Introduction
Sand accumulated by swash drifts with the wind on open beaches and piles up around vegetation patches and other obstacles. At the same time the plants are able to take nutrients from the beach. The flora of the dune fields is adapted to the poor surroundings, but in time the rain water washes the soil and the plant debris increases the amount of organic material, leading to an ecological succession. When sand brought by the waves welds onto the prograding beach, the plant communities form successive zones parallel to the shoreline. The geomorphological development of the shore depends on this succession. In a dune landscape the geomorphological processes are gradually impeded by plant biomass production (Van Der Meulen 1990). Both geomorphological and pedological processes are active and easy to detect on coastal dune fields.
The active coastal dune fields of Finland are located between the swash zone ecotone and the forest edge ecotone. The effect of waves during storms reaches far beyond the actual beach and can cause great changes to sandy beaches at an exceptional speed. On the other hand, there can be periods, perhaps decades, of quiet evolution between the more severe storms.
The coasts of Finland are characterized by a rapidly advancing shoreline, brackish water, a microtidal environment, a humid climate with moderate winds and long, severe winters. The aeolian sand, which originates from glaciofluvial sediments washed by the waves, is usually very infertile and acid because of the acidic nature of the bedrock, and this poor environment can support only a small number of species. There is only about 1300 ha of the open dune landscapes on the coasts of Finland remaining, and this area is continuously diminishing because of building and rapid invasion by pine forest. On these grounds one can very well speak of a threatened environment.
This work treats the coastal dune fields in Finland in their entirety, from the shoreline morphodynamics to the effectively stabilizing forest cover. The aim is to explain the evolution of these coasts during the present century and the most important factors affecting it. Topic of investigation is thus the geomorphological and the ecological succession of these areas and variations in this succession.
Coastal dune chronology in relation to climate and sea level
The wind has been shaping the sand dunes on the Finnish coasts since the beginning of the Holocene Epoch. Most of these dunes were formed in a periglacial environment (Aartolahti 1980: 81) and are now stabilized far away from the present coastline. They often have parabolic outlines and they are higher than the dunes on the present coasts. This older phase of dune formation ended about 8000 BP, when the climate was becoming warmer (Eronen & Olander 1990: 60) and more humid (Korhola 1992: 84, 1995), favouring plant growth. According to Stetler and Gaylord (1996) the interaction between climate and aeolian activity is sensitive even to small changes in precipitation. During the Atlantic stage (8000 - 5000 BP) the fine sand was mainly shaped into low beach ridges (Aartolahti 1973: 47; Hellemaa 1980: 92). Apart from wind, dry, cold climatic conditions also favour aeolian activity (Shuisky 1986: 37), although dunes can be formed in humid climates, too.
Judging from the height of the bases of the coastal dunes and the rate of land uplift, the younger phase in coastal dune formation in Finland began only about 1000 BP (Aartolahti 1990: 214), and most of these dunes were formed during the last 500 years (Aartolahti 1980: 82). This increase in aeolian activity was at least partly due to the intensification of human activity, as elsewhere in the world (Orme 1990: 334). The felling of dune forests and the grazing of domestic livestock on sandy sea-shore meadows have led to the formation of active dunes. The radiocarbon dates available suggest that the dunes in the Yyteri area formed during period 1000 - 1400 AD and again from the 18th century onwards (Wallin 1980), while the oldest dunes of Hailuoto seem to have begun to form about 1500 years ago (Alestalo 1986: 151). Dune activity on the barriers of Poland, which make up part of the southern coast of the Baltic, has been dated both archaeologically and by radiocarbon analysis of buried soils, and it has been shown that aeolian activity there came to a halt between 1500 and 500 BP (Borówka 1990a: 29, 1990b: 307-310), but that the dunes have become active again during the last 500 years. There are historical records of dry summers in the Netherlands in the 10th century and around 1800 AD (Klijn 1990a: 93), and it is known that after long, dry periods the vegetation is vulnerable to high winds and storms, which often follow droughts (Lamb 1977). According to Aartolahti (1976: 91; 1980: 82), the old stabilized dunes in Finland were activated during the climatically distinct Little Ice Age (1550 - 1850 AD), and it was then that the transgressive dunes located landward of the present coastal dunes were formed. The sea level was lower during that period (Christiansen et al. 1990) and the winds on the Finnish coasts were more powerful than nowadays, judging from the grain size of the aeolian material (Aartolahti 1976: 91). Dune formation in the Netherlands seems to have been connected with marine erosion and storms (Klijn 1990a), while Danish research (Christiansen et al. 1990) suggests that major dune building takes place in periods of falling or low sea level.
Aeolian activity is usually accelerated by abundant transport of sand towards the coast or by destruction of the plant cover on older stable dunes. Sand accumulation is related to falling sea level, when the groundwater level also falls at the coast. The reason for the destruction of the plant cover may be: 1) pedological, when leaching leads to extremely vulnerable stages in the vegetation succession (Klijn 1990a: 93), 2) climatic, when the climate is drier or cooler or the frequency of strong winds and storms increases, 3) marine erosion, which is often connected with a sea level rise, 4) biotic, e.g. plant diseases or 5) anthropic, such as the grazing of domestic livestock or felling of areas of forest. There are usually many factors operative at the same time, e.g. dry summers usually bring with them forest fires. The use of the dune vegetation as fuel in Denmark during the Little Ice Age was mainly a consequence of climatic deterioration (Christiansen et al. 1990: 63).
The transgressive dunes on the Finnish coast were still actively moving in the 1930's (Mattila 1938; Alestalo 1971), when average wind speeds seem to have been about the same as they are nowadays (Lemberg 1933: 20). It is difficult to compare wind records because of the progress made in measuring techniques, but according to Heino (1994: 90-91) the winds were more powerful at the beginning of this century, when the climate was warming up to the 1930's, than during the following period of cooling (1945-70). The present climatic warming (Heino 1994: 93) seems also to be accompanied by stormy winds (Kuusisto 1993), but nowadays the vegetation cover has spread over the old transgressive dunes and mostly stabilized them, so that their movement is insignificant. Land uplift with invasion of vegetation on the proximal deflation plains have separated these dunes from the present shoreline, so that they cannot get replenishment from the shore. In places the coastal sand accretion is abundant and new coastal dunes are still forming.
The coastal dune fields in Finland have been quickly covered by plants in recent times, especially around the shallow accumulation bays in the Gulf of Bothnia. The reason for this is mainly that the seashore meadows are no longer used for grazing (Heikkinen & Tikkanen 1987: 265). Afforestation of formerly grazed dune meadows is common in Europe (Hesp & Thom 1990: 280; Van Dijk 1992). According to Alestalo (1986: 153), when the grazing of sheep on shore meadows of Hailuoto stopped in 1954, the blown sand soon began to accumulate again to form consecutive dune ridges bound by lyme-grass (Leymus arenarius). The reason for the spread of such a vegetation cover may partly lie in acceleration of the greenhouse effect (e.g. Hulme 1994; Heino 1996). According to Heino (1994), this climate change cannot yet be seen in Finnish temperature records, but precipitation has increased over the last decade (Heino1994: 133) and the snow-free period has grown longer (Heino 1994: 177-178), as the spring has become warmer (Heino 1994: 94, 100). On the other hand, the average length of the growing season did not change between the two normal periods 1931-60 and 1961-90 (Climatological Statistics in Finland 1961-1990; Kolkki 1981). As far as the growth of mosses and lichens living on sand is concerned, it is enough that the temperature rises near zero (Kallio 1980: 335).
It has been predicted that acceleration of the greenhouse effect will lead to a more pronounced eustatic sea-level rise, caused mainly by thermal expansion of the ocean water and the melting of glaciers (e.g. Coker et al. 1989: 149; Carter 1991). There are many difficulties involved in studying sea level, however, because of tectonics, winds, currents, tides etc. that cause local variation. According to Woodroffe (1994: 440) there is no valid proof of an acceleration in sea-level rise this century. It has been estimated that the present rate of eustatic rise is 1.0-1.5 mm in a year (Coker et al. 1989), and although the estimates of an accelerating rate published by the Intergovernmental Panel on Climate Change will be between 1.9 and 6.4 mm a year for the period 1985-2030 (Carter 1991: 31), the latest figures are still only about 1.5 mm per year (Gorniz 1995). Åse (1994) claims that coastal formations caused by eustatic transgressions near Stockholm correlate well with warm climatic periods, and it has also been proposed that the increase in the use of groundwater, irrigation, forest felling and desertification is leading to a sea-level rise, too (Woodroffe 1994: 441).
Since the melting of the Scandinavian ice sheet the rate of land uplift has slowed down, so that the rate of shoreline displacement is now slower than earlier. Granö and Roto (1989), in their discussion of the duration of shore exposure along the Finnish coasts, estimate the eustatic sea-level rise to be only 0.8 mm per year, but if this rise should accelerate, it will also slow down shoreline displacement.
Sea level has a direct effect on the groundwater level and on the moisture content and vegetation of dune slacks, while pollution and increased amounts of nutrients in the seawater cause sandy beaches to be more eutrophic, and air pollution, both acidic and eutrophic, also makes a difference. A dense grass vegetation, dominated by Calamagrostis epigejos and others, has covered some dunes in the Netherlands as a result of recent acidic rains (Salman 1991: 30). By contrast, occasional high winds and the constantly increasing use of coasts as recreational areas are apt to destroy the vegetation cover.
Apart from the primary material the sand supply of a beach depends on the interacting coastal morphodynamics and the power of the waves. The wave type, inclination of the beach and grain size of the material are interdependent, and it is these factors that control sediment transport on the beach, the width of the surf zone and the type of nearshore bars. The "Australian School" of coastal geomorphology has classified coastal morphodynamics on these grounds (e.g. Short & Hesp 1982; Wright & Short 1984). At one extreme are coasts from which the wave energy is reflected and at the other accumulation coasts, over which the wave energy is dissipated. Reflective beaches are characterized by steep slopes, narrow breaker zones, beach cusps and the absence of nearshore bars, while dissipative beaches with a gently sloping profile and shoals are characterized by wide surf zones and multiple parallel bars. Intermediate beaches incorporate elements of both types (Carter 1988: 100-105).
Wave refraction and diffraction are features that bend the waves and redistribute their energy in the surf zone. Loose material drifts with the currents towards the areas of lower energy and accumulation. Successive wave refraction can shape embayment beaches into effective sediment traps (Carter 1988: 211-217). Aeolian sand beaches usually border on a wide wave-moulded cut-and-built terrace with many bars, which represent the amount of loose material. According to Keränen (1985: 13), who has studied sandy shore formations and processes in Lake Oulujärvi (Finland), the primary slope becomes gentler and the primary material finer as the width of a cut-and-built terrace increases, while gently sloping shores are characterized by multiple parallel bars. In longshore bars material can be transported either along the bar or shorewards (Keränen 1985: 32, Carter 1986), and a whole bar can also move towards the shoreline. Swash bars in particular can weld into a berm (Keränen 1985: 45-47; Carter 1986, 1988: 113-114). Longshore bars are an important source of material on drift sand beaches. Wave-induced sand transport is at its maximum on low gradient dissipative beaches (Short & Hesp 1982: 266), but berm and dune material is transported to the bars during storms, and the bars can move seawards (Keränen 1985: 43). Carter (1986) claims, that beach ridges are formed when the material eroded by storms is transported back onto the beach face. Sandy shores that are being moulded constantly by wave action usually have no vegetation cover.
When the sea-level fluctuates, there are wide, dry, sandy flats at times that provide material for wind transport, while at other times the waves can erode the foredune. The width of a beach correlates with the amount of sand carried by the wind (Davidson-Arnott & Law 1990: 198). According to Carter (1990: 34), the amount of sand on a beach on the Irish coasts, the local sediment supply, is closely associated with dune topography and development (Fig. 1). In low sediment supply systems, wave undercutting and deflation will lead to scarpment formation and crestal accumulation, creating landward-dipping sand sheets, while in high sediment supply systems consecutive parallel dune ridges are formed that can be either transgressive or bound by the vegetation. There is often a balance between seasonally alternating erosion and accumulation (Carter 1986: 192; Thom & Hall 1991; Sherman & Bauer 1993).
Fig. 1. The three basic types of Irish coastal dune topography associated with low, balanced and high sediment influxes. After Carter (1990: fig. 3), but the drawn structures conform to the observations made on the Finnish coasts.
A matrix constructed by Psuty (1992: 4-6) shows the topographic/morphological outcomes of foredune forms as combinations of the beach sediment budget and the foredune sediment budget, each of which can be either positive or negative. Psuty distinguishes the following types: beach ridge topography, dune ridge topography, foredune development and foredune attenuation. The last type includes blowouts, separate hummock dunes and washover by waves. According to this matrix, the foredune budget can be positive or balanced even while the total dune/beach budget is negative and the shore line is being eroded (Psuty 1990: 175). Coastal erosion caused by sea-level rise has been studied by Healy (1991).
Sand supply, wind strength and the density of the vegetation cover are usually the primary factors affecting dune topography, these factors are said to together determine the type of dunes (Hack 1941). The interaction of these three factors is obvious on coastal dune fields and is emphasized by mutual dependence on the seashore location. Apart from the primary grain size, sand supply, shape and width of the beach, the direction, strength and duration of both prevailing winds and extreme storms are also factors affecting the geomorphology of a dune field. On an open sandy field the wind velocity and direction together with the sand supply determine the type of dune. Sand accumulates around the plants on the coasts, and thus the plant species have an effect on the form and development of the dunes. On the other hand, the topography of a dune field will affect the evolution of its plant cover.
In Pye's (1990: 355-357) model, the coastal dune morphology as well as the shoreline dynamics will be dependent on wind energy and the sand trapping efficiency of the vegetation. The extreme types of dunes are thus parallel ridges bound by vegetation on a prograding beach and unvegetated, transgressive sand sheets on an eroding beach. Classifications of similar kind (Fig. 2) are also presented by Short & Hesp (1982) and Bird (1990: 22-23). Consecutive parallel dune ridges are formed whenever there is ample accumulating sand with a thriving vegetation. It is difficult to distinguish low dune ridges from beach ridges (Keränen 1986), which are often formed by waves and wind together. Beach ridges are the essential basis of prograding coastal landforms (Carter 1986: 210).
Fig. 2. Parallel dunes, which have developed as successive foredunes, are held in place by the vegetation. When its cover is interrupted, blowouts develop, and with further depletion of the vegetation, coastal dunes may become mobile and transgressive. Modified after Bird (1990: fig. 4).
In the zone behind the foredunes the influence of coastal processes will be of importance only during severe storms, usually combined with a sea-level rise. The foredune will prevent sand transportation landwards (Gomes et al. 1992). In this zone the evolution of the topography is mainly affected by deflation, rain erosion and the ecological succession, including soil formation. Mattila (1938: 5-12) described the zonation of the Vattajanhietikko dune field in Finland as one in which open beaches and foredunes were followed by a deflation surface shielded by gravel and small stones. Sand had become trapped around the plants in this zone to form small hummock dunes. In its distal part the deflation surface was covered with aeolian sand and bordered by a transgressive dune ridge, which was partly an open ridge and partly characterized by large dune hummocks and wind channels. Alestalo (1979: 117-119) distinguished similar zones on the dune field of Pajuperä, Hailuoto, where he regarded the geolittoral zone as including the lower beach, foredune and deflation surface, while the epilittoral zone contained small hummock dunes that increased in age further away from the beach, cover sand and dead deflation surfaces bordered distally by a parabolic dune.
The Sand Dune Inventory of Europe (Doody 1991) includes a zonation on successional grounds, into: strandline, foredune, dune grassland, dune slack, dune heath, scrub and woodland, while the ecological gradients in coastal dune fields as described by Carter (1988: 320, 323-332) comprise a pioneer stage followed by more stabilized intermediate dunes and then by mature plant communities. The successional sequences are determined by sand transport rates (Moreno-Casasola 1986; McLachlan 1990: 213), so that dunes can be classified on the grounds of both age and sand mobility (Carter & Wilson 1990).
Pedogenic processes such as leaching and the accumulation of organic matter affect the succession of the plant cover and the erodiblity of the soil. The intensity of water erosion depends on the water repellency of the sand surface, which in turn is brought about by a number of features that include fungal proliferation, litter decomposition, soil pH and texture (Dekker & Jungerius 1990: 174). The most important seem to be hydrophobic organic materials. Destruction of the vegetation cover, e.g. by human action, will intensify water erosion, deflation and the formation of secondary aeolian accumulations. Little research has been done into erosional landforms on coastal dunes (e.g. Ritchie 1972), but the topic is well summarized by Carter, Hesp and Norstrom (1990). Seppälä (1984: 47-48) lists possible reasons for deflation ceasing in Finnish Lapland.
An open, windy, barren habitat will affect both the flora and the life forms of the plants, and differences in flora and geomorphology between coasts can help us to evaluate the significance of the various factors. The most distant areas examined here are situated over 600 km apart in a north-south direction, so that the vegetational zone boundaries (Ahti et al. 1968) may be reflected in these features.
Lemberg (1933-35) provided quite a detailed description of the vegetation of the Finnish coastal dunes in his three-part work. Some of the coasts studied by him now have a forest or reedbed cover, but others have remained as separate patches of sparse vegetation surrounded by forest. Only those that have formed on the slopes of large eskers are still active. A more recent study of vascular plant communities on the Finnish coastal dunes has been made by Willers (1988), and a sand dune inventory has been produced for Europe that reports briefly on this vegetation, too (Hellemaa & Doody 1991). Willers (1988: 41-88) named the plant communities of sandy shores after: Juncus bufonius, Lathyrus japonicus, Honkenya peploides, Ammophila arenaria, Calamagrostis epigejos, Honkenya peploides - Leymus arenarius (the subassociations of which are characterized by Festuca ovina and Festuca rubra), Festuca polesica, Carex arenaria, Salix repens - Empetrum nigrum. By no means all of the plant communities studied by Willers are located on dune shores proper.
In Finland, Lumme (1934) made a review of earlier research on dune fields. Rosberg (1895a), Leiviskä (1905a) and Okko (1949) wrote in earlier times about the coastal dune fields beside the Gulf of Bothnia, and Alestalo (1971) studied the movement of these coastal dunes by means of dendrocronology and also described the vegetation and geomorphology of dunes (Alestalo 1982), and the history and ecological succession of the island of Hailuoto (Alestalo 1979; 1986). Vartiainen (1980) investigated the succession of the island vegetation of the northernmost parts of the Gulf of Bothnia, but her material did not include Hailuoto, and Heikkinen and Tikkanen (1987), Jämbäck (1995) and Tapaninen (1995) have all studied the history of the Kalajoki dune field. Fontell (1926) and Mattila (1938) wrote in their time about the dunes of the foreland of Vattaja, the former mainly about the vegetation and the latter about the geomorphology, and Wallin (1980) worked on the Yyteri dune field, an area also covered by the geomorphological interpretation of Tikkanen (1981). Skytén (1978) wrote about the dune vegetation on the Hanko Peninsula and the erosion caused by tourism.
The shore can be divided into the sublittoral zone, which is continuously submerged, the littoral zone, which is under water at times, and the epilittoral zone, which is above the highest water level. The littoral zone consists in turn of the hydrolittoral, lying below the mean water level, and the geolittoral above it. The lower limit of the epilittoral is often marked by wrack, driftwood and algal debris (Vartiainen 1980: 17). Above the waterline the shore can also be divided into a wetter lower beach and a drier upper beach, which forms the lower part of the windward dune slope. Water flows into the area behind the foredune only during severe storms.
Longshore bars are formed in a position where they are submerged by the orbital current (Keränen 1985: 20-23, 1986: 80). They are regularly grouped, long ridges, running parallel to the shore, and may be exposed at times, when the water level is low.
A swash bar (swash ridge) is a depositional form which reaches above the water level. It develops when a strong progressive current is shifting a bar towards the shore and the swash starts to deposit material above surface (Keränen 1985: 45-47, 1986: 80). Material of a swash bar can originate from the beach. In the early stage of its development, a swash bar is separated from the shoreline by a pool. Swash bars are easy to detect on aerial photographs.
A berm is accumulated mainly above the water level by the swash. The inclination of a seaward-sloping beach face depends on the grain size of its material and the power of the waves. If the sand supply is abundant, e.g. carried by beach drift, and the sediments cannot be transported further because of the curvature of the shoreline, then a berm can grow into a ridge (Keränen 1985: 44-45). The highest altitudes of ridge-like berms of this kind are the highest average water levels plus the highest levels of swash during their formation (Keränen 1986: 81). These beach ridges can be formed as a result of one storm, or they may have accumulated during numerous periods of high wind combined with high water levels. Short and Hesp (1982: 268) note that sediments on reflective coasts are stored in a subaerial, steeply sloping beach face berm.
According to Carter (1988: 121) a beach ridge is a berm that has survived erosion. Tanner (1993: 220) classifies beach ridges into four main types: 1) Swash-built ridges (the berms mentioned above) tend to be low, typically less than a metre in height, and uniform, and they occur in swarms or systems and contain low-angle cross-bedding. 2) Settling-lag ridges have the shape and size of the previous class but do not contain any cross-bedding, being formed from storms or other high-energy events by settling, as can be seen from their granulometric characteristics. 3) Dune ridges tend to be higher and more irregular in shape than the ridges in the previous classes and may be located within them or adjacent to them. They are built largely by wind and contain aeolian cross-bedding and usually a small swash-built core. 4) Finally storm-surge ridges occur singly rather than in swarms, are up to 8-10 m tall, are convex in shape and have concentric internal bedding. They are built by successive storm surges over a long period of time on a stable coastline.
In spits, bars and tombolos the beach drift carries material along these formations, whereas the material in submerged longshore bars is transported mainly towards the water line (Keränen 1985). Spits and tombolos reach above the water level and are attached to the shoreline at least at one end.
A transverse bar lies submerged, perpendicular or diagonal to the shoreline. It is formed by a progressive wave or current occurring in shallow water (Keränen 1985: 34). When the upper parts of bars of this kind are connected with the shoreline, the result is a cuspate shoreline. Their development may be connected with the wave refraction caused by shoals (Carter 1988: 118-119). Large surf cusps formed like this are quite permanent, remaining from year to year, and becoming oriented towards the greatest fetch or the prevailing winds (Keränen 1985: 34).
The smaller, regularly spaced swash cusps that are formed on shores at times are the product of stationary subharmonic or synchronous edge wave development within the reflective domain (Komar 1983: 102; Carter 1988: 124-125; Rasch et al. 1993). These cusps are usually shaped further when backwash and rip currents returning from the shore erode the area between the cusp horns. Cusps of this kind are erosional, and a berm or beach ridge may be breached in this way (Pyökäri 1982). Constructional cusps may also form, if rip currents pile material on the cusp horns. Swash cusps can also arise from the occurrence of erosion and deposition at the same time (Komar 1983: 102; Carter 1988: 120-125; Miller et al. 1989). Water does not penetrate sand as easily as gravel, and therefore gravel layers settle uppermost on cusps (Uusinoka 1984: 109). The formation of large swash cusps is favoured by long swells after a storm (Carter 1988: 130).
Dunes, formed by wind, have been classified according to their topography, vegetation cover, age, mobility, material and genesis (e.g. Van Dieren 1934: 202-212; King 1972: 180; Goldsmith 1978: 178; Hempel 1980: 430-431; Pye 1983: 539; Chorley et al. 1984: 415-424; Bird 1990; Pye & Tsoar 1990: 160-220). They also can be divided into primary and secondary dunes, the latter being formed when stabilized dunes are activated.
Small, phytogenic, semi-spherical or elongated embryo dunes are sometimes to be found on beaches. They are often destroyed by storms. Small shadow dunes evolve in the lee of boulders, clumps of vegetation or other obstacles. They are also called lee dunes, drift dunes or attached dunes. Some of them are horseshoe-shaped sand accumulations, as the air flow is deflected around and over the obstacle and a horseshoe vortex is created (Pye & Tsoar 1990: 163-165).
An incipient foredune (Carter & Wilson 1990: 129) lies nearest the beach. It is formed when small embryo dunes bound by vegetation grow into a continuous dune ridge, a transverse coastal dune, parallel to the shoreline. They can evolve on a beach ridge or on a drift line embryo dune. A dune has a distinct windward slope (stoss slope or proximal slope) where the sand is tightly packed by the wind and a leeward slope (distal slope) that consists of loose strata. If a leeward slope is high enough (Bagnold 1941: 200-203) it may evolve into a slipface, the inclination of which depends on the angle of repose, which is usually 34o for aeolian sand. Because of alternating aeolian accumulation and erosion, the windward slopes contain cross-bedding. The inclination of the foreset strata in the leeward slopes depends both on aeolian accumulation and on sand-binding vegetation, the amount of water erosion and niveo-aeolian accumulation.
On a prograding beach several foredunes may be formed one after another, or partly one above another, as a compound dune. Behind the youngest, incipient foredune, mosses and lichens stabilize the sand surface and lend it a more greyish colour. Intermediate dunes of this kind may begin to wear down because of erosion. The established dunes investigated by Carter and Wilson (1990: 145) on the Irish coast did not wear down, as they were quickly and effectively stabilized by vegetation.
A hummock dune, or coppice dune (Hesp & Thom 1990: 271-272), is a separate, usually low, semi-spherical, shield-like mound of sand bound by vegetation. It is usually formed by sand accumulation around a clump of vegetation on a deflation surface landward from the foredune, but it can also be formed when a dune ridge breaks down. Hummock dunes of this kind are usually more irregular in shape than the accumulation forms, and their slopes are steeper.
Transgressive dunes are ones that have moved from their original site. Their proximal slopes are deflation surfaces and their distal slopes are characterized by slipfaces. They may be transverse dunes, parallel to the coastline, or parabolic dunes, with arcuates opening to the coast (Pye & Tsoar 1990: 200-204).
Precipitation dunes are formed when the wind power is slowed down by the edge of a forest and sand is precipitated onto a dune ridge. These usually have long, high slipfaces (Pye & Tsoar 1990: 205-207).
A dune slack, a depression between dunes, can have a wet or dry surface. Water often penetrates the sand at times of high water levels and floods into the dune slacks.
A blowout can be a shallow, roundish, saucer-like or scooped hollow or a deeper deflation bowl. Transversal dune ridges are broken by wind-scoured gaps or transport corridors, and sand can accumulate in the lee of these to form rim dunes or fan-like blowover 'deltas' (Carter et al. 1990: 239-242). In this way blowouts produce bends in foredune ridges. The spiral vortex of an air flow can also erode smaller wind channels on the proximal sides of obstacles.
There are often deflation surfaces on windward dune slopes and windward of transgressive dunes, that are being reduced by deflation. Coarse material enriches on these surfaces in the form of lag deposits. One can find table-like erosion remnants on the deflation surfaces which grow higher as the sand originating from the deflation surface, accumulates among their vegetation.
More or less even cover sand layers can usually be found leeward of high dunes.