Set #4
Mineral Soils conditioned by Topography
Major landforms in alluvial
lowlands
Fluvisols
Gleysols
Major landforms in mountains and formerly
glaciated regions
Leptosols
Regosols
Major landforms in alluvial lowlands
In the present context, the term `lowlands' refers to flat and level wetlands, commonly situated at or near sea level and consisting of mainly Holocene sediments that are too young to be strongly weathered. Such lowlands are prominent in landscapes with a fluvial, lacustrine, marine or glacial signature.
Tectonic processes during the past 105-106 years had a decisive influence on the formation of lowlands, in particular in subsidence areas such as sedimentary basins and `grabens' where rivers lose much of their transport capacity as the river gradient decreases. Loss of gradient forces many rivers to deposit their bed load and suspension load in their lower reaches.
Prominent sea-level changes and climate fluctuations occurred during the past 103-105 years. Sea level changes strongly influence river behaviour and coastal development. Climate changes affect the discharge and sediment load characteristics of rivers and also the development of a protective vegetation cover that stabilises the landscape. Changes in sea level and/or climate induce changes in river type, channel pattern and sediment sequences. Climate changes during the (recent) Quaternary had a particularly strong impact. In (now) temperate regions, glacial climate was much drier and colder than today. Wet-tropical areas were drier as well. A belt of permafrost surrounded the large continental ice sheets and sea levels were some 120 m and perhaps even 135 m lower than today. There are strong indications that mass wasting processes such as frost weathering and `gelifluction' in high-relief areas were more prominent than today and that consequently the sediment loads of rivers were greater than at present. Once most of the Pleistocene ice cover had disappeared, some 6,000 years ago, the rate of sea level rise decreased and coasts adapted to the higher sea level. Today, (Pleistocene) glaciers that have long disappeared still influence some 15 percent of the land surface. It is impossible to understand the complexity of `landforms in lowlands' without referring to their present climate and their climatic history.
In the following discussion, the major landforms in alluvial lowlands will be aggregated to two broad categories:
River systems are complex systems and a catchment area will normally harbour a variety of landforms. These vary with the nature of the river and with the position in the river system. In most river systems, three zones can be distinguished (see Figure 1):
Figure 1
The upper, middle and lower reaches of a river course
Tectonic uplift or subsidence, climate changes and sea level changes are external controls on river behaviour. The main variables that control a fluvial system are gradient, discharge and sediment supply to the drainage network. They determine whether a river incises or aggrades and which channel type is most efficient in dissipating kinetic energy. Three main river types or (better) channel patterns can be distinguished that differ in the number of channels and channel sinuosity:
Note that river systems are highly dynamic and that channel patterns adapt to changes in regional or local conditions. This means that different channel patterns and landforms may exist in the same fluvial system. For instance, the Yukon River in Alaska has a braided pattern in its upper reach, but assumes a meandering pattern in the lower reach in response to its sharply decreased gradient.
Braided rivers have one single channel of low sinuosity and high gradient, with multiple `thalwegs' and bars. The bars are bare or vegetated sediment islands around which flow is diverted in the channel thalwegs. During times of maximum discharge, the channel is completely inundated and most of the bars conduct water. In times of low discharge, multiple thalwegs and bars reappear within the channel, which lets the braided river resemble a multiple-channel system. The pattern of bars and thalwegs may change profoundly during a flood stage, with vegetated bars remaining in place longer than barren ones.
Braided rivers occur in areas with
Such conditions are common in areas with sparse vegetation and torrential rains, e.g. in arid and semi-arid regions and in permafrost regions where sudden peaks in water and sediment discharge occur in spring when snow and ice melt. Braided rivers are particularly common near glacier fronts and in volcanic areas where glacier melting or volcanic eruptions are associated with sharp peaks in sediment supply.
Deposits by braided rivers contain alternating areas (lenses) of coarse gravel and sand with only minor inclusions of finer sediment. Gravel or coarse sand is preferentially deposited at the downstream ends of bars whereas the sides of bars are eroded as the thalwegs shift. Some braided rivers carry mainly sand and silt, e.g. the Yellow River in China and the Brahmaputra river in Bangladesh. Braided river deposits may become covered with a layer of fine-grained sediments (`overbank deposits', also referred to as `Hochflutlehm') once the river incises into deeper layers by deepening a major thalweg. Its former floodplain has then become a terrace. Overbank deposits are further thickened during exceptionally high floods of the incised and (often) already meandering river. Former gravel bars and gullies can normally still be seen in the micro-relief (less than a metre high) of areas underlain by braided river deposits. Figure 2 shows the basic elements of a typical braided river system.
Some authors consider `alluvial fans' as a special type of braided river. Nichols (1999) defined them as cones of detritus that form at a break in slope at the edge of an alluvial plain. Alluvial fans are not uncommon in humid regions but truly classical examples are found in arid regions such as Death Valley in Arizona, USA. Alluvial fans are therefore discussed in the chapter on landforms in arid regions. Some highlights:
Alluvial fans have a steep gradient and their unconfined channels shift frequently over the depositional cone. They develop at places where river gradients decrease sharply, for instance at mountain fronts where a tributary stream leaves the mountains and enters a level alluvial plain. The sediment load of the feeding river can no longer be carried and most of it is dropped right at the entrance to the plain. This rapidly blocks the channel, which then sweeps left and right to obviate the obstacle. Deposits tend to be coarser on the `proximal' part of the fan (close to the apex) than on the `distal' part (far into the plain). In contrast with `normal' braided river systems, mudflows are common on alluvial fans; their deposits are most numerous near the apex of the fan.
Meandering rivers consist of one single channel and thalweg, with a low gradient and high sinuosity. Sediments are preferentially deposited at the inner sides of meander bends. The channel is bordered by `natural levees' behind which `flood basins' or `backswamps' occur. Laboratory experiments have confirmed that meandering channels are typical of rivers with rather steady discharge rates and sediment loads of relatively fine-grained material (sand, silt, clay). See Figure 3.
Figure 2
Basic elements of the South Saskatchewan River, a typical
braided river. Bar A is driven toward the far bank, forming a protected part of
the channel (slough) in which mud is deposited. B is a complex sand flat.
Source: Blatt, Middleton & Murray, 1972; based on work by D.J. Cant
Meandering rivers are associated with vegetated areas under a humid climate where a dense vegetation cover curbs erosion. All this favours chemical weathering and formation of fine-grained soil material. Surface run-off is minimal and the river is fed by a steady base flow of groundwater. Note, however, that meandering river stretches are also found in cold regions (e.g. the lower Yukon River, Alaska). Meandering rivers do occur in arid environments but their sources are invariably outside the arid region. The rivers Euphrates and Tigris, for instance, originate in the snow-capped Taurus mountain range, Turkey, before they reach the Mesopotamian desert.
The discharge rate of a meandering river may fluctuate considerably over time, in line with seasonal fluctuations in rainfall, snow melt and evapotranspiration but the river runs never dry and frequent floods are uncommon. If the channel is filled to the top of its natural levees, the river is said to be at `bank-full stage'. When more water is supplied than can be held between the levees, the river spills into its basins and floods its backswamp areas.
Deposition of sediments takes place in the channel, on the levees and in the basins. Gravel and coarse sand are normally found on the channel floor (`lag deposits'). Finer sand settles along the inner bends of the river, on so-called `point bars'. During floods, fine sand or silt is deposited on top of the levees, and clay in the basins. (Peat may accumulate there as well, at the lowest/wettest positions in situations of low clastic input.)
Figure 3
The classical meandering stream. Source: Allen,
1964
The whole system of gravely lag deposits, sandy pointbars, sandy/silty levees and clayey backswamps shifts laterally as the meandering channel erodes the outer banks and deposits sediments on the inner-bend point bars. A `fining-upwards' sedimentary sequence may develop in this way. `Crevasse splays' formed when the river breaks through its levees and spills abruptly into a basin area during a flood, do not show the fining-upwards sequence. Eventually, meandering leads to channel cut-offs (so-called `oxbow lakes'), which become filled in with clay and/or peat.
Anastomosing rivers have multiple, interconnected channels that divide and rejoin around relatively stable areas of floodplain (in the case of lowland rivers) or bedrock (in upland streams). Anastomosing rivers differ from multiple-thalweg braided rivers in that the river has only one thalweg without bars. The river gradient is typically very low; the channels have stable banks and channel `avulsion' is common. Avulsion is the breaking of a river (channel) through its natural levees to find a new course in a lower part of the floodplain. The channel downstream of the point of avulsion may become an abandoned `paleo-channel' but in anastomosed river systems old channels normally remain active.
The very low gradient of anatomosed river reaches interferes with water discharge in times of floods: crevasse splays are abundant. These are sandy outwash fans that develop when the river breaks through its levee without forming a new channel. Crevasse deposits are predominantly clayey and silty in texture because the low gradient of the anastomosing river precludes transport of coarser particles. For the same reason, levees consist of fine sand and silt. Basin areas are filled in with clay and peat. Anastomosing rivers are typically found inside or near deltas or areas close to the coast such as the Late Holocene Rhine-Meuse delta in the Netherlands. Anastomosing rivers occur also in rapidly subsiding inland basins; examples are the Lower Magdalena Basin in Colombia and the river Niger south of Timbuktu in Mali.
Coastal lowlands contain a variety of landforms. Most of these are geologically young because the rapid post-glacial sea-level rise stabilised only some 6000 years ago. However, borings suggest that many deposits in fluvial deltaic plains and their marine counterparts had a long and turbulent history that was strongly influenced by sea-level fluctuations and changes in climate. Especially inundated, low-gradient continental shelves have complex histories. Some examples: the courses of the rivers Thames, Rhine-Meuse and Seine extended on the (then exposed) North Sea floor during glacial periods. They were tributaries to a big Channel River system, the delta of which was located far to the west of the present `English Channel' area. This indicates that deltas may shift laterally over tens or even hundreds of kilometres. Today's coastal landforms are quite recent; their formation and properties depend in part on whether fluvial processes interact/interacted with marine processes or not. In the first case, deltas and estuaries are the major landforms. In the second case, we distinguish between depositional and eroding coastlines. Three land-shaping factors are important (apart from global sea level fluctuations):
The combined actions of fluvial and marine processes determine whether or not a depositional body can form at the mouth of a river and what kind of body will form. If the rate of fluvial input of sediment exceeds marine sediment redistribution, the depositional sequence will `prograde' seawards to form a `delta'. If marine redistribution can handle the input of fluvial sediments, a depositional body may develop that is not prograding, but only aggrading: an `estuary'. The actions of waves and tides determine which type of delta, estuary or coastline is formed, and the landforms that occur. By and large, tidal regimes can be divided into three broad categories:
Deltas are prograding depositional bodies that form at the point where a river debauches in a lake or sea. The various sedimentary facies of delta bodies are indicative of one or more of the following (external) factors:
Basically, 3 types of delta are distinguished:
In river-dominated deltas, fluvial processes outweigh the influence of waves and tides. These deltas typically form in a microtidal regime at low coastal-shelf gradients. The low-gradient delta of the river Wolga (Russia) is a good example. The river Wolga drains into the Caspian Sea, which actually is not a sea but a fresh-water lake. Wave and tidal action are all but absent and fluvial processes shape the delta. Another classical example is the delta formed by glacially fed rivers debauching into the former (Late Glacial) Lake Bonneville (first described in 1885). Because no differences exist in the apparent densities of river water and lake water, and because the influence of waves and tides is negligible, a fan-shaped `Gilbert-type delta' is formed. In such river-dominated deltas, the coarsest (sand-sized) material is deposited in the `delta plain' close to the mouth. Finer (silty) sediments accumulate in the submerged `delta slope' and the finest (clay) particles travel farthest, to the `prodelta'. Thus, a gradation of grain sizes evolves across the delta. When the delta progrades under continuing sediment supply, progressively coarser sediments cover (or `onlap' as sedimentologists say) the finer sediments: a coarsening-upwards sedimentary sequence is the result. Such sequences may be tens or even hundreds of metres thick. The fine prodelta sediments are called `bottom-set beds', the sloping delta-front sediments are termed `fore-set beds', and the topmost delta plain sediments are the `top-set beds'.
The Mississippi River delta is a special, river-dominated delta. The considerable discharge and sediment load of the Mississippi River lead to rapidly extending channel levees and transport of suspended sediment far into the Gulf of Mexico. The system expands to form a `birdfoot delta', in which the distribution channels of the delta plain form the `toes' of the foot. They even continue below sea level as sub-aqueous pro-delta channels. Swamp lands and coastal marsh form between the toes; they are the brackish equivalents of backswamps in purely fluvial environments. Occasionally, a channel breaks through its levees and deposits a `crevasse splay': a small-scale fining-upwards sequence within the large-scale coarsening-upwards system. Because the channels transverse a very low-gradient delta plain, avulsions are frequent. When this happens, channels are abandoned and a new channel shifts laterally over the delta-plain: old delta lobes are left and new ones are built. See Figure 5 for illustrating block diagrams.
The second main type of delta is the wave-dominated river delta. Wave-dominated river deltas form in micro-tidal coastal regions where the coastal shelf has a steep gradient. Only then can waves attain the energy needed to attack and dissipate freshly deposited fluvial sediments. (Note that sub-aqueous prodelta channels cannot form!) If the main direction of the wind is perpendicular to the main delta axis, fluvial sediments will be redistributed laterally, in mouth bars and sandy beach ridges parallel to the coast. As a consequence, `wave-influenced delta plains' have commonly cuspate or straight outlines. Lagoons form in-between the beach ridges and aeolian dunes may develop on the ridges themselves. The deltas of the rivers Rhône, Ebro and Nile in the Mediterranean Sea, and of the river Danube in the Black Sea are examples of wave-dominated deltas.
Under a meso-or macrotidal regime, ebb- and flood-tide currents are strong enough to scour fluvial channels and redeposit fluvial sediments in them. If the delta is still prograding, a tide-dominated river delta forms. The tidal currents widen the channels in the delta plain and sediments are deposited in opposite directions. The delta plain itself consists of tidal channels that have the shape of funnels widening in seaward direction, with tidal flats on the overbank areas between channels. In the tropics, the tidal flats are normally colonised by mangroves, which add much organic debris to the fresh sediment. Sandy mouth bars develop at the delta front, perpendicular to the coastline and parallel to the tidal currents. The deltas of the Mekong and Ganges-Brahmaputra river systems are tide-dominated delta plains. The Shatt al Arab delta (combined Tigris-Euphrates system) is a prograding tide-dominated delta in an arid region.
Estuaries are aggrading depositional sedimentary bodies at the mouths of rivers. Many present river mouths became estuaries after the post-glacial sea-level rise inundated the former deltas. When that happened, fluvial channels became subject to tidal influence. Especially in meso- and macro-tidal regimes, the fluvial processes became subordinate to marine redistribution and tide-dominated estuaries could form.
Figure 4
Outline of the
basic structure of deltas
The major landforms in an estuary are tidal channels and creeks, in which predominantly sand is transported and deposited. In between them, tidal mudflats develop as a result of silt and clay deposition. The mudflats are flooded at high tide and fall dry again at low tide, when the suspended load stays behind as `slackwater deposits'. The mudflats are strongly saline, with halophytic vegetation on their most stable parts. Any vegetation helps to trap more sediment, which explains why overgrown mudflats grow faster than barren ones. In temperate climates, grasses are the main colonisers of mudflats whereas in tropical climates mangrove trees, with their air roots are the principal species. Mudflats in arid regions may fall dry for several months each year, which may result in net accumulation of evaporites. The Rhine-Meuse and Thames fluvial systems are examples of estuarine, now tide- dominated river mouths.
Where coastal land is eroded or clastic deposits are reworked without any fluvial influence, a marine coastline will form, shaped exclusively by the combined actions of waves and tides. A coastline demarcates the boundary between net erosion of the exposed earth surface and net deposition in the marine realm. As the sea level has fluctuated strongly during past glacial and interglacial periods, coastal landforms may be inherited and reflect former sea levels. Three types of coastline can be distinguished:
Drowned or uplifted coasts, with landforms conditioned by sea level fluctuations. During the penultimate interglacial (Eemian), the sea level was 6 m higher than it is today. Former coastlines that were equilibrated to this higher sea level indicate that gravely and sandy sediments were deposited. When the sea level fell again, these sediments became exposed to form marine terraces. A similar course of events took place along tectonically or isostatically uplifting coasts. For instance, coral reef terraces are found on the tectonically uplifted Huon Pensinsula of Papua New Guinea. These (terraces) have been dated to the Late Quaternary sea-level fluctuations. Raised beaches of up to 9000 years old are present on the Baltic Shield, which experienced strong isostatic uplift after the Scandinavian ice sheet disappeared.
During the last Pleistocene glaciation (the Weichselian), the sea level was 120 metres lower than today. Large parts of continental shelves fell dry. As the base levels of erosion and ground water levels were lower as well, rivers in upland areas could form deeply incised river valleys. The lower reaches of these river valleys became inundated with seawater during the subsequent post-glacial sea level rise, which gave coasts a variety of shapes. `Fjord coasts' formed in glaciated regions when the glaciers retreated from their U-shaped valleys that were inundated when the sea level rose. `Ria coasts' and `channel coasts' were formed in non-glaciated areas where sediment deposition in the lower river reaches was insufficient to match the rise in sea level and keep the sea out of the valleys. A ria coast (e.g. SW-Ireland) has valleys perpendicular to the coastline, whereas valleys run parallel to the coastline of a channel coast (e.g. the Dalmatian coast).
Erosional coastlines, with marine erosion as the major agent form where clastic sediment supply is low and wave attack and along-shore currents move detritus away from the coastline. Typically, a `cliff-face' is eroded away along rocky shores whereas the coastline retreats along former depositional beaches. Undermining of the cliff-face by high-energy waves is the main erosion process along rocky shores. Strong winds associated with hurricanes or deep extra-tropical depressions provide the required wave energy. Salt weathering assists in loosening the rocks as sea-salt crystallises in cracks and fissures. Calcium-carbonate dissolution accelerates erosion of calcareous rocks. Prolonged undermining lets entire blocks or coastline sections slump into the sea. The white cliffs of Dover with their `abrasion platforms' covered with flint and limestone pebbles are a good example. Individual `left-over' rocks stand out as arches or stacks along the coast.
Figure 5
Development of delta sequences in the Mississippi delta.
Source: Frazier, 1967
Depositional or constructional coastlines, with adequate sediment supply are stable or prograding coastlines where accumulation of clastic sediments outweighs erosion by the sea. Whether accretion will take place or transgression depends on such factors as along-shore sediment supply, slope of the foreshore and subsidence rate of the basin. All these factors influence sedimentary sequences. The clastic material may be provided by nearby rivers and transported by along-shore currents, or eroding stretches of upstream coastline may be the source of sediment material. Degrading or eroding coral reefs in tropical shallow seas provide limestone weathering and other bioclastic debris. In very arid regions, evaporites make up much of the beach deposits. The relative influences of sediment supply, waves and tides determine the shape of depostional coastal landforms. The tidal range is (again) important: coasts with a tidal amplitude of 2 metres or less are called `microtidal', coasts with 2 to 4 metres tidal amplitude are `mesotidal' and coasts with a tidal amplitude of more than 4 metres are `macrotidal' coasts.
Wave action rather than tidal action shapes microtidal coasts. Waves rolling up to the coast produce a surf (wash and backwash) resulting in net transport of sediment towards the coast. Compensating currents parallel to the shore are known as `longshore currents'. If the waves approach the coast under an angle, longshore currents and `beach drift' move sediment parallel to the coast. Wave-dominated microtidal coasts can have beach deposits attached to the hinterland or deposits can be separated from the hinterland by a narrow strip of sea, a `lagoon'.
Characteristic landforms of an attached beach are:
Typical landforms of a detached beach are:
The Dutch central coast, between Hoek van Holland and Den Helder, illustrates the formation of wave-dominated coastal lowland (the present tidal amplitude is only 1.5 metres). When the post-glacial sea level rise was still rapid, between 9,000 and 5,000 years BP, coastal formation was dominated by the tides. Then present beach barrier islands moved inland (during the transgression). Tidal inlets behind which clayey tidal flats were formed, the `Calais deposits', separated them. On the landward side, the Calais deposits interfinger with fluvial clays and peat. When the rise in sea level slowed down (5000-2000 years BP), the barrier system closed. A wave-dominated coast formed and grew seawards. The relict beach ridges (`Old Barriers', `Oude Strandwallen') and strand plains between them are particularly well developed in the stretch between the cities of The Hague and Haarlem. The slow rise of the groundwater level during this period created ideal conditions for accumulation of thick layers of peat (the `Holland peat') on top of the Calais deposits. Much of the peat was later mined for fuel, until the Calais deposits were reached. The resulting lakes have later been drained and form the lowest polders of The Netherlands.
Rivers debauching on meso- and macrotidal coasts have mouths that are strongly widened by incoming tides. It was mentioned earlier that estuaries have funnel-shaped channels separated by extensive tidal flats that are commonly brackish or saline. Where tides are high enough to flood large parts of a depositional coast, they build tidal flats. The Dutch-German-Danish Wadden Coast with a mesotidal range of 2-4 metres is a case in point.
Because tidal range conditions coastal development, the landforms found resemble those in estuaries. The tides enter and leave the tidal flats through deep tidal inlets between kilometre-scale barrier islands. Ebb-tidal and flood-tidal deltas may form near the inlets as in lagoons. The tidal channels are permanently submerged and conduct huge masses of seawater and (coarse) sediment at incoming and outgoing tides. The sediment reaches the interior parts of the flats through small gullies and creeks (Dutch: `prielen'). These meander and widen towards their mouths. The `intertidal mudflats' (Dutch: `platen') between mean high water level and mean low water level become submerged twice a day. Fine sandy and clayey material settles on these flats. The highest parts are flooded only during extreme spring tides; they are called `supratidal mudflats' (Dutch: `kwelders', `schorren').
In temperate regions, intertidal flats are normally barren whereas supratidal flats carry a halophytic vegetation of grasses, herbs and shrubs. This vegetation traps sediment when flooded; the resulting deposits are stratified with a typical alternation of fine and coarse layers (Dutch: `kweldergelaagdheid'). Tidal flats in the humid tropics carry a mangrove vegetation already in the intertidal zone and silt up much more rapidly than most supratidal flats in temperate regions. Shifting of creeks and gullies leads to fining-upwards sequences, with fine supratidal and intertidal sediments on top of coarser creek and gully sediments.
Seaward progradation is associated with coarsening-upwards sequences, with the sea bottom consisting of finer material than sediments added and with finer-grained sands near the shoreline than higher up. Alternatively, barriers may shift landwards and sand may be deposited on top of older lagoon material producing coarsening upwards sequences. Interpretation of sedimentary patterns in the field is complicated by the fact that regression and progradation have in places alternated during the Holocene. Figure 6 presents an outline of the main morphological components of the barrier model.
By and large, soils in alluvial lowlands show signs of prolonged wetness and young age. Where sedimentation is still going on, stratified Fluvisols may be expected. Gleysols are found in depression areas that do not receive regular additions of sediment; their profiles testify of a shallow water table during all or most of the year.
In places, Histosols, Arenosols, Solonchaks or/or Solonetz may occur in alluvial lowland areas; these Reference Soil Groups are discussed elsewhere is this text.
Figure 6
Main morphological components of the barrier model.
Source: Blatt,
Middleton & Murray, 1972
Fluvisols (FL)
(with special attention for Thionic Fluvisols)
The Reference Soil Group of the Fluvisols accommodates genetically young, azonal soils in alluvial deposits. The name `Fluvisols' is misleading in the sense that these soils are not confined to river sediments (L. fluvius means `river') but occur also in lacustrine and marine deposits. Many international soil names refer to this group, for example: `Alluvial soils' (Russia, Australia), `Fluvents' (USDA Soil Taxonomy), `Fluvisols' (FAO), Auenböden (Germany) and `Sols minéraux bruts d'apport alluvial ou colluvial' or `Sols peu évolués non climatiques d'apport alluvial ou colluvial' (France).
Definition of Fluvisols#
Soils having
Common soil units:
Thionic*, Histic*, Gelic*, Salic*, Gleyic*, Mollic*, Umbric*, Arenic*, Tephric*, Stagnic*, Humic*, Gypsiric*, Calcaric*, Takyric*, Yermic*, Aridic*, Skeletic*, Sodic*, Dystric*, Eutric*, Haplic*.
# See Annex 1 for key to all Reference Soil Groups.
@ Diagnostic horizon, property or material; see Annex 2 for full definition.
* Qualifier for naming soil units; see Annex 3 for full definition.
Connotation: soils developed in alluvial deposits; from L. fluvius, river.
Parent material: (predominantly) recent, fluvial, lacustrine or marine deposits.
Environment: periodically flooded areas (unless empoldered) of alluvial plains, river fans, valleys and (tidal) marshes, on all continents and in all climate zones.
Profile development: AC-profiles with evidence of stratification; weak horizon differentiation but a distinct Ah-horizon may be present. Redoximorphic features are common, in particular in the lower part of the profile.
Use: Fluvisols are planted to annual crops and orchards and many are used for grazing. Flood control, drainage and/or irrigation are normally required. Thionic Fluvisols suffer from severe soil acidity and high levels of noxious Al-ions.
Fluvisols occur on all continents and in all climates. They occupy some 350 million hectares worldwide of which more than half are in the tropics. Major concentrations of Fluvisols are found:
Major areas of Thionic Fluvisols (`Acid Sulfate Soils') occur in the coastal lowlands of southeast Asia (Indonesia, Vietnam, Thailand), West Africa (Senegal, the Gambia, Guinea Bissau, Sierra Leone, Liberia) and along the north-eastern coast of South America (Venezuela, the Guyanas). Figure 1 shows the world-wide distribution of Fluvisols.
Figure 1
Fluvisols world-wide
Fluvisols occur alongside other `typical' soils of aqueous sedimentary environments such as Arenosols, Cambisols, Gleysols and Solonchaks, and also with weakly developed soils such as Leptosols and Regosols.
Fluvisols are young soils that have `fluvic soil properties'. For all practical purposes this means that they receive fresh sediment during regular floods (unless the land was empoldered) and (still) show stratification and/or an irregular organic matter profile.
Fluvisols in upstream parts of river systems are normally confined to narrow strips of land adjacent to the actual riverbed. In the middle and lower stretches, the flood plain is wider and has the classical arrangement of levees and basins, with coarsely textured Fluvisols on the levees and more finely textured soils in basin areas further away from the river.
In areas with marine sediments, relatively coarse-textured Fluvisols occur on barriers, cheniers, sand flats and crevasse splays; finely textured Fluvisols are found on clayey tidal flats and in chenier plains. Where rivers carry only fine-grained material to the sea, coastal plains (and their Fluvisols) are entirely clayey.
Permanent or seasonal saturation with water preserves the stratified nature of the original deposits but when soil formation sets in, a cambic subsurface horizon will quickly form, transforming the Fluvisol into a Cambisol or Gleysol (depending on the water regime).
The only difference between the parent material of Thionic Fluvisols and that of other Fluvisols is the presence of pyrite (FeS2) in the former.
Formation of pyrite can take place during sedimentation in a marine environment if the following conditions are met:
The mechanism of pyrite accumulation is essentially as follows: Microbes reduce ferric (Fe3+) iron to ferrous (Fe2+) ions, and sulfate (SO42-) to sulfide (S2-) under oxygen-poor conditions (i.e. under water). Organic matter is decomposed in the process, ultimately to bicarbonates. Thus, a potentially acid compound (pyrite) and alkaline compounds (bicarbonates) are formed in an initially neutral system. Tidal flushing removes the alkalinity (HCO3-), and potentially acid pyrite remains behind.
Pyrite is formed in a number of steps, but the `overall' reaction equation reads:
Fe2O3 + 4 SO42- + 8 CH2O + 1/2 O2 = 2 FeS2 + 8 HCO3- + 4 H2O (1)
When pyritic sediment falls dry, oxygen penetrates and pyrite is oxidized, by microbial intervention, to sulfuric acid (H2SO4) and ferric hydroxide (Fe(OH)3). Soluble ferrous sulfate (FeSO4) and (meta-stable) `jarosite' (KFe(SO4)2(OH)6) and/or `schwertmannite' (Fe16O16(SO4)3(OH)10.10H2O) are intermediate products in this process. Jarosite has a typical straw-yellow colour; schwertmannite is yellowish brown. These minerals are easily recognized in the field and are indicative of `Actual' Acid Sulfate Soils, i.e. Thionic Fluvisols with a sulfuric horizon. `Potential' Acid Sulfate Soils contain sulfidic soil material that contains pyrite but has not oxidized to the extent that the soil-pH dropped to a value below 3.5.
The following reaction equations describe successive steps in pyrite oxidation. The mineral `jarosite' will be followed in this text; it is formed at soil-pH < 3.5 if sufficient K+-ions are present:
FeS2 + 7/2 O2 + H2O = Fe2+ + 2 SO42- + 2 H+ (2)
In the presence of carbonates, no lowering of the pH takes place even though much acidity is released.
CaCO3 + 2 H+ = Ca2+ + H2O + CO2 (3)
In strongly calcareous soils or in very dry conditions, gypsum may precipitate:
Ca2+ + SO42- + 2 H2O = CaSO4.2H2O (4)
(gypsum)
In the absence of carbonates, hydrogen produced is not neutralized and the pH of the sediment falls sharply. The soluble ferrous iron produced according to equation (2) is mobile; it oxidizes to jarosite in places where oxygen is present, i.e. in cracks and along root channels:
Fe2+ + 2/3 SO42- + 1/3 K+ + 1/4 O2 + 3/2 H2O = 1/3 jarosite + H+ (5)
With time, the hydrogen ions are removed from the system by percolation or lateral flushing, and jarosite is hydrolyzed to ferric hydroxide:
1/3 jarosite + H2O = 1/3 K+ + Fe(OH)3 + 2/3 SO42- + H+ (6)
The (poorly crystallized) ferric hydroxide will eventually transform to goethite:
Fe(OH)3 = FeOOH + H2O (7)
(goethite)
The stability of the various iron minerals formed in the process of pyrite oxidation is strongly influenced by the redox potential of the soil material (a measure of the quantity of oxygen present) and the soil-pH. Figure 2 presents a diagram of the stability of iron compounds in relation to redox potential (pe) and soil-pH.
The foregoing shows that aeration of fresh, non-calcareous pyritic sediments, e.g. by forced drainage, results in large quantities of H+-ions being released to the soil solution. These H+-ions lower the pH of the soil and exchange with bases at the cation exchange complex. Once the soil-pH has fallen to a level between pH 3 and pH 4, the clay minerals themselves are attacked. Mg, Fe and, in particular, Al are released from the clay lattices, and noxious Al3+ions become dominant in the soil solution and at the exchange complex.
Figure 3 shows the physiographic structure of the Mekong delta in Vietnam. The figure illustrates the relation between physiography and the occurrence of Fluvisols, and of Thionic Fluvisols in particular.
In the Mekong Delta, pyritic sediments are common in the depressions (map unit 4) where extensive areas of Thionic Fluvisols are found. There is no pyrite in the soils of map units 3 (fresh water deposits; no sulphate present during sedimentation), and 5 (sedimentation rate too high). The lower tiers of topogenous deltaic peat (map unit 6) may well contain pyrite.
Note that many soils in river deposits and marine sediments in the Mekong Delta have lost their fluvic soil properties, i.e. they are no longer stratified or have developed a cambic subsurface horizon. Most of these soils key out as Gleysols.
Figure 2
pe:pH diagram of pyrite, limonitic Fe2O3, jarosite and dissolved K+, SO42-, Fe2+ and Fe3+ at log[SO42-] = -2.3, log [K+] = -3.3, log [Fe2++Fe3+] = -5 and 1 atm. total pressure. Source: Miedema, Jongmans & Slager, 1973
Legend: (1) Granite hills (2) Pleistocene alluvial terraces (3) Holocene complex of river levees and basins (4) Holocene brackish water sediments in wide depressions (5) Holocene marine sand ridges and clay plains (6) Holocene peat dome.
Fluvisols are very young soils with weak horizon differentiation; they have mostly AC-profiles and are predominantly brown (aerated soils) and/or grey (waterlogged soils) in colour.
Their texture can vary from coarse sand in levee soils to heavy clays in basin areas. Most Fluvisols show mottling indicative of alternating reducing and oxidizing conditions. However, even if `gleyic colour patterns' occur in the upper 50 cm of the profile, the soils are not classified as Gleysols because their fluvic properties have priority in the `Key to Reference Soil Groups'.
It is evident that the characteristics of Fluvisols are dominated by their recent sedimentation and wetness: stratification, beginning ripening, chemical properties influenced by alternate reducing and oxidizing conditions, and in some environments also soil salinity. Rather special are the characteristics of Thionic Fluvisols; they will be discussed in some detail.
Most Fluvisols are wet in all or part of the profile due to stagnating groundwater and/or flood water from rivers or tides. Terraces are much better drained than the active flood plain; terrace soils are normally well-homogenized and lack fluvic properties.
The `ripening stage' of sedimentary material is judged by sqeezing a lump of soil material through one's fingers and interpreting the resistance felt. Wet clays and silt soils that have lost little water since deposition are soft and `unripe'. Such soils pose problems for agricultural use; they have a low `bearing capacity' and machines cannot be used on them.
Many coastal landforms were at one time colonized by pallustric vegetation (e.g. mangroves or reeds) that left large tubular pores in the sediment. Fluvisols on river levees and coastal sand ridges are porous and better drained than soils in low landscape positions.
Most Fluvisols have neutral or near-neutral pH values, which do not impair the availability of nutrients. Most coastal sediments contain some calcium carbonate (seashells !), and the exchange complex is saturated with bases from the sea water. High sodium saturation is not uncommon and high levels of electrolytes in the soil moisture can be a problem as well.
The hydrological and physical properties of Thionic Fluvisols are similar to those of other Fluvisols but their chemical characteristics are decidedly different. A distinction must be made between `Potential Acid Sulfate Soils', which are not yet oxidized but contain pyrite in the soil material, and `Actual Acid Sulfate Soils', which are oxidized and acidified.
Figure 3
Physiographic
sketch map of the Mekong Delta, Vietnam
Unfavourable properties of Potential Acid Sulfate Soils are:
5. Flooding of the land: flooding at spring tide may cause damage to crops.
6. Engineering problems arise when dikes, etc are constructed on soft mud. Acidity from oxidizing dike material attacks steel and concrete structures.
Unfavourable properties of Actual Acid Sulfate Soils are:
2 Fe2+ + ½ O2 + 5 H2O = 2 Fe(OH)3 + 4 H+
The good natural fertility of most Fluvisols (young soils!) and attractive dwelling sites on river levees and on higher parts in marine landscapes were recognized already by pre-historic man. Later, great civilizations developed in river landscapes and on marine plains.
Paddy rice cultivation is widespread on tropical Fluvisols with satisfactory irrigation and drainage. Paddy land should be dry for at least a few weeks every year, to prevent the soil's redox potential from becoming so low that nutritional problems (iron, H2S) arise. A dry period also stimulates microbial activity and promotes mineralization of organic matter. Many dryland crops are grown on Fluvisols, normally with some form of artificial water control.
Tidal lands that are strongly saline are best kept under mangroves or other salt tolerant vegetation. Such areas are ecologically valuable and can be used (with caution!) for fishing, hunting, salt pans, or woodcutting for charcoal or firewood.
There are two (opposite) strategies thinkable for `reclaiming' and using Potential Acid Sulfate Soils:
When discussing management and use of Actual Acid Sulfate Soils, a distinction must be made between areas with shallow inundation (less than 60 cm) and areas with deep inundation. Flap gates can be used for water control in areas with shallow inundation and at least some tidal influence in creeks or canals. In the rainy season, water can be discharged at low tide, or irrigation water can be applied as needed. Where tide water is fresh water, spring tides can be used for irrigation in the dry season. Where tides are not high enough to flood the land for rice cultivation, dry crops can be grown using flap gates to maintain a shallow groundwater table.
In many coastal lowlands, the `intensive shallow drainage system' is practiced: shallow ditches are dug at narrow spacings. This system relies on sufficient leaching of the surface soil at the start of the rainy season. Dryland crops can only be grown on raised beds whereby care must be taken not to turn the profile upside down and bring the most acid part (the subsurface soil) to the top. Rice is grown in the shallow depressions between the raised beds.
Table 1 presents theoretical lime requirements for complete neutralization of soils with various contents of oxidizable sulphur. The neutralizing capacity of a 10 cm layer (without lime; neutralization by exchange ions only) is also given. The table demonstrates the practical impossibility of the liming option: very few farmers can afford to apply to an average Acid Sulfate Soil (say 1.5 percent sulfur and an apparent density of 1.0 Mg/m3) a total of 28 (i.e. 47 minus 19) tons of lime per hectare. And that covers only the needs of the top 10-cm layer, assuming that no new acid forms during a subsequent dry spell.
Many lands with Actual Acid Sulfate Soils are not used for agriculture at all. Such hostile `wet desert' lands have an adapted vegetation, a limited fauna, and acid surface water, especially early in the wet season when acid substances dissolve. The situation seems to be a little less bleak in the equatorial climatic zone than in monsoon climates. There, Acid Sulfate Soils will not dry out as easily as in the monsoon area, and more water is available for management measures throughout the year.
TABLE 1
Lime requirements for complete neutralization of a 10-cm
soil layer. After Dent & Raiswell, 1982
|
Lime requirement of a 10 cm layer, |
Neutralizing capacity of the 10 cm layer, (no lime present) in clayey soil | |
|
Apparent soil |
percent oxidizable sulfur 0.5 1 1.5 2 3 4 |
|
|
0.6 0.8 1.0 1.2 |
9 19 28 37 56 74 12 25 37 50 74 112 16 31 47 62 19 37 56 74 |
11 14 19 22 |
The Reference Soil Group of the Gleysols holds wetland soils that, unless drained, are saturated with groundwater for long enough periods to develop a characteristic "gleyic colour pattern". This pattern is essentially made up of reddish, brownish or yellowish colours at ped surfaces and/or in the upper soil layer(s), in combination with greyish/bluish colours inside the peds and/or deeper in the soil. Common international names are `Gleyzems' and `meadow soils' (Russia), `Aqu-` suborders of Entisols, Inceptisols and Mollisols (USA), `Gley' (Germany), and `groundwater soils' and `hydro-morphic soils'.
Definition of Gleysols#
Soils,
Common soil units:
Thionic*, Histic*, Gelic*, Anthraquic*, Vertic*, Endosalic*, Andic*, Vitric*, Plinthic*, Mollic*, Gypsic*, Calcic*, Umbric*, Arenic*, Tephric*, Stagnic Tephric*, Abruptic*, Humic*, Calcaric*, Takyric*, Alcalic*, Toxic*, Sodic*, Alumic*, Dystric*, Eutric*, Haplic*.
# See Annex 1 for key to all Reference Soil Groups
@ Diagnostic horizon, property or material; see Annex 2 for full definition.
* Qualifiers for naming soil units; see Annex 3 for full definition.
Connotation: soils with clear signs of excess wetness; from R. gley, mucky mass.
Parent material: a wide range of unconsolidated materials, mainly fluvial, marine and lacustrine sediments of Pleistocene or Holocene age, with basic to acidic mineralogy.
Environment: depression areas and low landscape positions with shallow groundwater.
Profile development: mostly A(Bg)Cr or H(Bg)Cr profiles. Evidence of reduction processes with or without segregation of iron compounds within 50 cm of the surface.
Use: wetness is the main limitation of virgin Gleysols; these are covered with natural swamp vegetation and lie idle or are used for extensive grazing. Artificially drained Gleysols are used for arable cropping, dairy farming and horticulture. Gleysols in the tropics and subtropics are widely planted to rice.
Gleysols occupy an estimated 720 million hectares world-wide. They are azonal soils and occur in nearly all climates, from perhumid to arid. The largest extent of Gleysols is in sub-arctic areas in northern Russia, Siberia, Canada and Alaska, and in humid temperate and subtropical lowlands, e.g. in China and Bangladesh. An estimated 200 million hectares of Gleysols are found in the tropics, mainly in the Amazon region, equatorial Africa and the coastal swamps of Southeast Asia. See Figure 1.
Figure 1
Gleysols world-wide
Gleysols of the sub-arctic and temperate latitudes are associated with Histosols and with Fluvisols (in riverine and coastal areas). Gleysols at higher landscape positions are confined to depression areas with shallow groundwater where they occur adjacent to Luvisols and Cambisols. Gleysols in the steppe zone are found together with Chernozems and Phaeozems. Gleysols in arid regions occur predominantly in fluvial and marine lowlands, e.g. together with Solonchaks and Solonetz. A wide variety of soils (inter alia Calcisols, Gypsisols, Cambisols, Regosols, Arenosols and Leptosols) can be expected on adjacent uplands.
Gleysols in the humid tropics are confined to structural wetlands; Acrisols, Lixisols, Nitisols, Alisols and Ferralsols occur in (better-drained) adjacent uplands.
The formation of Gleysols is conditioned by excessive wetness at shallow depth (less than 50 cm from the soil surface) in some period of the year or throughout the year. Low-redox conditions, brought about by prolonged saturation of soil material in the presence of organic matter, cause reduction of ferric iron compounds to (mobile!) ferrous compounds. This explains why the permanently saturated subsoil layers of Gleysols have neutral whitish/greyish or bluish to greenish matrix colours: with the iron compounds mobilized and removed, the soil material shows its own colour, normally with a Munsell hue notation that is less red than 2.5Y.
Note that reduction of oxides to soluble Fe2+ and/or Mn2+ compounds will in practice only take place if the soil is periodically saturated with water that contains the dissolved products of organic matter decomposition.
Subsequent oxidation of transported Fe2+ and/or Mn2+ compounds (back) to oxides can take place near fissures or cracks in the soil, along living roots that have `aerenchym' (air ducts, e.g. in the roots of paddy) and along former root channels where there is supply of oxygen. Hysteresis between oxidation and (slower) reduction processes results in net accumulation of oxides near aerated spots. The soil develops a characteristic `gleyic colour pattern', with `redoximorphic features'. These comprise `reductomorphic' and `oximorphic' properties.
Note that `gleyic properties' are strictly associated with movement of the groundwater table; mottled, oxidized horizons occur on top of a fully reduced subsoil. A different type of mottling is found where perched water occurs on top of a slowly permeable subsurface horizon while the real groundwater table occurs at greater depth. In this case, the reduced horizon overlies an oxidized subsurface horizon. (Figure 2). This configuration occurs in soils that have `stagnic properties' and show a `stagnic colour pattern'.
The stagnic colour pattern differs from the gleyic colour pattern in that surfaces of peds (or parts of the soil matrix) are lighter and paler, and the interior of peds (or parts of the soil matrix) are more reddish and brighter than the (mixed average of the) non-redoximorphic soil material of the layer. The stagnic colour pattern may occur directly below the surface horizon or plough layer, or below an albic horizon.
Figure 2
Schematic horizon configurations of soils with gleyic properties and soils
with stagnic properties
The soils of rice paddies may be true Gleysols but are more often Anthraquic soil units in some other Reference Soil Group. By and large, there are two broad categories of `paddy soil':
Repetitive puddling of the surface soil of paddy soils has produced a thin, compacted and slowly permeable plough sole at the depth of cultivation. Most Anthraquic Gleysols have a reddish brown to black accumulation horizon of iron and/or manganese oxides at some depth below the plough sole. This accumulation layer formed as a result of intensive reduction processes in the puddled surface layer and translocation and precipitation of iron and manganese compounds in the (oxidized) subsoil. The accumulation layer may indurate to an impenetrable hardpan.
Soil morphological features are valuable indicators of a soil's water regime but correct interpretation of such features is not always easy.
Interpretation of oximorphic features is further complicated by the possibility that the features may be `fossil'. Soils, which were water-saturated and reduced in early Holocene or in Pleistocene times, may still show the signs today, even though they now have an entirely different water regime. The same may happen where soils are artificially drained.
Gleysol profiles have normally a spongy or matted litter layer resting on a dark grey Ah-horizon that changes sharply into a mottled grey or olive Bg-horizon. With depth, the Bg-horizon grades into a grey, olive or blue anaerobic Cr-horizon. Gleysols in fluvi-glacial sands in Central Europe may contain accumulations of hardened `bog iron' (smelted by Iron Age man).
Gleysols of the savannas have normally very dark grey to black, heavy clay surface layers that continue down to 2 metres or more in some profiles but give place to pale grey subsurface soil with prominent mottling in others (Fitzpatrick, 1986). The soil structure is medium blocky or even crumb near the soil surface but becomes coarsely prismatic at some depth. The soil material is hard when dry and sticky when wet.
Where Gleysols remain waterlogged throughout the year, except perhaps for short periods, the topsoil is typically a mixed organic and mineral (muck) H-horizon. It tops a mottled clay or sandy clay subsurface horizon over permanently anaerobic subsoil.
Reddish brown (ferrihydrite) or bright yellowish brown (goethite) mottles are indicators of oximorphic properties. Reductomorphic properties are not evidenced by specific minerals but rather by absence of iron (hydr)oxides and by `neutral whitish/greyish or bluish to greenish' matrix colours. Gleysols of coastal areas with salt meadows may feature `blue-green rust', i.e. the interior of structural peds is coloured greenish by Fe2+/Fe3+ -oxides.
The dominant attribute of Gleysols is prolonged saturation with water, associated with lack of aeration, poor rooting conditions for most crops and poor conditions for soil fauna.
Repeated wetting and drying may cause soil densification due to weakening of interparticle bonds during saturation and contraction of soil particles upon desaturation. The soil structure is likely to deteriorate if too wet a soil is tilled. Trafficability is a problem with many Gleysols.
Gleysols in depressions or at the lower ends of slopes are generally considered to be `comparatively fertile' because of their fine soil texture, slow rate of organic matter decomposition and influx of ions from adjacent uplands. They have more organic matter, greater cation exchange capacity, higher base saturation, and usually also higher levels of phosphorus and potassium, than adjacent upland soils.
The main obstacle to utilisation of Gleysols is the necessity to install a drainage system, designed to either lower the groundwater table, or intercept seepage or surface runoff water. Adequately drained Gleysols can be used for arable cropping, dairy farming or horticulture.
Soil structure will be destroyed for a long time if (too) wet soils are tilled. Gleysols in (depression) areas with unsatisfactory drainage possibilities are therefore best kept under permanent grass cover or (swamp) forest. Liming of drained Gleysols that are high in organic matter and/or of low pH value creates a better habitat for micro- and meso-organisms and enhances the rate of soil organic matter decomposition (and the supply of plant nutrients).
Gleysols can be put under tree crops only after the water table has been lowered with deep drainage ditches. Alternatively, the trees are planted on ridges that alternate with shallow depressions. In the `sorjan' system that is widely applied in tidal swamp areas with pyritic sediments in Southeast Asia, rice is grown in the inundated depressions between ridges. The difficulties discussed for Thionic Fluvisols apply also to Thionic Gleysols (see under Fluvisols).
Major landforms in mountains and formerly glaciated regions
Unstable rocky slopes and outcrops of bedrock are common features in mountainous and (formerly) glaciated regions, notably in
The highest mountains in the world are all relatively young; they started to form in the Oligocene during the period of `Alpine orogenesis' and belong to mountain belts situated at plate boundaries. Repeated levelling measurements indicate that these mountains are still actively forming today. Their properties depend in part on the nature of the plate boundaries. There are three basic plate-boundary types:
Rock outcrops and shallow, stony weathering mantles are prominent in all three mountain types, especially in the zone between the actual snow line and the lowermost extension of the ice during the last glaciation. (The present glaciers and snowfields are not considered here.) Typical erosion phenomena of this zone are the `cirques' (amphitheatre-like former snow-accumulation basins) and `U-shaped trough valleys', former fluvial valleys modified by glacial abrasion. Both have steep upper slopes, usually consisting of rough rock outcrops, and extensive lower slopes of rock debris. The sizes of rock fragments vary with the lithology: highly fissile rocks such as shale, schist and dolomite disintegrate to fine debris whereas granites give huge angular boulders.
The High Mountain Zone contains soil parent materials of glacial, fluvio-glacial and lacustrine origin alongside rock outcrops and in-situ rock weathering products:
Soil formation in these materials is slow (high altitude, low temperature) and has gone on for only 10,000 years or less. Steep slopes are subject to strong erosion, especially if stabilising vegetation is absent. Soils are regularly truncated and in a constant process of rejuvenation. As a result, (poorly developed) Leptosols and Regosols are strongly represented in high mountain areas.
The zone below the Pleistocene glaciation limit contains older weathering profiles, with more developed soils than occur in higher, more recently de-glaciated regions. They will be discussed later when Alpine fold belts lower than 3000 m. are discussed. Volcanic rocks and ash deposits can be widespread in high-mountain regions; volcanic regions and soils developed in volcanic materials are discussed elsewhere in these notes.
The `low-range mountains' of Central Europe and eastern North America are relicts of gigantic ancient continental collision structures. They formed before the Atlantic Ocean evolved and are attributed to one orogeny that produced the `Caledonian and Hercynian massifs'. They extend from northern Norway to Mauritania in N. Africa and from eastern Greenland to the southern USA. Their folded sedimentary and metamorphic rocks indicate a complex orogenic history. Both massifs have an inner zone of crystalline rocks, and an outer zone of low-grade metamorphic and sedimentary rocks. Typical inner-zone, high-grade metamorphic provinces are the Piedmont province of the USA, Brittany, the Central Massif, and the Vosges in France, the Black Forest of Germany and the Bohemian Massifs in central Europe. The Valley-and-Ridge province in the USA, the Ardennes (Belgium), the Rheinische Schiefergebirge (Germany) and Devonshire (U.K.) are largely made up of low-grade metamorphic and sedimentary rocks.
Palaeozoic orogenesis ended approximately 250 million years ago and most Caledonian and Hercynian Massifs have since degraded to low peneplains, e.g. in Brittany (France). In many areas peneplain formation went on until the Tertiary; relic (sub)tropical soils on some (high and flat) peneplain remnants still testify to the occurrence of warmer climates in the past.
The almost flat peneplains and the Mesozoic basins between them (e.g. the Paris and London Basins) were uplifted during the Late Tertiary (Alpine orogeny) and became subsequently incised by rivers. Present day differences in elevation are caused by:
The Caledonian mountain ranges of western Norway, Scotland and Ireland were strongly modified by glacial action. Glaciers protruding from the Scandinavian ice-sheet carved out deep valleys, which were later inundated as the sea level rose. The `fjords' of Norway are exceptionally deep: 1300 m, more than 1 km below the continental shelf. The Hercynian massifs in central Europe were hardly affected by glacial action because they remained below the snowline. However they became exposed to harsh periglacial conditions: there is ample evidence of `frost shattering', `frost heaving', `periglacial slope processes' and one finds indications of repeated thawing and freezing of surface soil on top of `permafrost' subsoil. All these processes promote mixing of soil material with fresh rock fragments from deeper strata, so that many soils became very stony. Remnants of Tertiary regoliths can still be found underneath periglacially affected surface tiers.
Rock outcrops are rare in plateau areas with sedimentary and low-grade metamorphic rocks. But they are common on the steep slopes of rejuvenated valleys. Leptosols and Regosols are the dominant soil groups there.
Glaciated shields such as the (Precambrian) Baltic and Canadian Shields are very complex structures; they consist of various high-grade metamorphic rocks and were strongly levelled/abraded in the past. Considering their old age, their present landforms are actually very recent; they were strongly influenced by Pleistocene glaciation. Till plains, end- and side-moraines and fluvio-glacial deposits are very common, as in glaciated mountain areas. Typical landforms such as `drumlins', `kames' and `eskers' were formed during phases of retreat, standstill and re-advance of the Pleistocene ice sheets. Glaciers or ice sheets became overridden and pushed upward depositing their `till' in elongate, elliptic hills.
Recall that `kames' and `eskers' are fluvio-glacial deposits. Streams carrying meltwater and depositing sediment on top of valley glaciers form kames whereas similar streams running beneath stagnant ice sheets form eskers. When an ice-sheet melts, any coarse sand and gravel will stay behind to form unstable terraces (kames) or well-founded long ridges (eskers) in the landscape.
In places, moving ice scraped off any saprolite that was already present in the landscape. Fresh rock outcrops and `roches moutonnées' are common, and relief became subdued. Many streams and (over-deepened) river valleys have orientations that indicate the flow direction of a former ice-sheet or glacier. The last glaciation ended only 10,000 years ago and soils are therefore recent and immature. Cryosols, Leptosols and Regosols are common soils in previously glaciated areas.
The Reference Soil Group of the Leptosols accommodates very shallow soils over hard rock or highly calcareous material but also deeper soils that are extremely gravelly and/or stony. Leptosols are azonal soils with an incomplete solum and/or without clearly expressed morphological features. They are particularly common in mountain regions. Leptosols correlate with the `Lithosols' taxa of many international classification systems (USA, FAO) and with `Lithic' subgroups of other soils groupings. In many systems, Leptosols on calcareous rock are denoted `Rendzinas'; those on acid rock are also called `Rankers'.
Definition of Leptosols#
Soils having
Common soil units:
Lithic*, Hyperskeletic*, Rendzic*, Gelic*, Vertic*, Gleyic*, Mollic*, Umbric*, Humic*, Aridic*, Gypsiric*, Calcaric*, Yermic*, Dystric*, Eutric*, Haplic*.
# See Annex 1 for key to all Reference Soil Groups
@ Diagnostic horizon, property or material; see Annex 2 for full definition.
* Qualifiers for naming soil units; see Annex 3 for full definition.
Connotation: shallow soils; from Gr. leptos, thin.
Parent material: various kinds of rock or unconsolidated materials with less than 10 percent fine earth.
Environment: mostly land at high or medium altitude and with strongly dissected topography. Leptosols are found in all climatic zones, particularly in strongly eroding areas.
Profile development: A(B)R or A(B)C profiles with a thin A-horizon. Many Leptosols in calcareous weathering material have a mollic A-horizon that shows signs of intensive biological activity.
Use: unattractive soils for arable cropping; limited potential for tree crop production or extensive grazing. By and large, Leptosols are best kept under forest.
Leptosols are the most extensive Reference Soil Group on earth, extending over approximately 1655 million hectares. Leptosols are found from the tropics to the cold polar regions and from sea level to the highest mountains. Leptosols are particularly widespread in mountain areas, notably in Asia and South America, in the Saharan and Arabian deserts, the Ungava peninsula of northern Canada and in the Alaskan Mountains. Elsewhere, Leptosols can be found on hard rocks or where erosion has kept pace with soil formation or removed the top of the soil profile. Lithic Leptosols (less than 10 cm deep) in mountain regions are the most extensive Leptosols. See Figure 1.
Figure 1
Leptosols
world-wide
The WRB definition of Leptosols refers specifically to shallow soils with continuous hard rock within 25 cm from the soil surface, excluding cemented layers such as a petrocalcic or petroplinthic horizon. However, the definition includes also deeper soils, provided that these have less than 10 percent fine earth over a depth of at least 75 cm. When discussing linkages between (diverse!) Leptosols and other Reference Soil Groups, one might make a distinction between Leptic soil units of other Reference Soil Groups and soil units of Leptosols.
The qualifier "Leptic" indicates continuous hard rock between 25 and 100 cm from the soil surface and is listed for many Reference Soil Groups but not for Histosols, Anthrosols, Vertisols, Fluvisols, Solonchaks, Gleysols, Podzols, Plinthosols, Ferralsols, Solonetz, Planosols, Chernozems, Kastanozems, Albeluvisols, Alisols, Nitisols and Arenosols.
Leptic units of Histosols, Vertisols and Podzols are known to exist and are likely to be included in the listing in the near future. Leptic units of Fluvisols, Solonchaks, Gleysols, Plinthosols, Ferralsols, Solonetz, Planosols, Chernozems, Kastanozems, Albeluvisols, Alisols and Nitisols are probably rare (or have not been sufficiently documented) or unlikely to occur; some of these may yet have to be included. Leptic Arenosols are excluded because continuous hard rock between 50 to 100 cm is not allowed in Arenosols; soils with a texture of sand or loamy sand above continuous hard rock between 25 and 100 cm would be Areni-Leptic Regosols.
Qualifiers distinguished for the Leptosol reference group link Leptosols to Cryosols (Gelic Leptosols), Gleysols (Gleyic Leptosols), Umbrisols (Umbric Leptosols), Phaeozems (Mollic Leptosols), Regosols (Hyperskeletic Leptosols) and Gypsisols (Gypsiric Leptosols). Because the central concept of Leptosols is one of weakly developed and/or very shallow soils, Leptosols are not supposed to have a histic, andic, spodic, argic or cambic horizon. For the same reason, soils with a petrocalcic, petrogypsic, petroduric or petroplinthic horizon within 25 cm of the surface are not classified as Leptosols but as Calcisols, Gypsisols, Durisols and Plinthosols respectively. Pedogenetic differentiation may still develop in very shallow soils. Thin histic surface layers occur over hard rock, especially in the boreal regions of the world and in areas with high precipitation (e.g. the "blanket peats" of the British Isles). If less than 10 cm thick, the shallow peat layer is ignored and the soils are classified as Lithic Leptosols; if more than 10 cm of peat are present, they key out as (Leptic) Histosols.
The first evidence of pedological development normally hints at formation of a cambic subsurface horizon. However, the soil ceases to be a Leptosol and becomes (classified as) a Cambisol as soon as the developing horizon meets all criteria for a real cambic horizon (at least 15 cm thick and its base deeper than 25 cm below the surface).
An andic horizon is not permitted in Leptosols but soils consisting of tephric soil material with less than 10 percent of fine earth represent intergrades of Leptosols to the Andosols.
The definition of the spodic horizon leaves the possibility open that a full-fledged Podzol exists within the depth limits of Leptosols. Such (mature!) soils are not Leptosols, despite the presence of hard rock within 25 cm from the surface, but are classified as (Leptic) Podzols.
The definition of the argic horizon precludes that an argic horizon can be present within the depth limits of Leptosols, unless there is an abrupt textural transition from the overlying horizon to the (more than 7.5 cm thick) argic horizon. In that case the soil is classified as a (Leptic) Luvisol (or Acrisol or Lixisol, depending on CECclay and base saturation).
"Gleyic properties" in Leptosols are exclusive to (deeper) soils that qualify as Leptosols because they have less than 10 percent (by weight) fine earth to a depth of at least 75 cm. "Stagnic properties" may have to be considered both in Leptosols with continuous hard rock within 25 cm of the surface and in Leptosols with less than 10 percent of fine earth to a depth of 75 cm, but Leptosols with stagnic properties have not been sufficiently documented so far.
Leptosols are genetically young soils and evidence of soil formation is normally limited to a thin A-horizon over an incipient B-horizon or directly over the unaltered parent material. The principal soil forming process in Rendzic and Mollic Leptosols is the dissolution and subsequent removal of carbonates. A relatively small residue remains behind and is thoroughly mixed with stable humifying organic matter and, in Rendzic Leptosols, fragments of limestone rock. Swelling and shrinking smectitic clays in the mineral residue are accountable for the dominance of blocky structures. Umbric Leptosols occur mostly on siliceous parent rock in (mountain) regions with a cool climate and a high precipitation sum.
Most Leptosols have an A(B)R or A(B)C configuration of only weakly expressed horizons. Rendzic and Mollic Leptosols have more pronounced morphological features. Their dark brown or black calcareous organo-mineral surface soil, in Rendzic Leptosols speckled with white limestone fragments, has a stable crumb or granular structure, or a vermicular structure with abundant earth worm casts. At the base of the soil profile, there is an abrupt change to the underlying rock or there may be a narrow transition horizon.
The Reference Soil Group of the Leptosols includes a wide variety of soils with greatly differing chemical and physical properties. By and large, Leptosols are free-draining soils with the exception of certain Hyperskeletic Leptosols that may have groundwater at shallow depth. Stagnic properties can occur in Leptosols on gentle slopes or in pockets but are rather exceptional.
The physical, chemical and biological properties of non-calcareous Leptosols are largely conditioned by the characteristics of the parent material and the climate. Calcareous Leptosols have generally better physical and chemical properties than non-calcareous ones and are also less diverse. Leptosols are normally free from noxious levels of soluble salts. However, their shallowness and/or stoniness, and implicit low water holding capacity, are serious limitations. The natural vegetation on Leptosols varies with the climate but is generally richer on calcareous Leptosols than on acid ones. Earthworms, enchytraeid worms and arthropods are the chief soil organisms. The soil fauna may be temporarily inactive in dry spells.
Leptosols have a resource potential for wet-season grazing and as forest land. Areas with Rendzic Leptosols in southeast Asia are planted to teak and mahogany; Rendzic Leptosols in the temperate zone are under (mainly) deciduous mixed forest whereas acid Lithic, Umbric and Dystric Leptosols are commonly under pine forest.
Erosion is the greatest threat to Leptosol areas, particularly in mountain regions in the temperate zone where high population pressure (tourism), overexploitation and increasing environmental pollution lead to increasing deterioration of forests and threaten large areas of vulnerable Leptosols.
Leptosols on hill slopes are generally more fertile than their counterparts on more level land. One or a few `good' crops could perhaps be grown on such slopes but at the price of severe erosion. Steep slopes with shallow and stony soils can be transformed into cultivable land through terracing, the removal of stones by hand and their use as terrace fronts. Agro-forestry (a combination of rotation of arable crops and forest under strict control) holds promise but is largely still in an experimental stage. The excessive internal drainage of many Leptosols can cause drought even in a humid environment.
The Reference Soil Group of the Regosols is a taxonomic rest group containing all soils that could not be accommodated in any of the other Reference Soil Groups. In practice, Regosols are very weakly developed mineral soils in unconsolidated materials that have only an ochric surface horizon and that are not very shallow (Leptosols), sandy (Arenosols) or with fluvic properties (Fluvisols). Regosols are extensive in eroding lands, in particular in arid and semi-arid areas and in mountain regions. Internationally, Regosols correlate with soil taxa that are marked by incipient soil formation such as `Entisols' (USA), `skeletal soils' (Australia), `Rohböden' (Germany), and `Sols peu évolués régosoliques d'érosion' or even `Sols minéraux bruts d'apport éolien ou volcanique' (France).
Definition of Regosols#
Being a taxonomic rest group, Regosols are not defined in terms of their soil properties but are rather described in terms of properties that they do not have. For all practical purposes, Regosols are soils in unconsolidated mineral material of some depth, excluding coarse textured materials and materials with fluvic properties, and have no diagnostic horizons other than an ochric horizon.
Common soil units:
Gelic*, Leptic*, Hyposalic*, Gleyic*, Thaptovitric*, Thaptoandic*, Arenic*, Aric*, Garbic*, Reductic*, Spolic*, Urbic*, Tephric*, Gelistagnic*, Stagnic*, Humic*, Gypsiric*, Calcaric*, Takyric*, Yermic*, Aridic*, Hyperochric*, Anthropic*, Skeletic*, Hyposodic*, Vermic*, Dystric*, Eutric*, Haplic*.
# See Annex 1 for key to all Reference Soil Groups
@ Diagnostic horizon, property or material; see Annex 2 for full definition.
* Qualifiers for naming soil units; see Annex 3 for full definition.
Connotation: soils in the weathered shell of the earth; from Gr. rhegos, blanket.
Parent material: unconsolidated, finely grained weathering material.
Environment: all climate zones without permafrost and at all elevations. Regosols are particularly common in arid areas, in the dry tropics and in mountain regions.
Profile development: AC-profiles with no other diagnostic horizon than an ochric surface horizon. Profile development is minimal as a consequence of young age and/or slow soil formation e.g. because of prolonged drought.
Use: land use and management vary widely. Some Regosols are used for capital-intensive irrigated farming but the most common land use is low volume grazing. Regosols in mountain areas are best left under forest.
Regosols cover an estimated 260 million hectares worldwide, mainly in arid areas in the mid-western USA, Northern Africa, the Near East and Australia. Some 50 million hectares of Regosols occur in the dry tropics and another 36 million hectares in mountain areas. The extent of most Regosol areas is only limited; consequently Regosols are common inclusions in other map units on small-scale maps. See Figure 1.
Figure 1
Regosols
world-wide
Being a taxonomic rest group, Regosols are found in association with a wide range of other Reference Soil Groups. However they are particularly common alongside other young or poorly developed soils in arid, degrading or eroding areas. As the listing of `common soil units' (see above) suggests, many Regosols are intergrades, with properties tending towards those of Cryosols (cold regions), Andosols, Leptosols or Umbrisols (mountain regions), Calcisols or Gypsisols (arid regions), Arenosols and Podzols (sandy deposits) or Cambisols.
Soil forming processes have had a minimal effect on the properties of Regosols. This may have been caused by
Profile development is limited to formation of a thin ochric surface horizon over (almost) unaltered parent material. The paucity of pedogenetic transformation products explains the low coherence of the matrix material and makes that soil colours are normally (still) determined by the composition of the mineral soil fraction. In regions with a considerable evaporation surplus over precipitation, some lime and/or gypsum may have accumulated at shallow depth in the profile but not to the extent that a calcic or gypsic horizon is present. Soils in recent deposits of mine waste, urban waste, landfills and dredgings that are (still) too young for soil formation to occur, are included in the Reference Soil Group of the Regosols.
The great variation among Regosols (taxonomic rest group!) makes it virtually impossible to give a generalised account of Regosol characteristics.
The central concept of a Regosol is a deep, well-drained, medium-textured, non-differentiated mineral soil that has minimal expression of diagnostic horizons (other than an ochric surface horizon), properties or materials.
Some general observations:
Regosols in desert areas have minimal agricultural significance. Regosols in the steppe region with 500-1000 mm rainfall per annum need irrigation for crop production. The low moisture holding capacity of these soils calls for frequent applications of irrigation water; sprinkler or trickle irrigation solves the problem but is rarely economic. Where rainfall exceeds 750 mm per annum, the entire profile is raised to its (low) water holding capacity early in the wet season; improvement of dry farming practices may then be a better investment than installation of costly irrigation facilities.
Many Regosols are used for extensive grazing. Many Regosols on colluvial deposits in the loess belt of northern Europe and North America are cultivated; they are planted to small grains, sugar beet or fruit trees. Regosols in mountain regions are quite delicate and are best left under forest.
