complex than the lake examples that Gilbert studied, for they include many different sedimentary environments including swamps, channels, floodplains, and submarine slopes. Sea-level changes may cause the positions of the different environments to move with time. Thus, deposits of an ocean-margin delta produce a great variety of sedimentary rock types (a in figure above).

Coastal beach sands

Now we leave the delta and wander along the coast. Oceanic currents transport sand along the coastline. The sand washes back and forth in the surf, so it becomes well sorted (waves winnow out silt and clay) and well rounded, and because of the back-and-forth movement of ocean water over the sand, the sand surface may become rippled (b in figure above). Thus, if you find well-sorted, medium grained sandstone, perhaps with ripple marks, you may be looking at the remnants of a beach environment.

Shallow-marine clastic deposits

From the beach, we proceed offshore. In deeper water, where wave energy does not stir the sea floor, finer sediment can accumulate. Because the water here may be only meters to a few tens of meters deep, geologists refer to this depositional setting as a shallow-marine environment. Clastic sedimentary layers that accumulate in this environment tend to be fine-grained, well-sorted, well rounded silt, and they are inhabited by a great variety of organisms such as mollusks and worms. Thus, if you see beds of siltstone and mudstone containing marine fossils, you may be looking at shallow-marine clastic deposits.

Shallow-water carbonate environments

In shallow marine settings relatively free of clastic sediment, warm, clear, nutrient-rich water hosts an abundance of organisms. Their shells, which consist of carbonate minerals, make up most of the sediment that accumulates (a and b in figure above). The nature of carbonate sediment depends on the water depth. Beaches collect sand composed of shell fragments; lagoons (protected bodies of quiet water) are sites where carbonate mud accumulates; and reefs consist of coral and coral debris. Farther offshore of a reef, we can find a sloping apron of reef fragments. Shallow-water carbonate environments transform into various kinds of limestone.

Deep-marine deposits

We conclude our journey by sailing far offshore. Along the transition between coastal regions and the deep ocean, turbidity currents deposit graded beds. In the deep-ocean realm, only fine clay and plankton provide a source for sediment. The clay eventually settles out onto the deep-sea floor, forming deposits of finely laminated mudstones, and plankton shells settle to form chalk (from calcite shells; a and b in figure above) or chert (from siliceous shells). Thus, deposits of mudstone, chalk, or bedded chert indicate a deepmarine origin.

Geology of the Antarctica

Geology of the Antarctica

kacimo samy   1:05 ص

ANDRILL Antarctica and the Antarctic Geological... by Learning Geology

Nonmarine Depositional Systems

Nonmarine Depositional Systems

kacimo samy   7:20 ص

The early sequence model of emphasized the accumulation of fluvial deposits during the late highstand phase of the sea-level cycle. The model suggested that the longitudinal profile of rivers that are graded to sea level would shift seaward during a fall in base level, and that this would generate accommodation for the accumulation of nonmarine sediments. This idea was examined critically and was described scenarios where this may and may not occur. The response of fluvial systems to changes in base level was examined from a geomorphic perspective in greater detail, and the sequence stratigraphy of nonmarine deposits was critically reviewed. In general, the importance of base-level change diminishes upstream. In large rivers, such as the Mississippi, the evidence from the Quaternary record indicates that sea-level changes affect aggradation and degradation as far upstream as the region of Natchez, Mississippi, about 220 km upstream from the present mouth. Farther upstream than this, source-area effects, including changes in discharge and sediment supply, resulting from tectonism and climate change, are much more important. In the Colorado River of Texas base-level influence extends about 90 km upstream, beyond which the river has been affected primarily by the climate changes of Neogene glaciations. At some point upstream rivers become completely independent of higher order relative changes in base level, and are responding to a tectonically controlled long-term average base level of erosion. The response of river systems to climate change is complex. Cycles of aggradation and degradation in inland areas may be driven by changes in discharge and sediment load, which are in part climate dependent. These cycles may be completely out of phase with those driven primarily by base-level change. The sequence boundary is commonly an incised valley eroded during the falling stage of the base-level cycle. This valley is filled with fluvial or estuarine deposits during the lowstand to transgressive part of the cycle, with the thickness and facies composition of these beds determined by the balance between the rates of subsidence, base-level change and sediment supply. Away from the incised valley, on interfluve areas, the sequence boundary may be marked by well-developed paleosol horizons. It is a matter of debate whether the fluvial fill of an incised valley should be assigned to the lowstand or the transgressive systems tract. The shape of the sea-level curve and the timing of these deposits relative to this curve are usually not knowable, and so this is a somewhat hypothetical argument. 

Transgression is commonly indicated by the appearance of abundant tidal influence in the fluvial succession. Sigmoidal cross bedding, tidal bedding (wavy, flaser and lenticular bedding), oyster beds and brackish to marine trace fossils are all typical indicators of tidal-marine environments. The transition from fluvial to tidal is typically diachronous, and the filling of the incised valley changes from aggradational to retrogradational. Inland from tidal influence, the change from the lowstand to the transgressive phase may be marked by a change in fluvial style or by the development of coal beds. Coal commonly occurs during an initial increase in accommodation, before this is balanced by an increase in clastic supply. Within the valleys of major rivers, the increase in accommodation can result in more loose stacking of channel sand bodies and greater preservation of overbank fines. Changes in fluvial style are also common, with braided rivers typifying lowstand systems and anastomosed or meandering rivers common during times of high rate of generation of accommodation, as during the transgressive phase of the base-level cycle. 

Meandering river

Anastamous

A highstand systems tract develops when base-level rise slows, and the rate of generation of accommodation space decreases to a minimum. There are two possible depositional scenarios for this phase of sequence development. Retrogradation of the river systems during transgression will have led to reduced slopes, and a low-energy landscape undergoing slow accumulation of floodplain deposits, limited channel aggradation, and closely-spaced, well-developed soil profiles. Given no change in source-area conditions, however, the sediment supply into the basin will continue, and vigorous channel systems will eventually be re-established. Under these conditions, channel bodies will form that show reduced vertical separation relative to the TST, leading to lateral amalgamation of sandstone units and high net-to-gross sandstone ratios. Basinward progradation of coastal depositional systems leads to downlap of deltaic and barrier-strandplain deposits onto the maximum flooding surface. It seems likely that the HST will be poorly represented in most nonmarine basins, because the highstand is usually followed by the next cycle of falling base level, which may result in the removal by subaerial erosion of much or all of the just-formed HST deposits. A minor increase in the sand-shale ratio immediately below the sequence boundary may be the only indication of the highstand phase, as in the Castlegate Sandstone of Utah. 

Care must be taken to evaluate all the evidence in interpreting such data as net-to-gross sandstone ratios. Changes in this parameter may not always be attributable to changes in the rate of generation of accommodation space. An increase in the proportion of channel sandstones in a section seems to have been related not to changes in the rate of generation of accommodation space, but to increased sediment runoff resulting from increased rainfall. In the case of sequences driven by orbital forcing mechanisms, where both base-level change and climate change may be involved, unravelling the complexity of causes and effects is likely to be a continuing challenge. Rising base level is the main control on the rate and style of channel stacking, the rate of generation of accommodation is small during transgression in inland areas while the coastline is still distant, and increases only once transgression has brought the coastline farther inland where the effect of baselevel rise on the lower reaches of the river produces a more rapid increase in accommodation. The rate of generation of accommodation is greater during the highstand than during transgression, and results in low net-to-gross sandstone ratios. However, this line of reasoning omits the influence of upstream factors, and must therefore not be followed dogmatically. One must also be cautious in using systems-tract terminology derived from marine processes for the labeling of nonmarine events. There may be a considerable lag in the transmission of a transgression upstream to inland positions by the process of slope reduction, aggradation and tidal invasion. The inland reaches of the river will not “know” that a transgression is occurring, and it is questionable, therefore, whether the deposits formed inland during the initial stages of the marine transgression should be included with the TST. An alternative terminology for the standard systems tract terms used for marine basins because, obviously, terms that include such words as transgressive, highstand, etc., are inappropriate for basins that are entirely nonmarine. For falling stage and lowstand deposits the term degradational systems tract, for transgressive deposits, transitional systems tract, and for highstand deposits, aggradational systems tract. These terms provide analogous ideas regarding changes in accommodation and sediment supply and their consequences for depositional style. 

In marine basins this will be marine base level (sea level). In inland basins it will be the lip or edge of a basin through which the trunk river flows out of the basin. The buffer is the zone of space above and below the current graded profile which represents the range of reactions that the profile may exhibit given changes in upstream controls, such as tectonism or climate change, that govern the discharge and sediment load of the river. For example, tectonic uplift may increase the sediment load, causing the river to aggrade towards its upper buffer limit. A drop in the buttress, for example as a result of a fall in sea level, may result in incision of the river system, but if the continental shelf newly exposed by the fall in sea level has a similar slope to that of the river profile, there may be little change in the fluvial style of the river. In any of these cases, the response of the river system is to erode or aggrade towards a new dynamically maintained equilibrium profile that balances out the water and sediment flux and the rate of change in accommodation. The zone between the upper and lower limits is the buffer zone, and represents the available (potential) preservation space for the fluvial system. Nowhere within coastal fluvial systems is there a single erosion surface that can be related to lowstand erosion. Such surfaces are continually modified by channel scour, even during transgression, because episodes of channel incision may reflect climatically controlled times of low sediment load, which are not synchronous with changes in base level. This is particularly evident landward of the limit of base-level influence. Post-glacial terraces within inland river valleys reveal a history of alternating aggradation and channel incision reflecting climate changes, all of which occurred during the last post-glacial rise in sea level. A major episode of valley incision occurred in Texas not during the time of glacioeustatic sea-level lowstand, but at the beginning of the postglacial sea-level rise, which commenced at about 15 ka. The implications of this have yet to be resolved for inland basins where aggradation occurs (because of tectonic subsidence), rather than incision and terrace formation. However, it would seem to suggest that no simple relationship between major bounding surfaces and base level change should be expected. Fluvial processes in relationship to glacially controlled changes in climate and vegetation, based on Dutch work. These studies, and those in Texas, deal with periglacial regions, where climate change was pronounced, but the areas were not directly affected by glaciation. A major period of incision occurred during the transition from cold to warm phases because run off increased while sediment yield remained low. Vegetation was quickly able to stabilize river banks, reducing sediment delivery, while evapotranspiration remained low, so that the run off was high. Fluvial styles in aggrading valleys tend to change from braided during glacial phases to meandering during interglacials. With increasing warmth, and consequently increasing vegetation density, rivers of anastomosed or meandering style tend to develop, the former particularly in coastal areas where the rate of generation of new accommodation space is high during the period of rapidly rising base level. Valley incision tends to occur during the transition from warm to cold phases. Reduced evapotranspiration consequent upon the cooling temperatures occurs while the vegetation cover is still substantial. Therefore run off increases, while sediment yield remains low. With reduction in vegetation cover as the cold phase becomes established, sediment deliveries increase, and fluvial aggradation is reestablished.

Channel incision

It is apparent that fluvial processes inland and those along the coast may be completely out of phase during the climatic and base-level changes accompanying glacial to interglacial cycles. Within a few tens of kilometres of the sea, valley incision occurs at times of base level lowstand, during cold phases, but the surface may be modified and deepened during the subsequent transgression until it is finally buried. Inland, major erosional bounding surfaces correlate to times of climatic transition, from cold to warm and from warm to cold, that is to say during times of rising and falling sea level, respectively. The Dakota Group of north-east New Mexico and southeast Colorado provides a good example of an internally architecturally complex fluvial unit generated by a combination of upstream tectonic controls and downstream sea-level cycles. At the coastline, progradation and retrogradation creating three sequences were caused by sea-level cycles on a 105-year time scale. Each of these sequences can be traced up dip towards the west,where they are composed of repeated cycles of aggradational valley-fill successions and mutually incised scour surfaces. These cycles reflect autogenic channel shifting within the limited preservation space available under conditions of modest, tectonically-generated accommodation. This space is defined by a lower buffer set by maximum local channel scour, and an upper buffer set by the ability of the river to aggrade under the prevailing conditions of discharge and sediment load.