Coastal acid sulfate soils in the Saloum River basin, Senegal.
Coastal landscapes around the world frequently contain acid sulfate soils (ASS). These are defined as soils that are severely acidic or have the potential to become acidic as a result of iron sulfide oxidation (van Breemen 1973; Fanning et al. 2002; Sullivan et al. 2009; Claff et al. 2010). Pyrite (Fe[S.sub.2]) is the most common sulfide in ASS (Sullivan and Bush 1997; Bush et al. 2000). Biologically mediated oxidation of pyrite is attributed to the activity of bacteria. Specifically, the bacterium Acidithiobacillus ferrooxidans has been identified in ASS (Durr et al. 2006). Much of the iron (Fe) released from pyrite oxidation is retained in sulfuric horizons via the precipitation of various [Fe.sup.111] minerals, as well as [Al.sup.111] minerals (van Breemen and Harmsen 1975; Johnston et al. 2011). Iron sulfates and iron oxides are ubiquitous in ASS environments (van Breemen 1973; Dent and Pons 1995; Fitzpatrick et al. 2005; Boman et al. 2008; Burton et al. 2008) and exert a major influence on biogeochemical processes (Bigham et al. 1996; Jones et al. 2006; Claff et al. 2010). Johnston et al. (2011) demonstrated that, during oxidative pedogenesis, long-term interactions between hydrology and topography influence the distribution of Fe mineral fractions in the landscapes of coastal ASS (CASS) (van Breemen and Harmsen 1975; Willett and Walker 1982; Lin and Melville 1993; Husson et al. 2000). In addition, short-term seasonal fluctuations in hydrology and redox conditions can stimulate reductive dissolution of Fe (van Breemen 1975; Johnston et al. 2005; Kraal et al. 2013), causing dynamic and potentially rapid mineralogical transformations (Burton et al. 2006, 2008). Rapid mineral transformations were also observed by Fanning et al. (2010) in the USA Mid-Atlantic region during both sulfidisation and sulfurisation processes, leaving a lasting imprint on the mineralogy and other properties of soils affected by these processes. Coastal plain soils of West Africa arc generally saline and acid, a result of seawater intrusion and pyrite oxidation, respectively (Vicillefon 1977; Marius 1985; Sadio 1991; Montoroi 1996). The according processes affect mineral weathering in different ways. Only a few studies involving ASS in the coastal area of West Africa have considered their mineralogical constitution, and most of these have been carried out in mangrove areas (intertidal zone). The most influential contribution was from Marius and Lucas (1991) for mangrove ecosystems on the shoreline along West Africa, where the predominance of quartz (Si[O.sub.2]), kaolinitc ([Al.sub.2][Si.sub.2][O.sub.5][(OH).sub.4]) and smectite in the soils, as well as the presence of pyrite, jarosite (K[Fe.sup.111.sub.3][(OH).sub.6](S[O.sub.4])) and halite (NaCl), were observed (Souza-Junior et al. 2008). Igwe et al. (2005) attributed the occurrence of advanced weathering products such as kaolinite and crystalline Fe and Al in recent parent materials of the River Niger floodplain (eastern Nigeria) to addition of intensive weathered material from the upland. Sylla et al. (1996) analysed the spatial variability of soil acidity in the mangrove ecosystem of West Africa for land-suitability planning and developed a conceptual framework for four West African estuaries: the Casamance River (Senegal), the Gambia River (Gambia), the Geba-Cacheu Rivers (Guinea Bissau) and the Great Searcies River (Sierra Leone). Extreme salinities and acidities were mentioned and had developed in many ASS along the coast of the Atlantic Ocean in Senegal, Gambia and Guinea Bissau, particularly in tidal flats away from tidal rivers, such as the Saloum, Gambia, Casamance and Geba-Cacheu (Dent and Ahmed 1995; Fanning and Burch 1997).
In Senegal, ASS issues have attracted considerable attention over the recent decades. Barbiero et al. (2005) demonstrated strong relationships between ASS and Vertisols in northern Senegal, in the Senegal River valley, and concluded that soil morphology in this area results from development and control of acidity and not from changes in past climates. Deckers et al. (1993) suggested that, in many instances, the resulting ASS are buried under more recent sediments of fluvial origin, and can therefore be considered as fossil soils. In southern Senegal, Vieillefon (1977), Marius (1985), Boivin (1991) and Montoroi (1996) observed changes in chemical properties of ASS in relation to Fe and sulfur (S) reactions within the Casamance River basin. In addition to the jarosite, Le Brusq et al. (1987) and Montoroi (1995) showed the presence of other iron sulfates, such as rozenite ([Fe.sup.II] S[O.sub.4], 4[H.sub.2]O), indicating a combined reduction and acidification process. CASS in the valley of the Casamance River arc connected to red soils (Ferralsols; IUSS Working Group WRB 2006) distributed atop the surrounding plateaus by a system of soil transformation where iron removal processes along the slopes dominate. In central Senegal (Saloum region), within Sudano-Sahelian climate zone, there is a lack of mineralogical information on CASS in contact with continental soils. Sadio (1991) investigated the agronomic and forestry potentialities of saline ASS in the Saloum River basin for soil-management purposes. He showed that acidification was not widespread in West Africa before 1971. It extended from the start of the catastrophic drought of the 1970s, which resulted in the oxidation of pyrite present in these coastal Holocene marine sediments.
The present study aims to: (i) describe the genesis and mineralogy of CASS in the Saloum River basin; (ii) display their spatial distribution and morphological features; and (iii) demonstrate the impact of environmental factors on mineral formation and crystallinity.
Materials and methods
The Saloum River basin, west-central of Senegal, West Africa, is located at 13[degrees]35'-14[degrees]20'N and 16[degrees]00'-16[degrees]50'W. It covers ~250 000 ha and is bounded on the west by the Atlantic Ocean (Fig. 1). The climate of the Saloum region is of Sudano-Sahelian type with a high mean annual temperature of 29[degrees]C, a high annual potential evapotranspiration ranging from 1500 to 2500 mm and a mean annual precipitation ranging from 600 to 800 mm [year.sup.-1], from north to south. Rainfall is limited to the period June--October (rainy season). The geology consists of Holocene marine sediments underlain by Miocene deposits known as the 'Continental Terminal', a thick series of ferruginised and argillaceous sandstones, mudstones and carbonaceous layers, deposited under fluviatile and lacustrine conditions (Wright et al. 1985), which has undergone a profound ferrallitic pedogenesis (Michel 1973; Kalck 1978; Lapparticnt 1985; Diop 1990). The hydrologic system is the Saloum River, a tide-influenced 'inverse estuary', fed only by a limited river flow for 2-3 months during the rainy season and characterised by high salinities of >80 [per thousand] resulting from seawater encroachment and evaporation (Faye et al. 2003). Hydrodynamics arc essentially governed by the penetration of the tidal wave and the strong evaporation regime, which develop in the vast system of the interconnected tributaries (locally known as 'bolons') and the mangrove forest within the delta (Barusseau et al. 1995).
The study site consists of a soil catena of -1.5 km length and is 50 km from the sea, 20 km west of Kaolack, a town beside a tributary of the Saloum River. The transect is oriented east--west and runs across the tributary with a general slope of 0.5%. At the lower part of the toposequence (floodplain), the soils are flooded and entirely devoid of vegetation. The area was formerly occupied by mangrove vegetation, with plant remnants as a prominent landscape feature (Lezine 1997). Nearby, salt extraction activities using tiny evaporated ponds indicate a high level of salinity in the groundwater system.
Field observations and measurements
The selected catena was placed in such a way that influences of the landscape position could be reflected in soil properties. Nine soil profiles were distributed along the toposequence, three on each landscape position, i.e. floodplain, the low terrace and the middle terrace. All soil profiles were described in the field during the dry season (January--February 2007 and April-May 2008) according to the World Reference Base (IUSS Working Group WRB 2006). The Munsell soil colours were determined on moist samples. The reaction of pyrite to 30% hydrogen peroxide ([H.sub.2][O.sub.2]) was measured. The peroxide test devised by van Beers (1962) involved noting the degree of effervescence after addition of 1 mL of 30% [H.sub.2][O.sub.2] to the soil (Smith et al. 2003). The reaction of carbonate to 10% hydrochloric acid (HC1) was also measured in all samples. Oxidation was prevented by keeping samples in anaerobic conditions (hermetic plastic bags) before analysis. In total, 45 soil samples, one per identified horizon, were collected for laboratory analyses. In addition, 15 samples enriched with concretions, mottles and iron pipes were carefully taken from subsoils where redoximorphic features (RMFs), such as jarosite and iron pedofeatures, appear more conspicuous.
The topography of the catena was determined using a theodolite. The groundwater level was estimated in the soil profiles by means of a measuring tape, twice a week during the dry season (January-February 2007 and April-May 2008). Electrical conductivity (EC) and pH were estimated with field EC and pH meters (model 8120; WTW GmbH, Germany,) before repetition in laboratory under standard conditions.
Six soil profiles were selected for mineralogical analysis. They are labelled P1 and P2 in the floodplain, P4 and P5 in the low terrace, P7 and P8 in the middle terrace. All are representative of local soil formation and mineral distribution. Transitional profiles (P3, P6, and P9) were excluded from this analysis because of the scarcity of iron and jarosite pedofeatures.
Soil and mineral analyses
Bulk soil samples were air-dried at room temperature and passed through a 2-ram mesh sieve to obtain the fine earth fraction. Particle-size analysis was performed after removal of carbonates and organic matter by treatment with HC1 (pH 4.5) and [H.sub.2][O.sub.2] (10%), respectively, and of excessive salts by repeated addition of deionised water, centrifugation and decantation until the EC dropped below 40pS [cm.sup.-1] (Schlichting et al. 1995). After subsequent addition of ammonia (N[H.sub.3]) for water dispersion, overnight shaking and ultrasonic treatment, the sand fractions (63-2000 [micro]m) were obtained by wet-sieving, and the silt (2-63 [micro]m) and clay (<2[micro]m) fractions were separated by pipette analysis after Kohn (Schlichting et al. 1995). Bulk density was measured for all horizons using a steel cylinder of volume 100 [cm.sup.3]. Soil samples were then oven-dried at 105[degrees]C and weighted to determine the soil bulk density (gem 3) (Schlichting et al. 1995).
Soil pH was measured in water at a soil: solution ratio of 1 :2.5. The EC was determined in a 1:5 soil: water extract. Total Fe ([Fe.sub.t]) was determined using a SRS 200 X-ray fluorescence spectrometer (Siemens AG, Munich, Germany). Dithionite-soluble Fe ([Fe.sub.d]) was analysed by atomic absorption spectrometry. Oxalate-extractable Fe ([Fe.sub.o]) was measured by inductive coupled plasma-optical emission spectrometry. Total sulfur (St) was measured with a Vario Max Elemental analyser after combustion at 1150[degrees]C. Water-soluble sulfate (S[O.sup.2-.sub.4]) was measured in a 1:20 soil-water extract and determined by using an ion chromatograph (Dionex 2000i; Dionex, Sunnyvale, CA, USA) (Schlichting et al. 1995).
X-Ray difffactometry of fine earth was determined as powder. Clay mineralogy was determined by X-ray diffraction of oriented clay specimens using a D-500 instrument (Siemens AG) with Cu Ka radiation. Different treatments of the clay samples included potassium (K) and magnesium (Mg) saturation, heating to 110[degrees]C, 220[degrees]C, 400[degrees]C and 600[degrees]C of the K-saturatcd samples and glycerol salvation of the Mg-saturated samples. The amounts of the different clay minerals were estimated on a semi-quantitative basis by using the computer package Diffrac AT V3.3 (Siemens AG). For this purpose, soil profiles were divided into three parts: topsoil, central horizons, and subsoil. Clay minerals are individually presented to assess their abundance in soil profiles. The method assumes that the available minerals in the sample sum to 100% and the individual mineral is a fraction of the total. Ulite--smcctite mixed layers (I-S) were excluded from this semi-quantification because of the trace amounts shown in the entire toposequence. This fraction was therefore subtracted from the total sample (100%). Percentages of clay minerals were then recalculated from the total of sample without the I-S content. A scanning electron microscope (SEM) (LEO 420; LEO Electron Microscopy Ltd, Cambridge, UK) equipped with a field emission cathode and coupled to an energy dispersive X-ray (EDX) INCA 400 system (Oxford Instruments, Abingdon, UK) was used to confirm the elemental composition.
The soils differ in solum thickness, texture, water saturation and mottling. Figure 2 displays a soil distribution pattern reflecting soil-formation processes on different topographic positions. The presence of Glcyic Solonchaks in the floodplain (PI, P2 and P3), Haplic Gleysols in the low terrace (P4, P5 and P6) and Endogleyic Arenosols in the middle terrace (P7, P8 and P9) shows the influence of hydrodynamic processes on local soil formation. Even the middle terrace shows endogleyic influence in all Arenosols.
Floodplain soils exhibit fluvic properties. However, strong salinisation and the high tidal influence to the watertable prevent Fluvisol development. Therefore, the Solonchaks are glcyic, hyposaline or hypcrsaline and generally sulfatic. High levels of water saturation and sulfidic material are prominent features of the floodplain soils. Also, the conspicuous yellow mottles along channel roots and on ped faces throughout PI and P2 profiles suggest the presence of jarosite (Fig. 2a). Iron enrichment and contrasted soil texture characterise the low terrace. The generalised reddish colour of the subsoil is due to the presence of well-crystallised iron oxides (Fig. 2b). The interesting feature in the lower terrace is that Gleysols are often strongly acidified and thionic. Because of the lower groundwater table in the middle terrace, iron mottles and concretions are limited to, and less extensive and visible in, the lower profile (Fig. 2c). The floodplain can therefore be considered as the more dynamic site regarding hydrological behaviour. In relation to solum thickness, soil texture and other morphological features, P3 can be considered as a transitional profile between floodplain and low terrace and P6 between low terrace and middle terrace.
Solonchaks in the floodplain show a fine texture (silt clay loam). The clay content is higher in the Bjv horizon of PI and the Cjr horizon of P2, but increases continuously with depth within P3. The silt fraction decreases with depth throughout all profiles, with higher content in the topmost horizons than the underlying horizons. The sand fraction shows irregular depth distributions throughout all profiles. However, a coarser texture with more sand confirms the transitional position of P3. The fine texture of the floodplain soils supports less permeability in the lowest landscape position, promoting thereby continuous upward fluxes of the saline groundwater (presence of salt crusts on soil surface). The low terrace soils arc of sandy (clay) loam and become finer at depth. Soil material becomes coarser in the Endogleyic Arenosols in the middle terrace, from loamy sand to sand (Fig. 3). Along the toposequencc, floodplain and middle-terrace soils are of more homogeneous texture than low terrace soils. The relatively low bulk density (BD) throughout floodplain profiles ([less than or equal to] 1.2g [cm.sup.-3] in the topsoil and subsoil) indicates a very loose soil material. However, it increases slightly ([greater than or equal to] 1.2g [cm.sup.-3]) in the central horizons of the profiles (Table 1). The low terrace profiles show increasing BD with depth, which is generally the case for natural soils. Values are [less than or equal to] 1.2g [cm.sup.-3] in the topsoil because of the presence of organic compounds, but range from 1.3 to 1.6g [cm.sup.-3] in the more compact subsoil (Table 1). Middle-terrace profiles show the highest BD values ([greater than or equal to] 1.5 g [cm.sup.-3] at all depths). This confirms the presence of compact soil material in this landscape position. Changes in soil BD therefore follow changes in soil texture (Table 1). Somewhat surprising are the higher BD values in middle-terrace soils (loamy sand to sand texture), which arc supposed to be looser than floodplain soils because of a more sandy texture. However, high BD values are often associated with loamy soils, where clay and silt can fill the large voids between sand grains, optimising the packing of the soil matrix (Schactzl and Anderson 2005), whereas clayey soils have many micropores, which helps to explain their lower bulk densities.
The groundwater depths follow changes in topography. The floodplain site is permanently inundated by hypcrsaline water from the Saloum River. The shallow, saline groundwater table drops daily up to 15-35 cm below the soil surface, creating aerobic conditions in the topsoil. The low-terrace soils arc seasonally flooded and the presence of a fluctuation zone in the profiles reflects the dynamic of the groundwater table. At ~25 cm below the soil surface in the rainy season, the groundwater table drops to a depth of at least 80 cm during the dry season. The groundwater table is deeper ([greater than or equal to] 150 cm) in the middle-terrace profiles. Its upper boundary is within the B1 horizons of soil profiles (Fig. 2). Variations in the groundwater level lead to changes in soil colour, indicating the presence of RMFs at various depths according to the landscape position. In this study, the RMFs are subdivided into redox concentrations or redox depletions. Redox concentrations are features formed when the Fe oxides or hydroxides have accumulated at a point or around a large pore such as a root channel. They have been defined as 'bodies of apparent accumulation of Fe-Mn oxides and hydroxides' and contain more [Fe.sup.3+] oxides and hydroxides than were found in the soil matrix originally (Vepraskas et al. 1995). Redox depletions are zones formed by loss of Fe and other components. They have been defined as "bodies of low chroma (<2), having values of 4 or more where Fe-Mn oxides alone have been stripped out or where both Fe-Mn and clay have been stripped out' (Vepraskas et al. 1995). The high groundwater level in the floodplain involves strong reduction processes. Sinking groundwater level generates alternate aerobic and anaerobic conditions in the soil upper horizons (oxidation--reduction). This hydrological behaviour supports the location of RMFs at shallow depth visible after natural drainage of the site. According to the colour of the matrix and mottles, and the above-mentioned RMF classification, the soils show mixed redox depletions and redox concentrations throughout profiles. RMFs remain more conspicuous in low-terrace soils, where they appear in a clear distribution trend throughout profiles. Redox depletions in the subsoil (grey to dark greyish colour) accompany higher redox concentrations within the fluctuation zone at 25-80 cm (with yellowish and reddish mottles). The deeper groundwater table in the middle-terrace position explains the location of RMFs only in the subsoil of soil profiles (~110cm depth); RMFs occur in the form of orange mottles and rusty iron concretions within B1 horizons (Fig. 2).
Soil pH decreases downward in the floodplain profiles. Values are in the neutral pH range, 6-8 throughout all profiles, but pH drops below 5 in the subsoil of P1 and P2 and reaches 5.5 in the B11 horizon of P3. Values of laboratory-measured pH are close to those field-measured (Table 1). The low-terrace soils show the lowest pH values (3-5), except for the topmost horizon of P6 (Ahz) where pH reaches 5.9. Apart from the latter profile where a downward decrease of pH was observed, values remain generally constant with depth in the intermediate landscape position (Table 1). The laboratory-measured pH values are lower than the field-measured values (3-5 v. 5-6), caused most probably by the oxidation of pyrite during storage or preparation of samples. Drying and storage can significantly change soil pH values, but it is not easy to predict the results of drying and storage (Bloom et al. 2005). Bartlett and James (1980) reported a 0.6-unit increase after drying a mineral soil of pH 5.0. Soil pH increases generally with depth in the middle-terrace profiles. The laboratory-measured pH values are 1 unit lower than the field-measured ones (5-6 v. 6-7) (Table 1).
The highest EC values are observed in the floodplain soils (range 5.5-55.5 dS [m.sup.-1]). Particularly high values in the upper part of profiles suggest a significant incidence of shallow saline groundwater and surface salt crusts (Table 1). This trend reflects intense evaporation caused by the high temperatures (mean average temperature 29[degrees]C), and subsequent capillary rise of the groundwater during the dry season on this lowest and bare landscape position. Low-terrace profiles show EC values of 0.5-10 dS [m.sup.-1]. No consistent distribution trend of EC is noted in these profiles (Table 1). The lowest EC ([less than or equal to] 2.5 dS [m.sup.-1]) values are in the middle-terrace profiles. Apart from the subsoil of P9, which shows the highest EC (2.5 dS [m.sup.-1]), values are generally low ([less than or equal to] 1 dS [m.sup.-1]) and almost stable with depth on this highest landscape position (Table 1).
Iron and sulfur in soils
The highest total Fe amounts ([Fe.sub.t]) are in the floodplain profiles (Table 1). The distribution trend is the same throughout all profiles, with the highest Fe, concentration in the central horizons ([Fe.sub.t], >30 g [kg.sup.-1] in Bjz and Bjv for P1 and Az and Cjr for P2). Low-terrace profiles show a global downward increase of [Fe.sub.t]. The highest values are obtained in the B1 horizons of all profiles (zone of capillary fringe mottling--gleying) (Table 1). Values of [Fe.sub.t], are, however, globally lower in this location than the floodplain. Middle-terrace profiles show decreasing (P7) and increasing (P9) [Fe.sub.t], values with depth (Table 1). An interesting feature in this highest landscape position is the high concentration of [Fe.sub.t], in the 2B1 horizon of P8. The massive presence of iron mottles and concretions at this depth most likely supports this finding. It may be concluded that [Fe.sub.t], is mainly concentrated in B horizons. However, difference in soil thickness along the toposequence induces some disparities in the depth of B horizons and supports, therefore, the high concentration of [Fe.sub.t], at shallow depth in the floodplain profiles, in the central horizons of low-terrace profiles, and in the subsoil of middle-terrace profiles (Fig. 2).
Pedogenic iron oxides and hydroxides ([Fe.sub.d]) arc the crystalline form plus the amorphous or poorly crystalline forms of [F.sub.e]. Floodplain soils show the highest [Fe.sub.d] values, in the Bj and Cj horizons of P1 and P2, respectively (Table 1). Values of [F.sub.ed] decrease with depth in the low-terrace profiles, except throughout P6 where they increase (Table 1). Middle-terrace profiles contain the lowest amounts of [Fe.sub.d]. In contrast to the [Fe.sub.t], distribution trend, the highest Fed value (4.6 g [kg.sup.-1]) was obtained in the Ap horizon of P7, rather than in the 2B1 horizon of P8 (Table 1). The values of [Fe.sub.d] arc far lower in the low-terrace profiles than the floodplain profiles. Because the [Fe.sub.d] in these sulfate-rich soils generally includes iron from jarosite and ferrous sulfate, as well as from iron oxides, the higher sulfate (and jarosite) content of floodplain profiles probably support this difference.
The [Fe.sub.o] represents the amorphous form of iron and is, therefore, included in the Fed fraction. Apart from P3, [Fe.sub.o] values decrease continuously with depth in the floodplain profiles (Table 1). The highest values (>3 g [kg.sup.-1]) were recorded in the topmost horizons, characterised by important accumulation of salt crusts. The opposite trend noted within P3 is probably linked to the accumulation of organic matter in the subsoil of this profile (Bhr horizon). Low-terrace profiles show decreasing [F.sub.o] values with depth. Middle-terrace profiles show no discernible trend regarding the distribution of Fen with depth (Table 1).
The values of [S.sub.t] increase seaward. The highest concentrations (7-14 g [kg.sup.-1]) were obtained in the topmost horizons of floodplain profiles. Values decrease strongly in the underlying horizons but increase again in the sulfidic subsoil (Table 1). Values of S, are lower in low-terrace profiles (0.2-5 g [kg.sup.-1]) and decrease with depth on this intermediate landscape position, except for P6, which shows a significant concentration of [S.sub.t] (8g [kg.sup.-1]) in the central B11 horizon (Table 1). The lowest [S.sub.t] values were obtained in middle-terrace profiles (Table 1). Sulfate measurements show a similar distribution trend, with the highest value (2.6 [cmol.sub.c] [kg.sup.-1]) in floodplain profiles (Azm horizon of P2) and the lowest value (0.001 [cmol.sub.c] ]kg.sup.-1]) in the middle terrace (Table 1). This confirms the marine origin of elemental sulfur and sulfate in the studied soils.
Soil and clay mineralogy
The bulk mineralogy is dominated by quartz. Downstream soils (floodplain and low terrace) contain high amounts of halite. Feldspars, mainly albite (NaAl[Si.sub.3][O.sub.8]), are identified only in the floodplain and low terrace profiles. The proximity to the estuary explains the greater amounts of halite and albite in the floodplain soils. Pyrite is present in the floodplain, whereas jarosite is detected in the floodplain and the low terrace. Hematite ([Fe.sub.2][O.sub.3]) is detected only in the low terrace, and lepidocrocite is present in all the profiles (Fig. 4).
Table 2 shows a semi-quantitative estimation of the clay mineral suite in the toposequence. Kaolinite dominates the minerals present, with strong peaks and nearly constant amounts with depth throughout all profiles (Table 2). However, differences arc observed in its spatial distribution. The greatest amounts were obtained in the low-terrace profiles, with ~84% (on average) of the clay mineral assemblage, compared with 72% in the floodplain, and 62% in the middle terrace (Table 2). Kaolinite is mainly inherited from the Continental Terminal and not only dominating in abundance, but also having a high crystallinity. Smectite is also found in all of the soils but it shows by far the highest values in the middle terrace--up to 35% (Table 2). In the low terrace, because of the aggressive environment, smectite is strongly reduced, Illite is found in the floodplain, which shows that it may be brought in by the marine--coastal environment. Its presence in the low terrace and middle terrace is restricted to traces; 1-S shows only trace amounts in the entire toposequence.
Iron oxide dynamics in soils
The presence of iron oxides in the low-terrace profiles was first suggested by the prominent RMFs within the soil matrix of the sublayers. It was then investigated through the [Fe.sub.o]/[Fe.sub.d] ratio, sometimes called the Fe 'activity ratio' (Blume and Schwertmann 1969); a ratio [less than or equal to] 0.2 may be indicative of the formation of lepidocrocite, hematite and goethite (Igwe et al. 2005). A comparison of the order of the [Fe.sup.III] oxides, as revealed by the active iron ratio ([Fe.sub.o]/[Fe.sub.d]), in the iron accumulation zones (B horizons) of the low-terrace and floodplain profiles provides interesting results. In the floodplain profiles, the [Fe.sup.III] oxides (in the presence of jarosite) have active iron ratios [less than or equal to] 0.8 and are therefore relatively poorly crystalline. In low-terrace profiles, the active iron ratios in the zone of maximum iron accumulation were [less than or equal to] 0.1, representing a high degree of crystallinity. This difference may be due to continuing pyrite oxidation and upward movement of [Fe.sup.2+] in the floodplain profiles, and cyclic reduction and oxidation caused by movements in the watertable. Both processes would be expected to favour a poorly crystalline form, because [Fe.sup.III] oxides recently formed by the oxidation of [Fe.sup.2+] are poorly crystalline (Willett and Higgins 1978; Willett and Walker 1982). In low-terrace profiles, the scarcity of jarosite in the iron accumulation zone suggests an almost complete migration of [Fe.sup.2+] by oxidation of the pyritic sediment. Therefore, the [Fe.sup.III] oxides of low-terrace profiles have had an environment favouring crystallisation because of a higher landscape position than the floodplain, which reduces inundations by seawater. Soils subjected to frequent, periodic surface flooding contain more oxalate-extractable iron than rarely flooded soils (Willett et al. 1978), and iron oxides of rarely flooded soils arc more crystalline than those frequently flooded (Habibullah et al. 1971). Hematite was detected only on the low-terrace soils. Torrent et al. (1982) investigated the effect of the relative humidity (RH) on the crystallisation of goethite and hematite from ferrihydrite at 45[degrees]C. They found that with decreasing relative humidity, more hematite was formed. The aggressive environment in the low terrace (pH 3-5) may also support the presence of hematite in such a topographic position, because soil acidity is demonstrated to favour iron oxide formation. The effect of pH in a given [Fe.sup.III] system on the formation of goethite and hematite from ferrihydrite was demonstrated by Knight and Sylva (1974) and by Schwertmann and Murad (1983). The latter authors found that at 25[degrees]C as the pH drops from 8 to 4, the hematite/(hematite + goethite) ratio decreased from 0.70 to 0.04, whereas below pH 4, hematite again increased and above pH 8 goethite increased. In addition, the lower organic matter content in the iron-accumulated horizons of the low-terrace profiles (0.2%) compared with the floodplain ones (1.4%) may explain the difference in iron oxide crystallinity between the two landscape positions. As well as indirectly affecting [Fe.sup.III] oxide crystallinity as a reductant, organic matter also directly inhibits crystallisation (Schwertmann 1966; Schwertmann et al. 1968).
Jarosite formation and crystallinity
The X-ray patterns (Fig. 4) show higher amounts of jarosite in the upper layers of the floodplain profiles than in the upper layers of the low terrace. Soils in low topographic positions contain pyrite, which converts to jarosite upon drying (oxidation), giving rise to bright-yellow streaks along root channels in the original black-grey matrix (Fig. 2a). Jarosite presence in the floodplain and in low terrace soils is confirmed by the SEM and EDX spectra of some pale-yellow mottles from the oxidising zone of the floodplain and low-terrace profiles. Differences in crystallinity are noted according to the landscape position. Jarosite crystals appear nearly amorphous and less structured in the floodplain samples (Fig. 5a), contrasting with the definite angular blocky structure they show in low terrace samples (Fig. 5b). The hydrological behaviour supports formation of low-crystalline jarosite. The low-terrace soils undergo annual cycles of waterlogging and drying, promoting periodic reduction and oxidation caused by shifting of soil materials between anaerobic and aerobic conditions. Pyrite is not detected in these soils but its past presence is inferred by the presence of jarosite. The high crystallinity of jarosite (Fig. 5b) and the low pH values (3-5) due to a somewhat higher topographic position support the maturation of the soil material, with almost complete pyrite oxidation and sulfuric acid production. The soils in the low terrace have probably passed through the potential ASS stage to the ripe one, marked by advanced crystallinity of jarosite and iron oxides. In agreement with a theoretical stability relationship, jarosite is considered stable only under relatively oxidised (Eh >400 mV) and acid (pH range 2-4) conditions (van Breemen 1982). At pH >4, jarosite is mctastable and ultimately hydrolysed to iron oxide, predominantly goethite and sometimes hematite (van Breemen 1976; Dent 1986; Furian el al. 2011). Keene et al. (2010) noted that with increasing pH and decreasing Eh under tidal inundation, reduction of [Fe.sup.III] minerals occurred and jarosite became increasingly unstable. Hydrolysis is generally enhanced by leaching and by a supply of bases.
The oxidation of pyrite and the redistribution of the main oxidation products ([Fe.sup.2+] and S[O.sub.4.sup.2-]) seem the most critical pcdogenic process in the formation of jarosite and hematite in saline CASS. Assuming that pyrite is the precursor mineral of jarosite and hematite in our studied soils, we suggest that [Fe.sup.2+] produced during pyrite oxidation probably migrates upwards from the reducing sublayers to the oxidised layers, where it accumulates in the form of immobile [Fe.sup.2+] oxides (Fig. 6). This explains the presence of hematite in the mottled horizon starting from ~20 cm below the surface of the low-terrace profiles. The [Fe.sup.2+] may also migrate towards the soil surface and then become immobile through the formation of jarosite after combination with S[O.sub.4.sup.2-]. The addition of potassium from the seawater completes the process. Because the iron in jarosite is [Fe.sup.III], as in iron 'oxides', both jarosite and iron 'oxides' tend to occur in the more oxidised parts of soils rather than in the more anaerobic parts (Fanning and Fanning 1989).
Major soil-forming processes
The formation of a soil profile as viewed by Simonson (1978) represents the combined effect of additions to the ground surface, transformations within the soil, vertical transfers (upward or downward) within the soil and removals from the soils (Schaetzl and Anderson 2005). Our results suggest that soil formation in the study area essentially proceeds through transfers and matter additions. The transfers include intrapedon as well as interpedon transfers, whereas matter additions essentially concern acolian dust inputs (Fig. 6).
Intrapedon processes encompass the internal pedogenic processes, i.e. translocations and transformations within the soil profile (Schaetzl and Anderson 2005); in the present study, these processes centre on the dynamics of Fe and S ions in soils. They bring about three major soil-forming processes characteristic of local soil development: gleysation, sulfidisation and sulfurisation.
Gleysation is a common process in the studied soils, and significantly affects local soil formation. It indicates the effects of wetness on soil development and, in particular, soil morphological reflections of varying degrees of wetness (Buol et al. 1997). The presence of Gleyic Hypersalic Solonchaks or Flyposalic Solonchaks (Sulfatic) in the floodplain, Haplic Gleysols (Thionic) in the low terrace and Endogleyic Arenosols in the middle terrace substantiates the importance of gleyic processes for all landscape positions. The main characteristic of these processes is reflected by the occurrence of RMFs throughout all of the soil profiles. They occur in the B1 (capillary fringe mottling, gleying) or Bj (jarosite accumulation) horizons. The solum thickness introduces some differences and explains the location of RMFs at various depths within the soil profiles.
Sulfidisation explains the presence of potential ASS in the floodplain and represents the major soil-forming process in the low topographic position. However, it is important to note that sulfidisation and gleysation generally co-occur in these highly reduced soils. Unlike gleysation, which simply involves Fe ions in the present study, both Fe and S ions are involved in sulfidisation processes. Sulfidic material also forms the subsoil of low-terrace profiles. It was not identified in the middle terrace profiles, however, because of the predominance of aerobic conditions. The weak reduced conditions prevailing in the subsoil promote the formation of [Fe.sup.2+] ions but not [S.sup.2-] ions. Stronger reduced conditions are needed for the production of S2 ions than for the production of [Fe.sup.2+] ions (Fanning and Fanning 1989).
Sulfurisation remains a contemporary pedogenic process occurring in the floodplain and low-terrace profiles. It supports the presence of jarosite and lepidocrocite in topsoils and central horizons of floodplain soils, with jarosite hydrolysis being more favourable in low-terrace soils (actual ASS) than in the floodplain (potential ASS). Iron changes between [Fe.sup.2+] in the reduced subsoil and [Fe.sup.3+] in the oxidised topsoil, and pyritc is almost completely oxidised to jarosite in the low terrace soils. Jarosite hydrolysis leads to further sulfuric acid production and iron oxide crystallisation, mainly hematite. This supports the lowest pH values (3-5) yielded in the low terrace soils.
The importance of intrapedon processes was largely demonstrated in the above analysis; however, interpedon processes, defined as losses from and additions to soil profiles (Schaetzl and Anderson 2005), are also important in the coastal landscapes. They promote the transport and deposition of elements from topographic high position to topographic low position through drainage and leaching. Salinisation, colluvio-alluviation and lateral eluviation are the main interpedon processes identified in the studied coastal area.
Salinisation is a continuous process in the Saloum River basin. Our study demonstrated the fundamental linkage between landscape position and soil salinity, exemplified by a lateral salt distribution pattern appreciable at landscape scale. The concentration of water-soluble cations and anions in soils decreases steadily from the floodplain to uplands along the investigated toposequence. This reflects the influence of the topographic setting, which conditions the intensity of seawater intrusions. The floodplain profiles are more affected by salinisation processes, and the presence of Hyposalic Solonchaks and Hypersalic Solonchaks indicates the predominant saline features. Frequent tidal intrusions and a saline, shallow groundwater table support raised salt concentrations in both topsoils and subsoils. Strong evaporation processes induced by high temperatures and a continuous capillary rise due to a fine soil texture support the permanent build-up of salt crusts on the soil surface. Because of a somewhat higher landscape position, the low-terrace soils are only seasonally inundated by seawater. The groundwater table is also lower and the soil texture less uniform than the floodplain soil profiles. This reduces the intensity of capillary rise and therefore the formation of surface salt crusts. The distal position to the sea and the high landscape position support the limited influence of seawater inundations in the middle-terrace soils. Salinity is mainly caused by lateral water flow, though considerably reduced. Although the vertical processes are linked to the fluctuations and depths of the groundwater, they remain subsidiary and depend largely on the intensity of tidal inundations, which is governed by the landscape position.
Colluvio-alluviation remains the major interpedon process regarding the build-up and characteristics of soil material in the Saloum River basin. The sulfidic material, identifiable in the floodplain by a soft and buttery texture, mid-dark grey colour, a positive peroxide test and an offensive [H.sub.2]S odour, forms the subsoil of low-terrace profiles. Topography, as the most determining parameter, controls the marine influence. This supports the absence of any sulfidic material in the middle terrace. The latter site is marked by a homogeneous, sandy parent material resulting from the erosion of the nearby uplands (high terrace and plateau). The colluvial material, transported by wind and deposited downslope, builds the topsoil of the low-terrace profiles. It was also identified, to a lesser extent, in the floodplain, where it partly explains the sand- and silt-rich topsoil. The low-terrace profiles provide a better illustration of interpedon transfers, with alluvial deposits overlain by colluvial sediments. The intermediate landscape position therefore favours the contribution of marine and terrestrial sediments on the formation of the low-terrace pedon.
The influence of the lateral eluviation on element transport is demonstrated in the present study. Depletion and lateral transport of bases and silica (Si[(OH).sub.4]) from the upland to the lowland soils is a critical process in the formation of pedogenic clay minerals, mainly kaolinite and smectite, from tropical regions with wet-dry seasons. Kaolinite forms in uplands soils subjected to desilication, introduced by intense leaching processes, whereas silica and bases leached from the higher landscape positions accumulate downslope and enhance smectite formation. Clay mineralogy distribution may therefore give clear insight into geochemical processes in catenas along a climatic transect. The above considerations suggest that lateral eluviation processes are mostly accountable for the formation of pedogenic kaolinite in the middle terrace soils and smectite in the floodplain soils. Silication and smectite formation also help to explain the clayey floodplain soils compared with the sandy middle-terrace soils. The effects of lateral eluviation processes are, however, less perceptible in the low-terrace soils. The intermediate landscape position makes these soils less base-depleted than the middle terrace and less base-enriched than the floodplain. Differences are introduced by the contrasting redox cycles of Fe and S ions and the acid pedoenvironment, which promote the formation of kaolinite and iron oxides over smectite in this landscape position.
Colluvio-alluviation and lateral eluviation appear as active processes in the composition of the clay mineral suite for the studied soils. However, their contribution is different regarding the nature and origin of clay minerals. Colluvio-alluviation is more conducive to the distribution of terrigenous clay minerals (aeolian dust inputs), whereas lateral eluviation mainly affects the formation of pedogenic clay minerals.
A toposequence of CASS has been studied in the region of the Saloum River (west-central Senegal), from a soil perspective and from a mineralogical point of view, taking into account the topographic position of the studied soils. Soil distribution is established according to the following pattern: Gleyic Hyposalic and Hypersalic Solonchaks (sulfatic) in the floodplain, Haplic Gleysols (thionic) in the low terrace and Endogleyic Arenosols in the middle terrace. Intrapedon processes such as gleysation, sulfidisation and sulfurisation, as well as interpedon processes such as salinisation, colluvio-alluviation and lateral eluviation are identified and discussed. Strongly dependent on the landscape position, they combine and occur with varying intensities. In an unfavourable climatic and geographical context, the salinity and acidity constraints of CASS in the Saloum River appear very strong. Finally, knowledge of iron and sulfur reactions in soils enhances understanding of pedogenesis in CASS environments. Such knowledge may also help develop sustainable soil-management strategies, including optimising water management, at local and regional level.
Received 4 February 2014, accepted 19 June 2014, published online 10 October 2014
Barbiero L, Mohamedou AO, Roger L, Furian S, Aventurier A, Remy JC, Marlet S (2005) The origin of vertisols and their relationship to acid sulfate soils in the Senegal valley. Catena 59, 93-116. doi: 10.1016/j.catena.2004.05.007
Bartlett R, James B (1980) Studying dried stored soil samples. Some pitfalls. Soil Science Society of America Journal 44, 721-724. doi: 10.2136/sssaj1980.03615995004400040011x
Barusseau JP, Ba M, Descamps C, Diop KHS, Giresse P, Saos JL (1995) Coastal evolution in Senegal and Mauritania at [10.sup.3], [10.sup.2] and [10.sup.1]-year scales: Natural and human records. Quaternary International 29-30, 61 73. doi: 10.1016/1040-6182(95)00008-7
Bigham JM, Schwertmann U, Pfab G (1996) Influence of pH on mineral speciation in a bioreactor simulating acid mine drainage. Applied Geochemistry 11, 845-849. doi:10.1016/S0883-2927(96)00052-2
Bloom PR, Skyllberg UL, Sumner ME (2005) Soil acidity. In 'Chemical processes in soils'. Book Series 8. (Eds MA Tabatabai, DL Sparks) pp. 411-459. (Soil Science Society of America: Madison, WI, USA)
Blume HP, Schwertmann U (1969) Genetic evaluation of profile distribution of aluminium, iron, and manganese oxides. Soil Science Society of America Proceedings 33, 438-444. doi:10.2136/sssaj1969. 03615995003300030030x
Boivin P (1991) 'Caracterisation physique des sols sulfates acides de la vallee de Katoure (Basse Casamance, Senegal). Etude de la variabilite spatiale et relation avec les caracteristiques pedologiques.' Etudes et Theses. (Orstom: Paris)
Boman A, Astrom M, Frojdo S (2008) Sulfur dynamics in boreal acid sulfate soils rich in metastable ironsulfide--the role of artificial drainage. Chemical Geology 255, 68-77. doi:10.1016/j.chemgeo.2008.06.006
Buol SW, Hole FD, McCracken RJ (1997) 'Soil genesis and classification.' (Iowa State University Press: Ames, IA, USA)
Burton ED, Bush RT, Sullivan LA (2006) Sedimentary iron geochemistry in acidic waterways associated with coastal lowland acid sulfate soils. Geochimica et Cosmochimica Acta 70, 5455-5468. doi: 10.1016/j.gca. 2006.08.016
Burton ED, Sullivan LA, Bush RT, Powell B (2008) Iron-sulfide and trace element behaviour in sediments of Coombabah Lake, southern Moreton Bay (Australia). Marine Pollution Bulletin 56, 1353-1358. doi: 10.1016/ j.marpolbul.2008.04.012
Bush RT, Sullivan LA, Lin C (2000) Iron monosulfide distribution in three coastal floodplain acid sulfate soils, eastern Australia. Pedosphere 10, 237-245.
Claff SR, Sullivan LA, Burton ED, Bush RT (2010) A sequential extraction procedure for acid sulfate soils: Partitioning of iron. Geoderma 155, 224-230. doi: 10.1016/j.geoderma.2009.12.002
Deckers J, Raes D, Merckx R, Diallo A (1993) The fate of salic and thionic fluvisols under irrigated rice in the Senegal river delta. Pedologie 43, 389-401.
Dent DL (1986) 'Acid sulfate soils: a baseline for research and development.' (ILRI: Wageningen, The Netherlands)
Dent DL, Ahmed F (1995) Resurrection of soil survey. A case study of acid sulfate soils on the floodplain of the River Gambia. I. Data validation, taxonomic and mapping units. Soil Use and Management 11, 69-76.
Dent DL, Pons LJ (1995) A worldperspective on acid sulfate soils. Geoderma 67, 263-276. doi: 10.1016/0016-7061(95)00013-E
Diop S (1990) 'La coteouestafricaine. Du Saloum (Senegal) a la Mellacoree (Rep. de Guinee).' Etudes et Theses. (Orstom: Paris)
Durr M, Wakelin SA, Rogers SL, White I, Macdonald BCT, Welch S (2006): Archealdiversityof an acid sulfate soil in coastal northern NSW. In 'Regolith 2006--Consolidation and dispersion of ideas'. (Eds RW Fitzpatrick, P Shand) pp. 63-66. (CRC LEME: Bentley, W. Aust.)
Fanning DS, Burch SN (1997) Acid sulfate soils and some associated environmental problems. Advances in GeoEcology 30, 145-158.
Fanning DS, Fanning MCB (1989) 'Soil: morphology, genesis, and classification.' (John Wiley and Sons: New York)
Fanning DS, Rabenhorst MC, Burch SN, Islam KR, Tangren SA (2002) Sulfides and sulfates. In 'Soil mineralogy with environmental applications'. (Eds JE Amonette, WF Bleam, DG Schulze, JB Dixon) pp. 229- 256. (Soil Science Society of America, Inc.: Madison, WI, USA)
Fanning DS, Rabenhorst MC, Balduff DM, Wagner DP, Orr RS, Zurheide PK (2010) An acid sulfate perspective on landscape/seascape soil mineralogy in the U.S. Mid-Atlantic region. Geoderma 154, 457-464. doi: 10.1016/j.geoderma.2009.04.015
Faye S, Faye SC, Ndoye S, Faye A (2003) Hydrogeochemistry of the Saloum (Senegal) superficial coastal aquifer. Environmental Geology 44, 127-136.
Fitzpatrick RW, Baker AKM, Raven M, Rogers S, Dcgens B, George R, Kirby J (2005) Mineralogy, biogeochemistry, hydro-pedology and risks of sediments, salt efflorescences and soils in open drains in the wheatbelt of Western Australia. In 'Regolith 2005--Ten Years of CRC LEME'. (Ed. 1C Roach) pp. 97-101. (CRC LEME: Bentley, W. Aust.)
Furian S, Mohamedou AO, Hammecker C, Maeght JL, Barbiero L (2011) Soil cover and landscape evolution in the Senegal floodplain: a review and synthesis of processes and interactions during the late Holocene. European Journal of Soil Science 62, 902 912. doi:10.1111/j.1365-2389.2011.01398.x
Habibullah AK, Greenland DJ, Brammer H (1971) Clay mineralogy of some seasonally flooded soils of East Pakistan. Soil Science 22, 179-190. doi:10.1111/j.1365-2389.1971.tb01605.x
Husson O, Verburg PH, Phung MT, van Mensvoort MEF (2000) Spatial variability of acid sulphate soils in the Plain of Reeds, Mekong delta, Vietnam. Geoderma 97, 1-19. doi: 10.1016/S0016-7061(00)00016-1
Igwe CA, Zarei M, Stahr K (2005) Mineral and elemental distribution in soils formed on the River Niger floodplain, eastern Nigeria. Australian Journal of Soil Research 43, 147-158. doi:10.1071/SR04046
IUSS Working Group WRB (2006) 'World Reference Base for soil resources. A framework for international classification, correlation and communication.' World Soil Resources Reports, 103. (FAO: Rome)
Johnston SG, Slavich PG, Hirst P (2005) Changes in surface water quality after inundation of acid sulfate soils of different vegetation cover. Australian Journal of Soil Research 43, 1-12. doi: 10.1071/SR04073
Johnston SG, Keene AF, Bush RT, Burton ED, Sullivan LA, Isaacson L, McElnea AE, Ahem CR, Douglas Smith C, Powell B (2011) Iron geochemical zonation in a tidally inundated acid sulfate soil wetland. Chemical Geologv 280, 257-270. doi:10.1016/j.chemgeo.2010.11.014
Jones EJP, Nadeau T-L, Voytek MA, Landa ER (2006) Role of microbial iron reduction in the dissolution of iron hydroxysulfate minerals. Journal of Geophysical Research 111, G01012. doi: 10.1029/2005 JG000089
Kalck Y (1978) Evolution des zones de mangroves a Senegal au Quatemaire. Etudes geologiques et geochimiques. PhD Thesis, Institut de Geologie Strasbourg, France.
Keene AF, Johnston SG, Bush R, Sullivan L, Burton L (2010) Reductive dissolution of natural jarosite in a tidally inundated acid sulfate soil: geochemical implications. In '19th World Congress of Soil Science, Soil Solutions for a Changing World'. 1-6 August 2010, Brisbane, Australia. (International Union of Soil Sciences)
Knight RJ, Sylva RN (1974) Precipitation in hydrolysed iron(III) solutions. Journal of Inorganic and Nuclear Chemistry 36, 591-597. doi: 10.1016/0022-1902(74)80119-3
Kraal P, Burton ED, Bush RT (2013) Iron monosulfide accumulation and pyrite formation in eutrophic estuarine sediments. Geochimica et Cosmochimica Acta 122, 75-88. doi: 10.1016/j.gca.2013.08.013
Lappartient JR (1985) Le continental terminal et le pleistocene ancien du bassin senegalo-mauritanien. Stratigraphie, sedimentation, diagenese, alterations. Reconstitution des paleorivages au travers des cuirasses. PhD Thesis, Aix-Marseille III University, Marseille, France.
Le Brusq JY, Loyer JY, Mougenot B, Carn M (1987) Nouvelles parageneses a sulfates d'aluminium, de fer et de magnesium, et leur distribution dans les sols sulfates acides du Senegal. Science du Sol 25, 173-184.
Lezine A-M (1997) Evolution of the West African mangrove during the Late Quaternary: A review. Geographiephysique et Quaternaire 51, 405-414. doi: 10.7202/033139ar
Lin C, Melville MD (1993) Control of soil acidification by fluvial sedimentation in an estuarine floodplain, eastern Australia. Sedimentary Geology 85, 271-284. doi:10.1016/0037-0738(93)90088-M
Marius C (1985) 'Les mangroves du Senegal et de la Gambie.' Travaux et Documents 193. (Orstom: Paris)
Marius C, Lucas J (1991) Holocene mangrove swamps of West Africa: sedimentology and soils. Journal of African Earth Sciences 12, 41-54. doi: 10.1016/0899-5362(91)90056-5
Michel P (1973) 'Les bassins des fleuves Senegal et Gambie: etude geomorphologique.' Memoires de I'Office de Recherche Scientifque et Technique d'outre-Mer (ORSTOM J 63). (Orstom: Paris)
Montoroi JP (1995) Mise en evidence d'une sequence de precipitation des sels dans les sols sulfates acides d'une vallee amenagee de Basse-Casamance (Senegal). Comptes Rendus de l'Academie des sciences, Paris 320, I la, pp. 395 402. [extended abstract in English]
Montoroi J-P (1996) 'Gestion durable des sols de la mangrove au Senegal en periode de secheresse. Dynamique de l'eau et geochimie des sels d'un bassin versant amenage.' Etudes et Theses. (Orstom: Paris)
Sadio S (1991) 'Pedogenese et potentialites forestieres des sols sulfates acides sales des tannes du Sine Saloum, Senegal.' (Orstom: Paris)
Schaetzl RJ, Anderson S (2005) 'Soils: genesis and geomorphology.' (Cambridge University Press: Cambridge, UK)
Schlichting E, Blume HP, Stahr K (1995) 'Bodenkundliches Praktikum.' (Blackwell Wissenschafts-Verlag: Berlin)
Schwertmann U (1966) Inhibitory effect of soil organic matter on the crystallization of amorphous ferric hydroxide. Nature 212, 645-646. doi: 10.1038/212645b0
Schwertmann U, Murad E (1983) The effect of pH on the formation of goethite and hematite from ferrihydrite. Clays and Clay Minerals 31, 277 284. doi: 10.1346/CCMN. 1983.0310405
Schwertmann U, Fischer WR. Papendorf H (1968) The influence of organic compounds on the formation of iron oxides. In 'Transactions 9th International Congress of Soil Science'. Adelaide, S. Aust. Vol. 1, pp. 645-655. (International Society of Soil Science and Angus and Robertson: Sydney)
Simonson RW (1978) A multiple-process model of soil genesis. In 'Quaternary soils. 3rd Symposium on Quaternary Research'. Toronto, Canada, 21-23 May 1976. (Ed. WC Mahaney) pp. 1-25. (Geo Abstracts: Norwich, England)
Smith J, van Oploo P, Marston H, Melville MD, Macdonald BCT (2003) Spatial distribution and management of total actual acidity in an acid sulfate soil environment, McLeods Creek, northeastern NSW, Australia. Catena 51, 61-79. doi: 10.1016/S0341-8162(02)00069-3
Souza-Junior VS, Vidal-Torrado P, Garcia-Gonzalez MT, Otero XL, Macias F (2008) Soil mineralogy of mangrove forests from the State of Sao Paulo, Southeastern Brazil. Soil Science Society of America Journal 72, 848-857. doi: 10.2136/sssaj2007.0197
Sullivan LA, Bush RT (1997) Quantitative elemental microanalysis of rough-surfaced soil specimens in the scanning electron microscope using a peak-to-background method. Soil Science 162, 749-757. doi: 10.1097/00010694-199710000-00008
Sullivan LA, Ward NJ, Bush RT, Burton ED (2009) Improved identification of sulfidic soil materials by a modified incubation method. Geoderma 149, 33-38. doi: 10.1016/j.geoderma.2008.11.019
Sylla M, Stein A, van Mensvoort MEF, van Breemen N (1996) Spatial variability of soil actual and potential acidity in the mangrove agroecosystem of West Africa. Soil Science Society of America Journal 60, 219-229. doi: 10.2136/sssaj1996.03615995006000010034x
Torrent J, Guzmann R, Parra MA (1982) Influence of relative humidity on the crystallization of Fe(III) oxides from ferrihydrite. Clays and Clay Minerals 30, 337-340. doi:10.1346/CCMN.1982.0300503
van Beers C (1962) 'Acid sulphate soils.' (ILRI: Wageningen, The Netherlands)
van Breemen N (1973) Soil forming processes in acid sulphate soils. In 'Proceedings International Symposium on Acid Sulfate Soils'. 11-20 August, 1972. (Ed. H Dost) pp. 66-129. (ILRI Publication: Wageningen, The Netherlands)
van Breemen N (1975) Acidification and deacidification of coastal plain soils as a result of periodic flooding. Soil Science Society of America Proceedings 39, 1153 1157. doi:10.2136/sssaj1975.036159950039000 60035x
van Breemen N (1976) 'Genesis and solution chemistry of acid sulfate soils in Thailand.' Agricultural Research Report 848. (Centre for Agricultural Publishing and Documentation: Wageningen, The Netherlands)
van Breemen N (1982) Genesis, morphology and classification of acid soils in coastal plains. In 'Acid sulfate weathering'. Special Publication 10. (Eds JA Kittrick, DS Fanning, LR Hossner) pp. 95-108. (Soil Science Society of America: Madison, WI, USA)
van Breemen N, Harmsen K (1975) Translocation of iron in acid sulfate soils: I. Soil morphology and mineralogy of iron in a chronosequence of acid sulfate soils. Soil Science Society of America Journal 39, 1140-1148. doi: 10.2136/sssaj1975.03615995003900060033x
Vepraskas MJ, Teets SJ, Richardson JL, Tandarich JP (1995) Development of red oximorphic features in constructed wetland soils. Wetlands Research, Inc. Technical Paper No. 5, pp. 1-12.
Vieillefon J (1977) 'Les Sols des mangroves et des tannes de Basse Casamance (Senegal).' Memoires 83. (Orstom: Paris)
Willett IR, Higgins ML (1978) Phosphate sorption by reduced and reoxidized rice soils. Australian Journal of Soil Research 16, 319-326. doi: 10.1071/SR9780319
Willett IR, Walker PH (1982) Soil morphology and distribution of iron and sulfur fraction in a coastal floodplain toposequence. Australian Journal of Soil Research 20, 283 294. doi:10.1071/SR9820283
Willett IR, Muirhead WA, Higgins ML (1978) The effects of rice growing on soil phosphorus immobilization. Australian Journal of Experimental Agriculture and Animal Husbandry 18, 270-275. doi: 10.1071/EA9780270
Wright JB, Hastings DA, Jones WB, Williams HR (1985) 'Geology and mineral resources of West Africa.' (Springer: Dordrecht, The Netherlands)
Aidara C. A. Lamine Fall (A,D), Jean-Fierre Montoroi (B), and Karl Stahr (C)
(A) Department of Geography, University of Ziguinchor, BP 523 Ziguinchor, Senegal,
(B) Institute of Research for Development (IRD), UMR 242, IEES Paris, 32 Avenue Henri Varagnat, 93143 Bondy, France.
(C) Institute of Soil Science and Land Evaluation, University of Hohenheim, 70593 Stuttgart, Germany. Corresponding author. Email: firstname.lastname@example.org
Table 1. Soil characteristics along the investigated toposequence BD, Bulk density; EC, electrical conductivity; [Fe.sub.t], [Fe.sub.d], [Fe.sub.0]: total, dithionite-soluble, oxalate- extractable iron; [S.sub.t] total sulfur Depth Horiz. Colour Texture (%) (cm) (moist) Sand Silt Clay Floodplain (PI), Gleyic Hyposalic Solonchak (sulfatic), 14[degrees] 04'27.4" N, 16[degrees] 11'17.4" W 0-1 Az 7.5Y4/1 24.5 56.7 18.8 1-4 Bjz 7.5YR5/6 25.6 37.4 37 4-23 Bjv 7.5YR4/2 17.9 34.8 47.2 23-60 Bv 10 YR 3/3 29.4 31.4 39.2 Floodplain (P2), Gleyic Hypersalic Solonchak (sulfatic), 14[degrees] 04'23.8" N, 16[degrees] 11'16.6" W 0-1 Azm 5 Y 3/1 + 22.8 45.7 31.5 5 YR 5/8 1-8 Az 2.5 Y 4/1 27.6 32.9 39.4 8-30 Cjr 2.5 YR 3/1 + 21.3 30.7 47.7 10 YR 7/6 30-60 Cr 7.5 YR 2/2 42.3 30.5 27.2 Floodplain (P3), Gleyic Hyposalic Solonchak (sulfatic), 14[degrees] 04'21.4" N, 16[degrees] 11'06.0" W, altitude 0.41 m a.s.l. 0-5 Ahz 7.5YR5/4 75.2 21.6 3.2 5-11 Cz 10YR7/2 69.8 25.9 4.3 11-17 C 10YR6/4 79.5 17.8 3.1 17-26 Bl1 7.5YR5/8 83.6 10.7 6.1 26-34 Bl2 Mottled 78.9 12.2 8.8 34-70 Bhr 7.5YR2/1 64.3 11.0 24.7 Low terrace (P4),Haplic Gleysol (thionic), 14[degrees] 04'19.4" N, 16[degrees] 10'59.2" W, altitude 1.15 m a.s.l. 0-2 Ah1 7.5 YR 3/3 57.2 26.4 16.3 2-16 Ah2 10 YR 4/4 57.9 18.3 23.7 16-38 Bl1 7.5 YR 4/3 61.8 18.3 20 38-62 2Bl2 10 R 4/3 45.8 23.3 30.9 62-90 2Bl3 5 Y 6/1 46.0 21.3 32.7 90-100 2Br1 2.5 Y 5/2 41.8 23.6 34.6 Low terrace (P5),Haplic Gleysol (thionic), 14[degrees] 04'20.3" N, 16[degrees] 10'59.5" W, altitude 1.17 m a.s.l. 0-2 Ahzl 10 YR 3/3 70.4 19.7 9.7 2-11 Ahz2 10 YR 3/4 62.5 18.3 19.1 11-29 Bwz 10 YR 4/6 46.7 22.4 31 29-51 Bl1 2.5 YR 4/6 50.0 21.4 28.6 51-81 Bl2 2.5 Y 6/2 48.7 21.3 30 81-92 Br1 2.5 Y 6/2 38.6 22.9 38.3 Low terrace (P6), Haplic Gleysol (thionic), 14[degrees] 04'21.2" N, 16[degrees] 10'51.6" W. altitude 1.51 m a.s.l. 0-4 Ahz 10YR2/3 82.1 12.8 5.1 4-32 Ahlz 10YR2/2 67.6 21.8 10.5 32-62 Bl1 10YR4/6 71.0 20.0 8.8 62-82 Bl2 10YR5/6 68.6 17.2 14.3 82-110 Br1 10YR4/6 61.7 12.4 26 Middle terrace (P7), Endogleyic Arenosol, 14[degrees] 04'15.9" N, 16[degrees] 10'47.5" W, altitude 4.15 m a.s.l. 0-26 Ap 10 YR 3/4 85.1 9.7 5.1 26-60 Ah 10 YR 3/4 79.8 14.0 6.2 60-108 Cw 7.5 YR 8/3 90.4 8.6 0.8 108-160 B1 7.5 YR 7/4 88.4 9.7 1.8 Middle terrace (P8), Endogleyic Arenosol, 14[degrees] 04'14.0" N. 16[degrees] 10'47.1" W, altitude 4.19 m a.s.l. 0-31 Ap 10 YR 2/2 77.6 15.3 7.1 31-62 AhC 10 YR 3/2 78.8 18.3 3 62-107 lCw 7.5 YR 7/3 82.1 15.9 2 107-138 2B1 10 YR 6/3 72.9 10.7 16.4 138-160 1B1 10 YR 8/2 83.5 12.5 4.3 Middle terrace (P9), Endogleyic Arenosol, 14[degrees] 04'14.0" N, 16[degrees] l0'47.2" W, altitude 4.34 m a.s.l. 0-36 Ap 10YR2/3 67.5 18.2 14.3 36-67 Ah 10YR3/3 64.6 20.3 15.1 67-88 Cv 10YR4/3 70.1 17.4 12.5 88-126 Be 10YR5/4 64.9 20.5 14.6 126-160 B1 2.5Y3/2 61.4 14.5 24.1 Depth BD pH EC (cm) (g (dS [cm.sup.-3]) Ca[Cl.sub.2] [H.sub.2]O [m.sup.-1]) Floodplain (PI), Gleyic Hyposalic Solonchak (sulfatic), 14[degrees] 04'27.4" N, 16[degrees] 11'17.4" W 0-1 0.8 8.2 8.4 26.5 1-4 1.0 7.8 7.9 11 4-23 1.2 6.1 6.1 12 23-60 1.0 4.5 4.5 15.5 Floodplain (P2), Gleyic Hypersalic Solonchak (sulfatic), 14[degrees] 04'23.8" N, 16[degrees] 11'16.6" W 0-1 1.2 7.8 7.9 55.5 1-8 1.4 7.8 8.0 9 8-30 1.2 7.3 7.5 8.5 30-60 1.0 5.0 5.0 14 Floodplain (P3), Gleyic Hyposalic Solonchak (sulfatic), 14[degrees] 04'21.4" N, 16[degrees] 11'06.0" W, altitude 0.41 m a.s.l. 0-5 1.2 7.6 7.8 12.5 5-11 1.5 7.0 7.3 7 11-17 1.5 6.2 6.5 5.5 17-26 1.4 5.4 5.5 5.5 26-34 1.4 6.2 6.3 7 34-70 1.2 6.2 6.2 16 Low terrace (P4),Haplic Gleysol (thionic), 14[degrees] 04'19.4" N, 16[degrees] 10'59.2" W, altitude 1.15 m a.s.l. 0-2 1.0 4.0 4.1 0.5 2-16 1.6 4.8 5.0 2.5 16-38 1.5 4.2 4.5 4.5 38-62 1.5 4.2 4.3 7.5 62-90 1.5 4.5 4.5 9.5 90-100 1.5 4.9 4.7 1 Low terrace (P5),Haplic Gleysol (thionic), 14[degrees] 04'20.3" N, 16[degrees] 10'59.5" W, altitude 1.17 m a.s.l. 0-2 1.2 4.5 4.4 1 2-11 1.5 3.9 3.9 0.5 11-29 1.5 3.8 3.9 4.5 29-51 1.6 3.7 3.7 3 51-81 1.5 3.8 3.9 2.5 81-92 1.5 4.0 4.1 2 Low terrace (P6), Haplic Gleysol (thionic), 14[degrees] 04'21.2" N, 16[degrees] 10'51.6" W. altitude 1.51 m a.s.l. 0-4 1.0 5.9 5.9 10 4-32 1.3 4.4 4.4 6 32-62 1.3 4.0 3.9 5 62-82 1.5 3.9 3.8 4 82-110 1.5 3.8 3.7 5.5 Middle terrace (P7), Endogleyic Arenosol, 14[degrees] 04'15.9" N, 16[degrees] 10'47.5" W, altitude 4.15 m a.s.l. 0-26 1.5 4.1 4.7 0.01 26-60 1.5 4.2 4.7 0.04 60-108 1.6 4.6 5.3 0.02 108-160 1.5 4.9 5.4 0.03 Middle terrace (P8), Endogleyic Arenosol, 14[degrees] 04'14.0" N. 16[degrees] 10'47.1" W, altitude 4.19 m a.s.l. 0-31 1.5 4.6 4.9 0.03 31-62 1.6 5.2 5.7 0.03 62-107 1.5 5.6 6.0 0.03 107-138 1.5 5.8 5.8 0.20 138-160 1.5 6.0 6.4 0.05 Middle terrace (P9), Endogleyic Arenosol, 14[degrees] 04'14.0" N, 16[degrees] l0'47.2" W, altitude 4.34 m a.s.l. 0-36 1.5 5.8 6.0 0.15 36-67 1.6 6.2 6.5 0.10 67-88 1.6 6.7 7.6 0.30 88-126 1.6 6.8 7.8 1.00 126-160 1.5 6.5 6.8 2.50 Depth [Fe.sub.t] [Fe.sub.d] [Fe.sub.o] [s.sub.t] (cm) (g [kg.sup.-l]) Floodplain (PI), Gleyic Hyposalic Solonchak (sulfatic), 14[degrees] 04'27.4" N, 16[degrees] 11'17.4" W 0-1 22.5 12.6 36 10 1-4 34.6 24.4 1.8 4 4-23 38.2 36.2 1.6 6 23-60 18.4 11.1 0.8 8 Floodplain (P2), Gleyic Hypersalic Solonchak (sulfatic), 14[degrees] 04'23.8" N, 16[degrees] 11'16.6" W 0-1 10.9 4.3 4.1 14 1-8 32.9 21.2 4.0 3 8-30 32.0 26.3 2.8 7 30-60 18.6 16.5 1.1 5 Floodplain (P3), Gleyic Hyposalic Solonchak (sulfatic), 14[degrees] 04'21.4" N, 16[degrees] 11'06.0" W, altitude 0.41 m a.s.l. 0-5 4.5 1.5 0.3 7 5-11 4.9 1.4 0.5 5 11-17 4.6 1.9 0.4 6 17-26 13.7 2.2 1.8 2 26-34 18.2 7.7 2.4 2 34-70 11.8 6.4 1.5 4 Low terrace (P4),Haplic Gleysol (thionic), 14[degrees] 04'19.4" N, 16[degrees] 10'59.2" W, altitude 1.15 m a.s.l. 0-2 15.5 7.1 4.6 2 2-16 19.9 8.5 1.3 2 16-38 18.9 8.0 0.7 0.3 38-62 23.1 7.5 0.3 0.3 62-90 16.8 5.3 0.2 0.3 90-100 19.0 5.9 0.2 1 Low terrace (P5),Haplic Gleysol (thionic), 14[degrees] 04'20.3" N, 16[degrees] 10'59.5" W, altitude 1.17 m a.s.l. 0-2 8.6 3.9 1.2 1 2-11 15.7 8.0 2.0 3 11-29 25.9 7.7 1.3 0.3 29-51 27.9 8.3 0.6 0.3 51-81 20.6 6.9 0.5 0.4 81-92 22.7 5.7 0.3 0.2 Low terrace (P6), Haplic Gleysol (thionic), 14[degrees] 04'21.2" N, 16[degrees] 10'51.6" W. altitude 1.51 m a.s.l. 0-4 5.6 3.2 0.4 2 4-32 10.2 4.0 1.1 4 32-62 11.7 4.2 0.6 8 62-82 13.8 6.3 0.5 5 82-110 21.7 6.5 0.8 0.2 Middle terrace (P7), Endogleyic Arenosol, 14[degrees] 04'15.9" N, 16[degrees] 10'47.5" W, altitude 4.15 m a.s.l. 0-26 6.6 4.6 0.9 0.03 26-60 7.6 3.7 0.7 0.03 60-108 2.3 1.7 0.01 0.02 108-160 4.0 3.5 0.1 0.01 Middle terrace (P8), Endogleyic Arenosol, 14[degrees] 04'14.0" N. 16[degrees] 10'47.1" W, altitude 4.19 m a.s.l. 0-31 6.2 2.5 0.3 0.02 31-62 3.8 1.9 0.1 0.01 62-107 2.3 1.0 0.05 0.01 107-138 29.3 3.5 0.4 0.02 138-160 3.0 1.4 0.1 0.01 Middle terrace (P9), Endogleyic Arenosol, 14[degrees] 04'14.0" N, 16[degrees] l0'47.2" W, altitude 4.34 m a.s.l. 0-36 10.5 1.9 0.3 0.3 36-67 10.8 1.8 0.2 0.03 67-88 10.4 2.0 0.2 0.03 88-126 9.6 1.6 0.3 0.04 126-160 178 0.9 0.6 0.05 Depth S[O.sub.4.sup.2-] Water soluble ions (cm) ([cmol.sub.c] [kg.sup.-1]) [Cl.sup.-] [Na.sup.+] Floodplain (PI), Gleyic Hyposalic Solonchak (sulfatic), 14[degrees] 04'27.4" N, 16[degrees] 11'17.4" W 0-1 1.4 14.5 10 1-4 0.5 5.6 4.7 4-23 0.3 5.7 5.2 23-60 0.4 8.2 7 Floodplain (P2), Gleyic Hypersalic Solonchak (sulfatic), 14[degrees] 04'23.8" N, 16[degrees] 11'16.6" W 0-1 2.6 37.7 33.2 1-8 0.4 4.3 3.8 8-30 0.3 4.5 4.2 30-60 0.4 7.8 6.5 Floodplain (P3), Gleyic Hyposalic Solonchak (sulfatic), 14[degrees] 04'21.4" N, 16[degrees] 11'06.0" W, altitude 0.41 m a.s.l. 0-5 1.4 5.5 4.2 5-11 0.7 2.5 2.4 11-17 0.6 2 1.4 17-26 0.2 2.3 2 26-34 0.2 3.8 3 34-70 0.5 7 7.3 Low terrace (P4),Haplic Gleysol (thionic), 14[degrees] 04'19.4" N, 16[degrees] 10'59.2" W, altitude 1.15 m a.s.l. 0-2 0.2 2.3 1.7 2-16 0.03 0.08 0.1 16-38 0.03 0.1 0.2 38-62 0.03 0.3 0.3 62-90 0.04 0.4 0.4 90-100 0.04 0.4 0.4 Low terrace (P5),Haplic Gleysol (thionic), 14[degrees] 04'20.3" N, 16[degrees] 10'59.5" W, altitude 1.17 m a.s.l. 0-2 0.2 5.5 2.8 2-11 0.5 3 1.2 11-29 0.04 2.4 1.1 29-51 0.03 1.6 0.8 51-81 0.03 1.1 0.7 81-92 0.04 1.5 0.7 Low terrace (P6), Haplic Gleysol (thionic), 14[degrees] 04'21.2" N, 16[degrees] 10'51.6" W. altitude 1.51 m a.s.l. 0-4 0.09 5.4 2.2 4-32 0.06 3 1.2 32-62 0.04 2.6 0.9 62-82 0.03 2.1 0.7 82-110 0.03 2.8 1.9 Middle terrace (P7), Endogleyic Arenosol, 14[degrees] 04'15.9" N, 16[degrees] 10'47.5" W, altitude 4.15 m a.s.l. 0-26 0.001 0.002 0.001 26-60 0.001 0.002 0.001 60-108 0.001 0.002 0.0002 108-160 0.001 0.002 0.0004 Middle terrace (P8), Endogleyic Arenosol, 14[degrees] 04'14.0" N. 16[degrees] 10'47.1" W, altitude 4.19 m a.s.l. 0-31 0.001 0.002 0.001 31-62 0.001 0.002 0.001 62-107 0.001 0.002 0.0004 107-138 0.002 0.004 0.007 138-160 0.002 0.002 0.001 Middle terrace (P9), Endogleyic Arenosol, 14[degrees] 04'14.0" N, 16[degrees] l0'47.2" W, altitude 4.34 m a.s.l. 0-36 0.001 0.003 0.001 36-67 0.001 0.002 0.002 67-88 0.002 0.004 0.002 88-126 0.002 0.03 0.002 126-160 0.003 0.09 0.002 Depth Water soluble ions (cm) ([cmol.sub.c] [kg.sup.-1]) [K.sup.+] [Ca.sup.2+] [Mg.sup.2+] Floodplain (PI), Gleyic Hyposalic Solonchak (sulfatic), 14[degrees] 04'27.4" N, 16[degrees] 11'17.4" W 0-1 0.2 1 2.4 1-4 0.1 0.3 0.9 4-23 0.1 0.2 0.9 23-60 0.1 0.2 1.2 Floodplain (P2), Gleyic Hypersalic Solonchak (sulfatic), 14[degrees] 04'23.8" N, 16[degrees] 11'16.6" W 0-1 0.3 2 5.1 1-8 0.2 0.3 0.9 8-30 0.1 0.2 0.7 30-60 0.1 0.2 1.9 Floodplain (P3), Gleyic Hyposalic Solonchak (sulfatic), 14[degrees] 04'21.4" N, 16[degrees] 11'06.0" W, altitude 0.41 m a.s.l. 0-5 0.08 1.4 1.5 5-11 0.06 0.6 0.7 11-17 0.05 0.6 0.4 17-26 0.06 0.3 0.5 26-34 0.07 0.4 0.7 34-70 0.2 0.4 1.5 Low terrace (P4),Haplic Gleysol (thionic), 14[degrees] 04'19.4" N, 16[degrees] 10'59.2" W, altitude 1.15 m a.s.l. 0-2 0.02 0.6 0.6 2-16 0.004 0.003 0.001 16-38 0.006 0.01 0.01 38-62 0.006 0.02 0.02 62-90 0.008 0.03 0.03 90-100 0.01 0.03 0.03 Low terrace (P5),Haplic Gleysol (thionic), 14[degrees] 04'20.3" N, 16[degrees] 10'59.5" W, altitude 1.17 m a.s.l. 0-2 0.1 0.9 2.2 2-11 0.07 1.8 1.2 11-29 0.06 0.3 0.9 29-51 0.04 0.2 0.6 51-81 0.03 0.2 0.3 81-92 0.03 0.1 0.2 Low terrace (P6), Haplic Gleysol (thionic), 14[degrees] 04'21.2" N, 16[degrees] 10'51.6" W. altitude 1.51 m a.s.l. 0-4 0.01 1.6 1.4 4-32 0.008 1 0.8 32-62 0.005 0.8 0.7 62-82 0.004 0.8 0.7 82-110 0.006 0.8 0.8 Middle terrace (P7), Endogleyic Arenosol, 14[degrees] 04'15.9" N, 16[degrees] 10'47.5" W, altitude 4.15 m a.s.l. 0-26 0.002 0.001 0.0004 26-60 0.001 0.001 0.0004 60-108 0.001 0.001 0.0001 108-160 0.001 0.001 0.0002 Middle terrace (P8), Endogleyic Arenosol, 14[degrees] 04'14.0" N. 16[degrees] 10'47.1" W, altitude 4.19 m a.s.l. 0-31 0.001 0.001 0.001 31-62 0.001 0.001 0.001 62-107 0.001 0.001 0.001 107-138 0.002 0.002 0.002 138-160 0.001 0.0005 0.0002 Middle terrace (P9), Endogleyic Arenosol, 14[degrees] 04'14.0" N, 16[degrees] l0'47.2" W, altitude 4.34 m a.s.l. 0-36 0.001 0.001 0.001 36-67 0.002 0.002 0.002 67-88 0.002 0.001 0.002 88-126 0.002 0.003 0.003 126-160 0.002 0.006 0.005 Table 2. Semi-quantitative estimation of clay mineral abundance Horizon Depth Kaolinite Smectite Illite (cm) (%) (%) (%) Floodplain, P1 Topsoil 0 4 75 23 3 Central horizons 4 23 67 32 1 Subsoil 23 60 71 28 1 Floodplain, P2 Topsoil 0-8 83 14 3 Central horizons 8-30 64 35 1 Subsoil 30-60 70 29 1 Low terrace, P4 Topsoil 0 16 81 17 2 Central horizons 16-62 94 5 I Subsoil 62-100 87 12 1 Low terrace, P5 Topsoil 0-11 85 13 2 Central horizons 11-51 91 8 1 Subsoil 51 100 67 32 1 Middle terrace, P7 Topsoil 0 60 82 17 1 Central horizons 60-108 72 26 1 Subsoil 108-160 75 23 1 Middle terrace, P8 Topsoil 0-62 41 58 1 Central horizons 62 107 47 52 1 Subsoil 107 160 58 40 2
|Printer friendly Cite/link Email Feedback|
|Author:||Fall, Aidara C.A. Lamine; Montoroi, Jean-Pierre; Stahr, Karl|
|Date:||Oct 1, 2014|
|Previous Article:||Fate of urine nitrogen through a volcanic vadose zone.|
|Next Article:||Soil fertility, physical and chemical organic matter fractions, natural [sup.13]C and [sup.15]N abundance in biogenic and physicogenic aggregates in...|