Soil evolution over the Quaternary period in a Mediterranean climate (SE Spain)
I. Ortiz, , M. Simón, C. Dorronsoro, F. Martín and I. García
Departamento Edafología y Química Agrícola, Facultad de Ciencias, Universidad de Granada, Campus de Fuentenueva, s/n 18071, Granada, Spain
2. Site details
3. Materials and methods
4.2. Analytical features
4.3. Micromorphological features
5.1. Buried soils
5.2. Surface soils
5.3. Soil–time relationships
According to Jenny (1941), soils and their properties are the product of the different soil-forming factors (climate, organisms, relief, parent material and time) that control the degree of soil development, as indicated by comparisons with the parent material (Harden, 1990). Because the soil-forming factors also govern geomorphic processes, landscape evolution is intimately related to soil development (McFadden and Kneupfer, 1990).
Over time, soil-forming factors, especially climate and vegetation, may change in such a way that many old soils, palaeosols, are not related to the present climate and vegetation. Palaeosols, defined as soils formed in a landscape of the past (Ruhe and Yaalon), include both relict and buried soils (Bronger and Catt, 1989). Relict soils are surface soils which show inactive characteristics inherited from past periods when soil-forming conditions were sufficiently different from those of the present to produce features unlike any of those developing currently in the same area. They are likely to have properties similar to those of buried soils that formed during the same past periods (Catt, 1989). It is possible to reconstruct relief and/or palaeoclimate as soil-forming factors on the basis of the processes inferred from palaeosol properties (Bronger and Catt, 1989).
Properties of the different soil horizons have also been used to determine the age of soils (Harden; Levine and Harrison) and thus the approximate age of the landforms (Semmel, 1989). For this reasoning to be valid, climatic conditions must remain relatively stable over the entire soil-forming period, for only then do the soil properties increase constantly with time (Bockheim and Birkeland). However, dating becomes complex on surfaces subject to long-term climatic fluctuations. This can be solved if relationships between specific soil properties and climatic fluctuations are known.
Among soils up to 85,000 years old (estimated from the clay accumulation index of Levine and Ciolkosz, 1983) in Sierra Nevada (SE Spain), Simón et al. (2000) distinguished two well-differentiated groups: (a) soils approximately 85,000 years old (early Late Pleistocene), with strongly developed Bt horizons, which are red in colour, clayey in texture and contain abundant clay coatings and much kaolinite; and (b) soils younger than 15,000 years (Late Pleistocene–Holocene), with less developed Bw horizons, which are brown in colour, without evidence of clay illuviation and with small kaolinite contents. The degree of development of the latter group was less in the younger surface. No soils of intermediate ages (between 85,000 and 15,000 years old) were found, apparently because of unstable surfaces and a cold climate during this time interval, which would have discouraged chemical weathering and soil development (Catt, 1989). In Sierra Nevada, SE Spain, some surfaces older than 85,000 years are preserved, probably formed during the Riss glacial period (Hempel; Messerli and Lhenaff), but their soils are strongly eroded and cannot be used for determining soil–time relationships. However, lower elevation alluvial fan deposits around Sierra Nevada that resulted from the tectonic activity in the Late Pliocene exhibit stable surfaces with soils of Early Pleistocene age (Estévez and Sanz de Galdeano, 1983). Also, the reorganization of the relief during the Middle–Late Pleistocene formed unstable surfaces on which successive depositional episodes alternated with pedogenic episodes (Sanz de Galdeano and López-Garrido, 1999).
In this paper, we compare soil development on two different types of surfaces in SE Spain: (1) geomorphically stable surfaces with old soils also showing younger pedogenic overprinting (surface soils) and (2) unstable surfaces with successive erosion–deposition episodes, forming sequences of buried soils in which the successive pedogenic stages are spatially distinct. The aim is to reconstruct soil development over the Quaternary period in a Mediterranean climate.
2. Site details
The Granada Basin is located in the central sector of the Betic Cordillera (SE Spain) in the contact area between the External and the Internal Zones (Fig. 1). The External Zones, located to the north of the Granada Basin, are made up of Mesozoic and Tertiary carbonate rocks (limestones and dolomites). The Internal Zones, located to the east of the Granada Basin, are made up of two complexes: the Nevado–Filabride Complex, occupying the central sector of Sierra Nevada and composed mainly of mica schists and quartzites; and the Alpujarride Complex, forming a ring around Sierra Nevada and composed of phyllites, quartzites, limestones and dolomites. During the Late Miocene, Sierra Nevada was strongly uplifted whereupon massive erosion gave rise to major alluvial fans containing large blocks reworked from the Nevado–Filabride Complex on the borders of the basin, and lacustrine formations were deposited in subsiding areas within the basin (Fernández and Soria, 1986–1987). In the Late Pliocene, there was renewed uplift (Estévez and Sanz de Galdeano, 1983), and significant new coarse detrital inputs from the External and the Internal Zones were deposited in the basin during the Early Pleistocene. Consequently, the northern half of the basement of the Granada Basin is made up of Mesozoic and Tertiary carbonate materials from the External Zones, and the eastern half consists of Paleozoic and Triassic materials from the metamorphic complexes of the Internal Zones (Fernández et al., 1996). Further reorganization of the relief occurred during the Middle?–Late Pleistocene. Low areas on the borders of Sierra Nevada rose and subsequent erosion episodes, probably activated by cold episodes, have left behind abundant coarse-grained deposits. In summary, the uplift of Sierra Nevada and the different sedimentation episodes in the Granada Basin were not continuous processes but resulted from pulses of tectonic activity separated by periods of relative quiescence (Sanz de Galdeano and López-Garrido, 1999).
The present climate of the area (Table 1) is typically Mediterranean (hot, dry summers; cold, wet winters; temperate autumns and springs with variable rainfall). The natural vegetation (Valle and Ruiz) is oak forest (Quercus rotundifolia) with shrubs (Juniperus oxycedrus, Ruscus aculeatus, Daphne gnidium, Clematis flammula, Lonicera etrusca and Hedera helix) and herbaceous plants (Paeonia coriacea, P. broteroi, Primula vulgaris and Viola sp.). However, in many sectors, this has been replaced by crops such as olive and almond trees.
3. Materials and methods
We have studied soils that developed on three alluvial fans in the Granada Basin that have remained relatively stable over time (Fig. 1). Two of these, Dúrcal (DUR) and Llano de la Perdiz (LLP), date from the Early Pleistocene (Aguirre; Ruiz and Sanz), and consist of gravels with mica schists and quartzites from the Nevado–Filabride Complex and a small proportion of limestones and dolomites from the Alpujarride Complex. The third alluvial fan, Colomera (COL), also dates from the Early Pleistocene (Fernández and Soria, 1986–1987) but has gravel clasts of limestones and dolomites from the External Zones.
Similarly, in a sector adjacent to the area that rose during the Middle?–Late Pleistocene, Nigüelas (NIG), we studied a vertical section approximately 11.5 m high where four depositional episodes of gravelly materials, equivalent to those from DUR and LLP (mica schists and quartzites from the Nevado–Filabride Complex with a small proportion of limestones and dolomites from the Alpujarride Complex), were distinguished. These depositional episodes alternated with pedogenic episodes. We identified and studied four buried soils and designated them (from bottom to top) as NIG-1, NIG-2, NIG-3 and NIG-4 (Fig. 2). The original surfaces of soils NIG-1 and NIG-2 were tilted before the deposition of the parent material of soil NIG-3.
Field descriptions of the soils were based on procedures of the Soil Survey Staff (1990). The micromorphological study was based on thin sections (Bullock et al., 1985). The Munsell soil colour chart was used to describe the soil colours. Particle size distribution was determined by the pipette method after the removal of organic matter with H2O2 and dispersion by shaking with sodium hexametaphosphate (Loveland and Whalley, 1991). The organic carbon content was determined using the method of Tyurin (1951). The pH was measured potentiometrically in a 1:2.5 soil/water suspension. The CaCO3 equivalent was determined according to Williams (1948). For the determination of the cation exchange capacity (CEC), 1 N Na–acetate was used at pH 8.2. Exchangeable bases were extracted with 1 N NH4–acetate at pH 7.0 and measured by atomic absorption spectroscopy (Ca and Mg) and flame photometry (Na and K). Discs of soil and lithium tetraborate (0.6:5.5) were prepared and the total contents of Si, Fe and Al were measured by X-ray fluorescence using a Philips PW-1404 instrument. X-ray diffraction patterns for the clay fraction were obtained with a Philips PW-1700 instrument using CuK radiation, and the diffraction intensities used in the quantitative analysis were taken from Schultz (1964) and Barahona (1974). Total iron oxides (Fed) were extracted with citrate–dithionite (Holmgren, 1967), and the amorphous forms (Feo) with ammonium oxalate (Schwertmann and Taylor, 1977). Iron in the extracts was measured by atomic absorption spectroscopy. A redness index (Rr) was calculated as (hueÅ~chroma)/value (Hurst, 1977). For this index, hue is converted to the following values: 10YR=0.0, 7.5YR=2.5, 5YR=5.0, 2.5YR=7.5 and 10R=10.0.
To estimate the degree of development of each profile, a clay accumulation index (CI) was calculated as (B-C)T, where B=B horizon clay content (%), C=C horizon clay content (%) and T=thickness (cm) of the B horizon (Levine and Ciolkosz, 1983). Similarly, an iron oxide accumulation index (FedI) was calculated using the same equation as for the clay, where B=B horizon Fed content (%) and C=C horizon Fed content (%). To estimate the age of the soils, the equation of Levine and Ciolkosz (1983) was used:
The C horizons of all the soils ranged in colour from pink to light brownish grey and retained the original deposit structure although in the surface soils, the mineral particles are cemented by calcium carbonate. None of the buried soils contains a clear A horizon, suggesting that they were disturbed or truncated at the time of burial. Surface soils also seemed to be truncated, especially soils COL (under olive cultivation) and DUR (under almond cultivation), where the Ap horizons are part of the previous Bt horizons that are disturbed by ploughing. Soil LLP was the least disturbed. All soils have a well-developed Bt horizon (Table 2) characterized by a red colour and a moderate to strong angular–subangular blocky or prismatic structure. The redness indices of the most strongly developed Bt horizons in each soil ranged from 7.5 to 15, being greatest in the surface soils COL, LLP and DUR and the buried soil NIG-3, intermediate in soils NIG-1 and NIG-2 and least in NIG-4.
4.2. Analytical features
All the Bt horizons contained more clay than the C horizons (Table 3). The clay accumulation index (CI) was greatest in soils COL, DUR, LLP and NIG-3, intermediate in soils NIG-1 and NIG-2, and least in soil NIG-4 (Fig. 3). The pH in the Bt horizons was mostly alkaline (Table 3) except for soil LLP, where slightly acid pH values appeared in the upper weakly calcareous horizons. In soil NIG-2, the Bt horizons seem to have been recalcified by the leaching of calcium carbonate from soil NIG-3. CaCO3 has accumulated in the Ck horizons, particularly in the surface soils where the mineral particles are cemented to form a Ckm horizon. For this reason, they could be designated as Bk or Bkm, but because most or all of the original parent material structure has not been obliterated (Soil Survey Staff, 1990), they are labelled as C horizons. The gravel content was similar in all C horizons (Table 3). The Bt horizons of all the soils contained very little organic C, indicating that the mineralization of organic matter predominated during the development of these soils.
The cation exchange capacity (CEC) was related to the clay and organic C contents by the multiple-regression equation:
CEC (cmolc kg-1)=5.008Å~OC (%)+0.345Å~ Clay (%) (r=0.965)
The regression coefficients show that the influence of organic C on the CEC values was roughly 15 times greater than that of the clay. Exchangeable bases are dominated mainly by Ca2+ and Mg2+, with lesser amounts of Na+ and K+. Only the Ap horizons of soils DUR and COL show relatively high contents of K+ which is attributable to fertilizing. All soils are eutric. The high base saturation (equal or close to 100%) can be attributed to basification by the runoff of waters rich in Ca2+ and Mg2+ originating from the surrounding terrain of limestone and dolomite. Only soil LLP, whose surface was isolated from the surrounding terrain by an incision of the rivers probably during the Middle?–Late Pleistocene (Sanz de Galdeano and López-Garrido, 1999), was less affected by this runoff and has a base saturation of less than 80% in its upper horizons. This basification must have occurred subsequent to soil formation and the original pH of the Bt horizons should have been more acidic than that at present. The neutral or nearly neutral pH values and high base saturation of these red soils have also been explained by the occurrence of dry periods during which there was a capillary rise of bases (Lamouroux, 1971). Whatever the mechanism causing basification, the red soils are usually less acidic than the brown soils that developed over the similar parent material (Duchaufour, 1977).
The contents of the total iron (Fet) that was extracted by dithionite (Fed) and that extracted by oxalate (Feo) were all greater in the Bt horizons than the C horizons (Table 4). The values of the iron oxide accumulation index (FedI) showed a similar pattern to those of the clay accumulation (Fig. 3) and redness indices (Fig. 4). It was greatest in the Bt horizons of soils COL, LLP, DUR and NIG-3, least in soil NIG-4 and intermediate in soils NIG-1 and NIG-2. The values of the Feo/Fed ratio in the Bt horizons were very small (<0.05), indicating an almost total crystallization of the hydrous Fe oxides that were formed by the weathering of silicates (Arduino et al., 1986).
In all the soils, the Fet/Sit and Alt/Sit ratios were greater in the Bt horizons than the C horizons (Table 4), indicating that Si was more mobile than Fe and Al. As the original pH of these soils should not have been <5.0 (Loughnan, 1969), the difference in the Fet+Alt/Sit ratio between the Bt horizons and C horizons should increase with greater weathering and leaching. These differences show approximately the same patterns as the CI, FedI and Rr indices (Fig. 5).
The semi-quantitative analysis of the clay minerals (Table 4) also revealed differences between the soils. In soils COL, LLP, DUR and NIG-3, smectite is less abundant and kaolinite was more abundant in the Bt horizons than in the C horizons. However, in soils NIG-1, NIG-2 and NIG-4, the upward increase in kaolinite was weaker than in other soils. The kaolinite neoformation in the Bt horizons must have occurred before basification when the soils were more acidic.
4.3. Micromorphological features
The Bt horizons show a porphyric-related distribution, with stipple-speckled b-fabric in soils NIG-1, NIG-2 and NIG-4 and mono-granostriated b-fabric in NIG-3 and the surface soils. In NIG-4, the Bt horizons show many thin red clay coatings in the channels and other voids (Fig. 6a). They are somewhat thicker and more common in NIG-1 and NIG-2 (Fig. 6b), and even thicker and more abundant in NIG-3 and in the surface soils (Fig. 6c). The surface soils contain many fragments of these red clay pedofeatures embedded in the matrix (Fig. 6d). The red clay coatings in soil LLP have scattered yellowish zones (Fig. 6e), indicating weak hydromorphic iron depletion. Distinct and rather frequent calcitic coatings appear in the Bt horizon of NIG-2 (Fig. 6f). Similar features also appear, though with less clarity and less frequency, in the surface soils.
The main pedogenic processes that affected these soils were the mineralization of organic matter, leaching of carbonates, strong weathering of smectite to kaolinite, clay illuviation and rubification, which formed strongly developed red Bt horizons of clay texture with abundant clay coatings. These soil properties must have developed under a wetter climate than that at present. In addition, the pH>7.0, the high contents of exchangeable bases and the presence of CaCO3 in the Bt horizons suggest subsequent calcification, the latter also being evident in the micromorphological study of soils NIG-2, DUR and COL.
5.1. Buried soils
Continuous deep oceanic sedimentary records can be used as a chronological and paleoclimatic reference for long-term climatic fluctuations (Kukla and Bradley). Because the oxygen isotopic record of the oceanic sequences provides an integrated summary of global ice-volume changes, it has been argued that the isotopic stages should be used as standard reference units for both marine and terrestrial deposits (Shackleton and Opdyke, 1973). Various authors have used this marine record to date and correlate the episodes of soil development (Bronger; Bronger; Bronger; Markewich; Stremme; Olsen; Frechen; Dearing and Antoine).
According to the ages of the isotopic events in the low-latitude oxygen-isotope sequence (Bassinot et al., 1994), the deposit on which the heavily eroded soil NIG-5 developed probably formed during the last cold episodes between 11,000 and 71,000 BP (stages 4–2). Consequently, the parent material of soil NIG-4 should have formed in the former cold episode, between 127,000 and 186,000 BP (stage 6), and soil NIG-4 during the warm periods between 71,000 and 127,000 BP (stage 5). In addition, the formation of soil NIG-4, estimated by the clay accumulation index (Levine and Ciolkosz, 1983), must have begun around 85,000 BP or even earlier, given that erosion decreased the thickness of the Bt horizons. This supports the suggestion that this soil was formed during stage 5. The deposit on which soil NIG-3 developed probably formed during the cold episode between 242,000 and 301,000 BP (stage 8) and soil NIG-3 during the warm episode between 186,000 and 242,000 BP (stage 7). The Bt horizons of soil NIG-2 probably formed during the warm period between 301,000 and 334,000 BP (stage 9) and its parent material dates from the cold episode between 334,000 and 364,000 BP (stage 10). Finally, soil NIG-1 probably formed during the warm period between 364,000 and 427,000 BP (stage 11) and its parent material was probably deposited during the cold episode between 427,000 and 474,000 BP (stage 12). Consequently, the tilting of both deposits and soils NIG-2 and NIG-1, which were related to an uplift of Sierra Nevada, must have occurred around 300,000 BP in the Middle Pleistocene.
Based on the CI and FedI indices (Fig. 3), the differences in the Fet+Alt/Sit ratio between Bt and C horizons (Fig. 5), the extent of kaolinite neoformation (Table 4) and the micromorphological features, soils NIG-2 and NIG-1 show similar degrees of development although less than that of soil NIG-3 and greater than that of NIG-4. In addition, the duration of the warm periods in which these soils developed was around 63,000 years (NIG-1), 33,000 years (NIG-2) and 56,000 years (NIG-3 and NIG-4). Therefore, the time factor appears not to account for the different degrees of development of the buried soils, especially NIG-1, NIG-3 and NIG-4. The greater development of soil NIG-3 may therefore be attributed to a different, probably moister, climate. Greater moisture would also account for the leaching of carbonates from the Bt horizons of NIG-3 through the C horizon to form calcitic coatings in the Bt horizons of soil NIG-2. Consequently, in our region, the different degrees of soil development during the last 474,000 BP indicate that the wettest climate of the later Quaternary warm periods dates from between 186,000 and 242,000 BP (stage 7), and the driest from 71,000–127,000 BP (stage 5). The warm periods older than 242,000 BP (stages 9 and 11) probably had climates with intermediate wetness.
5.2. Surface soils
The parent materials of the surface soils, dating from the Early Pleistocene (between 788,000 and 1,650,000 BP; Birkeland, 1999), must have been deposited during one of the cold episodes before 788,000 BP, and the soils on them were formed during subsequent warm periods. The CI and FedI indices (Fig. 3), the differences in the Fet+Alt/Sit ratio between Bt and C horizons (Fig. 5), the extent of kaolinite neoformation (Table 4) and the micromorphological features were similar in all of the soils, indicating an equivalent degree of weathering and development. The minor differences in the indices of these soils could be attributed to parent-material differences or to waterlogging in some profiles. The greater weatherability of carbonate materials (limestones mainly) compared with metamorphic materials (mica schists and quartzites) could account for the slightly stronger development of soil COL, and the hydromorphic processes that affected soil LLP could explain its slightly weaker development.
The extent of development of the surface soils is similar to that of NIG-3, but the FedI index of the latter is slightly less (Fig. 3). Nevertheless, the surface soils present two basic differences from soil NIG-3. First, most of the red clay coatings are fragmented and incorporated into the soil matrix, and second, a strong accumulation of CaCO3 in the C horizons cements the mineral particles, forming the Ckm horizon. The fragmentation of the clay coatings suggests frost disturbance (Catt, 1987) and may be attributed to the cold episodes (Kemp and Van) following the formation of the Bt horizons. The origin of the large carbonate contents of the Ckm horizons of soils LLP and DUR, which were formed on a parent material similar to NIG-3 and also have decalcified Bt horizons, cannot be explained purely by leaching from the upper horizons, rather, this carbonate content must be attributed to the infiltration by the runoff water that is rich in Ca2+ and HCO3- ions. The presence of calcitic coatings in the Bt horizons of the surface soils indicates recalcification after the formation of the Bt horizons. This implies that CaCO3 accumulation and cementation in the Ckm horizons increased over time as described in the soils of fluvial terraces (Dorronsoro and Alonso, 1994). However, the accumulation of CaCO3 in the Ckm horizons was far greater in COL because this soil was surrounded by and formed over carbonate materials. Consequently, in periods after their formation, these soils were partially truncated, disturbed and recalcified to form polygenetic soils (Tarnocai and Valentine, 1989).
Because the extent of development of the surface soils is similar to that of NIG-3, we cannot rule out that it occurred during stage 7. Nevertheless, it could also have taken place during earlier warm episodes with similar climatic conditions to stage 7, such as stages 13 (between 474,000 and 528,000 BP) and 15 (between 568,000 and 621,000 BP). Bt horizons with clay illuviation are known to have formed elsewhere in these early interglacials (Bronger et al., 1998a).
5.3. Soil–time relationships
As the CI and FedI indices, the differences in Fet+Alt/Sit ratio between Bt and C horizons, the extent of kaolinite neoformation and the micromorphological features of the soils formed during stage 7 (186,000–242,000 BP) are similar to surface soils formed on deposits of the Early Pleistocene, these features cannot be used to date surfaces older than 242,000 BP. In contrast, the degree of development of the soils formed in stage 5 and later is less than that in stage 7 and decreases progressively towards the youngest surfaces (Simón et al., 2000), showing a clear relationship between the degree of development and the age of the surfaces on which they formed. Consequently, these soils can be used for the approximate dating of landforms.
The depositional and soil development episodes during the Pleistocene were not continuous but were governed by pulses of tectonic uplift, giving rise to sedimentation, separated by periods of relative quiescence with soil development. From the Early to the early Late Pleistocene, the main pedogenic processes were the leaching of carbonates, weathering, illuviation and rubification, but the degree of development of the Bt horizons varied over time. The surface soils that formed over the deposits from the Early Pleistocene show the strongest development although in periods after their formation, they were partially truncated, disturbed and recalcified, resulting in polygenetic soils. The different degrees of development of the buried soils during the last 474,000 years indicate that the wettest warm period was stage 7 (186,000–242,000 BP), and the driest, stage 5 (71,000–127,000 BP). Stages 9 (301,000–334,000 BP) and 11 (364,000–427,000 BP) had climates with intermediate wetness. Given that the CI and FedI indices, the differences in Fet+Alt/Sit ratio between Bt and C horizons, the extent of kaolinite neoformation and the micromorphological features of the soils that were formed during stage 7 are similar to the surface soils that were formed on deposits of the Early Pleistocene, these features cannot be used to date surfaces older than 242,000 BP. However, from stage 7, the degree of soil development progressively declines with the decreasing age of the surfaces so that these soils can be used to estimate the age of landforms.
This study was supported by DGICYT Project No. PB96-1385.
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