The influence of weathering processes on labile and stable organic matter in Mediterranean volcanic soils

Available online at www.sciencedirect.com

Geoderma 143 (2008) 191 – 205 www.elsevier.com/locate/geoderma

The influence of weathering processes on labile and stable organic matter in Mediterranean volcanic soils Vito Barbera a , Salvatore Raimondi a , Markus Egli b,⁎, Michael Plötze c a

Dipartimento di Agronomia Ambientale e Territoriale, University of Palermo, Viale delle Scienze 12, 90128 Palermo, Italy b Department of Geography, University of Zurich, Winterthurerstrasse 190, 8057 Zurich, Switzerland c ETH Zurich, Institute for Geotechnical Engineering, 8093 Zurich, Switzerland Received 8 January 2007; received in revised form 31 October 2007; accepted 2 November 2007 Available online 3 December 2007

Abstract The relationship and mechanisms among weathering processes, cation fluxes, clay mineralogy, organic matter composition and stability were studied in soils developing on basaltic material in southern Italy (Sicily). The soils were transitions between Phaeozems and Vertisols. Intense losses of the elements Na, Ca and Mg were measured indicating that weathering has occurred over a long period of time. The main weathering processes followed the sequence: amphibole, mica, volcanic glass or if ash was the primary source → smectite → interstratified smectite–kaolinite → kaolinite. Kaolinite formation was strongly related to high Al, Mg and Na losses. The good correlation between oxyhydroxides and kaolinite in the soils suggests that (macro)aggregates have formed due to physical or electrostatic interactions between the 1:1 clay minerals and oxides. The stability of organic matter was investigated with a H2O2-treatment that assumes that chemical oxidation mimics the natural oxidative processes. The ratio of C after the H2O2 treatment to the total organic C ranged from 1–28%. No correlation between clay content and organic matter (labile or stable fraction) was found. The refractory organic fraction was enriched in aliphatic compounds and did not greatly interact with the kaolinite, smectite or poorly crystalline Fe or Al phases. A part of this fraction (most probably proteins) was bound to crystalline Fe-oxides. In contrast, the oxidisable fraction showed a strong relationship with poorly crystalline oxyhydroxides and kaolinite. Surprisingly, smectite did not contribute to the stabilisation of any of the organic C fractions. The stabilisation of organic matter in the soils has, therefore, two main mechanisms: 1) the protection of labile (oxidisable with H2O2) organic matter, including also aromatic-rich compounds such as charcoal, by the formation of aggregates with oxyhydroxides and kaolinite and 2) the formation of a refractory fraction enriched in aliphatic compounds. © 2007 Elsevier B.V. All rights reserved. Keywords: Weathering; Mass balance; Clay mineralogy; Organic matter stability; Mediterranean soils; FT-IR spectroscopy

1. Introduction The Mediterranean climate is characterised by a strong seasonal winter/summer rainfall contrast with the result that the soils and the root zone are dry during the summer, often for several months. It borders the (cool) temperate region to the North and the subtropical desert region to the South. Similar environmental conditions prevail in all other continents, but to a lesser extent (Yaalon, 1997). Due to the dry summers, bush fires sporadically occur in many regions and exert a significant effect on landscape evolution (Naveh, 1990). ⁎ Corresponding author. E-mail address: [email protected] (M. Egli). 0016-7061/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.geoderma.2007.11.002

Regarding soil development and weathering, the excess of rainfall during several of the winter months enables moderately hydrolytic weathering of silicate minerals and the formation of 1:1 and/or 2:1 clay minerals. Soil hydrology and the lateral movement of soil moisture in through flow and surface flow are both the engine for clay illuviation and carbonate fluxes forming catenary soil changes on slopes (Yaalon, 1997). When layered clays (2:l) accumulate on lower slopes and flood plains, the seasonal swelling and shrinking rapidly produces the typical slickenside structure in the B horizon and deep surface cracking during summers, both characteristic of vertic horizons (Yaalon, 1997). Another important aspect regarding soil formation in this area is the anthropogenic factor. The long continuous settlement and intense cultivation by man, extending to more than

192

V. Barbera et al. / Geoderma 143 (2008) 191–205

5000 years, had important implications on the evolution of the landscapes (Brice, 1978; Brueckner, 1986). The climax vegetation is now an open forest or “interrupted” woodland and the vegetation biomass is about 4 to 5 kg C/m2, an intermediate value between closed forests and grasslands (Olson et al., 1983). Mean soil organic carbon stocks in the different life zones vary from 4 to over 10 kg C/m2 (Post et al., 1982). Soil organic matter is an important factor of the global carbon cycle. Depending on the turnover rates in the soil, three conceptual C fractions can be distinguished: active-labile and active-intermediate fractions – which may remain in soil for years or decades – and passive or refractory organic matter (OM) remaining in soil from centuries to millennia. Neither its pool size, important for C-turnover models, nor the stabilisation process responsible for its refractory is well known (Lützow et al., 2006). Small variations of the proportion of the C-forms in soils may have a significant effect on the global C balance and therefore on climate change (González-Pérez et al., 2004). Studies about OM stability in soil have shown that many factors like climate, soil matrix, human impact etc. influence humification and decomposition processes. Soil organic matter content is often positively correlated with the clay content of the soil (Bosatta and Ågren, 1997). Sorption processes play an important role in soil organic carbon (SOM) preservation (Wiseman and Püttmann, 2006) and consequently influence also the mean residence time of SOM (Saggar et al., 1996). Metal oxides have been demonstrated to be particularly effective in adsorbing and stabilising organic matter in soils (Kaiser and Zech, 1999; Wiseman and Püttmann, 2006). In soils developing from volcanic parent materials, secondary Al and Fe are hypothesised to stabilise SOM giving rise to its accumulation and to the typical black A horizon in Andosols (Shoji and Fujiwara, 1984; Rasmussen et al., 2005). Clay minerals or phyllosilicates can also help to preserve SOM (Wiseman and Püttmann, 2006). There is no unequivocal agreement about their role in SOM preservation. Stability of organic matter and its resistance to oxidation can be due to the capacity of soil matrix to preserve organic matter (Theng et al., 1986; Righi et al., 1995), but also to some natural compound recalcitrance because of their biochemical characteristics, like aliphatic macromolecules (lipids, cutans, algaenans, suberans), charcoal, sporopollenins and lignins. Only a little data on carbon and organic matter dynamics is available for volcanic Mediterranean areas. Extreme environmental conditions alternate through the year. Under such conditions, the relative importance of abiotic constraints such as fire (natural or man inducted) and irreversible dehydration favoured by intense solar radiation and drastic drying cycles, may be important factors in the formation of stable OM in soil. The introduction of fire-generated, highly condensed and resilient materials into soils is assumed to increase the refractory soil OM pool (Knicker and Skjemstad, 2000). Our aim was to decipher the process involved in OM stabilisation in selected Mediterranean soils by the analysis of the relationship between soil inorganic materials (oxyhydroxides, clay minerals), weathering mechanisms and SOM fractions (including the analysis of functional groups).

2. Investigation area and sites We selected six soils (Fig. 1, Table 1) which developed on basaltic substrates in the Monti Iblei region (Iblean Plateau, south-eastern part of Sicily, Italy). The chosen soils differed in the vegetation coverage, morphology and land use. Two profiles were close to each other and had a different vegetation (maquis and oak forest). Three other profiles developed on a gentle slope and were used as pasture (M. Lauro 1–3), however not intensively. The sixth profile M. Lauro 4 was in the vicinity of M. Lauro 1–3 and was used more intensively as pasture. The region is a carbonate plateau with substantial intercalations of volcanic material. Volcanism occurred from the Late Triassic through the early Pleistocene. Chemically, the products of Iblean volcanism are both (subalkalic) tholeiitic and alkalic basalts. Volcanic deposits in the Monti Iblei show a great variety of submarine facies while subaerial deposits are almost entirely sheeted lava flows with rare occurrences of partially welded spatter. The diversity of submarine deposits is due to different eruptive and depositional processes both during and after eruptive events (Lentini et al., 1990). According to the soil map of Sicily (Fierotti et al., 1988) the area is characterised by Regosols to (Andic) Cambisols. This association is found prevalently on volcanic material in the Iblean zone from the top of Monte Lauro (986 m a.s.l.) up to the sea shore (in north-eastern direction). It covers an area of approximately 29,000 ha. Its morphology is variable with undulating and hilly areas, gently slopes and steeper areas at high altitudes (near the top of Monte Lauro). In the more rugged areas, the prevalent agronomic use is pasture land which changes gradually to herbaceous cultures and more shrubs and trees on the gentler slopes. The vegetation type is, therefore, a mixture of open forests of deciduous oak, garrigue, grassland and rock plants. 3. Materials and methods 3.1. Soil sampling Soil profile ditches were made down to the C-horizon. In total, 6 sites were investigated. As soil horizon is, more or less, an identical compartment with typical chemical and mineralogical processes, sampling was bound to the morphology of the soils. From 2 to 3 kg soil material was collected per soil horizon. Soil bulk density was determined with a soil core sampler. Taking advantage of the profile pits, undisturbed soil samples were taken down to the C-horizon. 3.2. Soil mineralogy and grain sizes The clay fraction (b 2 μm) was obtained from the soil after destruction of organic matter with dilute and Na-acetate buffered H2O2 (pH 5) by dispersion with Calgon and sedimentation in water. Oriented specimens on glass slides were analysed by X-ray diffraction using CuKα radiation from 2 to 15°2θ with steps of 0.02°2θ at 2 s/step (PHILIPS PW1820; CuKα, 40 kV, 30 mA). The following treatments were performed: Mg saturation, ethylene glycol solvation (EG) and K-saturation, followed by heating

V. Barbera et al. / Geoderma 143 (2008) 191–205

193

Fig. 1. Investigation area and sites.

for 2 h at 335° and 550 °C. Digitised X-ray data were smoothed and corrected for Lorentz and polarisation factors (Moore and Reynolds, 1997). Peak separation and profile analysis were carried out by the Origin PFM™ using the Pearson VII algorithm after smoothing the diffraction patterns by a Fourier transform function. Background values were calculated by means of a nonlinear function (polynomial 2nd order function; Lanson, 1997). The semi-quantitative estimation of phyllosilicate concentration was done with two different methods: 1) Peak-area approach: This procedure allows a semi-quantitative estimation of phyllosilicate composition by the combination of the areas of the ethylene glycol solvated, Mg-saturated, the K-saturated and heated (335 °C and 550 °C) samples. On the basis of these integrals, an estimate of clay minerals composition was performed. The sum of the areas between 2

and 15° 2θ, which were attributed to HIV (hydroxylinterlayered vermiculites), smectite, vermiculite, mica, chlorite and kaolinite, were standardised to 100%. For the Mgsaturated and for the ethylene glycol solvation treatment, the area of the following peaks (d-spacings) were corrected by a weighting factor F: 1.6 nm with F = 0.453, 1.4 nm with F = 0.478, and 0.71 nm with F = 0.16 (Schwertmann and Niederbudde, 1993; Gjems, 1967; Laves and Jahn, 1972; Niederbudde and Kussmaul, 1978). For the determination of the kaolinite peak intensity, the possible contribution of chlorite to this signal was taken into account. Although the (semi)quantification of clay minerals in soils is bedevilled by manifold problems (Kahle et al., 2002), the applied and standardised (sample preparation, treatments, measurement and calculation) procedure enabled the assessment of the variability of clay mineral assemblage amongst the sites.

Table 1 Characteristics of the study sites Location

Elevation Slope Parent material Land use Vegetation (m a.s.l.) (%)

Soil type (soil taxonomy) Soil type (WRB a)

Buccheri Radura Buccheri Querceto Monte Lauro 1 Monte Lauro 2 Monte Lauro 3 Monte Lauro 4

890 900 909 904 897 850

Ultic Haploxerrol Lithic Haploxerrol Lithic Dystroxerept Vertic Dystroxerept Typic Hapludert Typic Hapludert

a

45 32 10 7 6 5

WRB = World Reference Base.

Alkali-basalt Alkali-basalt Alkali–basalt Alkali–basalt Alkali–basalt Alkali–basalt

Pasture Forest Pasture Pasture Pasture Pasture

Pyrus amigdaliformis, Crataegus monogina Quercus suber, Quercus pubescens Hedysarum coronarium, Hordeum vulgare Hedysarum coronarium, Hordeum vulgare Hedysarum coronarium, Hordeum vulgare Hedysarum coronarium, Hordeum vulgare

Haplic Phaeozem Leptic Phaeozem Leptic Umbrisol Vertic Cambisol Grumic Vertisol Thermic Vertisol

194

V. Barbera et al. / Geoderma 143 (2008) 191–205

2) The obtained results were cross-checked with the CEC (cation exchange capacity) approach. Most soils were dominated by kaolinite and smectite. If one assumes that a smectite layer contributes 100 and a kaolinite layer 10 parts to the total CEC, then the relative percentages of these minerals can be estimated.

analysed for C and N concentration. Functional groups and compounds were determined by FT-IR measurement (Table 2). FT-IR spectra were recorded over the range of 4000 to 250 cm− 1 on pellets made with 1 mg of sample and 250 mg of KBr previously heated at 150 °C. The pellets were heated at 60 °C prior to the measurements (to reduce the influence of water).

The main mineralogical composition of the clay fraction was checked using XRD patterns of randomly orientated specimens (2−30° 2θ, 0.02°2θ steps, 3 s/step, automatic slits). ‘Imogolite-type material’ (ITM), henceforth referred to the sum of imogolite and proto-imogolite allophane, was estimated (assuming that the Al/Si molar ratio is close to 2.0) according to Parfitt and Hemni (1982). The presence of ITM and kaolinite was checked with FT-IR measurements. ITM was further investigated by the molar ratio (Alo − Alp)/Sio with Alo as the oxalateextractable Al, Alp as the pyrophosphate-extractable Al and Sio as the oxalate-extractable Si. IR-Spectra were recorded over the range of 4000 to 250 cm− 1 on pellets made with 1 mg of sample and 250 mg of KBr that had been previously heated at 150 °C. After a pre-treatment of the samples with H2O2 (3%), particle size distribution of the soils was measured by a combined method consisting of sieving the coarser particles (2000–32 μm) and the measurement of the finer particles (b32 μm) by means of an X-ray-sedimentometer (SediGraph 5100).

3.5. Calculation of weathering rates

3.3. Soil chemistry Element pools in the fine earth and skeleton fraction (Ca, Mg, K, Na, Fe, Al, Mn, Si, Ti and Zr) were determined applying a method of total disintegration. Oven-dried (70 °C) samples were dissolved using a mixture of HF, HCl, HNO3 and H3BO3 as in Hossner (1996) and in a closed system (microwave-oven and at 25 bar). The concentrations of Ca, Mg, K, Na, Fe, Mn, Al, Si, Ti were determined by atomic absorption spectroscopy. Additionally, the dithionite-, pyrophosphate- and oxalate-extractable fractions were measured for the elements Fe, Al and Si (McKeague et al., 1971). The pyrophosphate extraction is often used to characterise the organically-bound Al fraction. Total C and N contents of the soil were measured with a C/H/N analyser (Elementar Vario EL). Soil pH (in 0.01 M CaCl2) was determined on air-dried fine earth samples using a soil:solution ratio of 1:2.5. Organic C was calculated as difference between total C and CaCO3 values. Carbonates were determined by dissolution with HCl. 3.4. SOM fractions Acting on the assumption that chemical oxidation mimics the natural oxidative processes, we treated the soils with H2O2 to eliminate labile organic material from more refractory organic matter. Wet oxidation was performed according to Eusterhues et al. (2005), by suspending 2 g of soils in 20 ml of 30% H2O2, at room temperature. Fresh H2O2 was added daily or twice a day, after centrifugation and removing of the old solution, if frothing was not visible. After 65 days when no reaction occurred after adding fresh H2O2, samples were washed twice with deionized water and oven dried at 60 °C for 24 h. The dried samples were

Long-term weathering rates of soils were derived from the calculations of enrichment/depletion factors determined using immobile element contents (e.g. Ti). The derivation of massbalance equations and their applications to pedologic processes are discussed in detail by Brimhall and Dietrich (1987) and Chadwick et al. (1990), and revised by Egli and Fitze (2000). Volumetric changes that occur during pedogenesis were determined by adopting the classical definition of strain, εi,w (Brimhall and Dietrich, 1987): ei;w ¼

Δzw 1 Δz

ð1Þ

with Δz as the columnar height (m) of a representative elementary volume of protore ‘p’ (or unweathered parent material) and Δzw is the weathered equivalent height (m) ‘w’. The calculation of the open-system mass transport function τj,w is defined by (Chadwick et al., 1990) !  qw Cj;w
sj;w ¼ ei;w þ 1  1 ð2Þ qp Cj;p with Cj,p (kg/t) as the concentration of element j in protolith (e.g. unweathered parent material, bedrock), Cj,w as the concentration of element j in the weathered product (kg/t), and with ρp and ρw being the bulk density (t/m3) of the protolith and the weathered soil, respectively. With n soil layers the calculation of changes in mass of element j is given by (Egli and Fitze, 2000) P

mj; flux ðzw Þ ¼

n X

 Cj;p qp

a¼1

 1 sj;w Δzw ei;w þ 1

ð3Þ

where τj,w corresponds to the mass transport function, εi,w to the strain, Cj,p (kg/t) to the concentration of element j in protolith Table 2 Major IR absorption bands and assignments (Stevenson, 1994; Senesi et al., 2003; Tan, 2003) Wave number cm− 1

Assignment

1725–1710 1660–1630

C_O stretching of COOH, aldehydes and ketones C_O stretching of amide groups, quinone C_O and/or C_O of H-bonded conjugated ketones Aromatic C_C, strongly H-bonded C_O of conjugated ketones Aromatic rings Aliphatic C–H stretching OH deformation and C–O stretching of phenolic groups Aliphatic, alcoholic OH C–O stretching of polysaccharide

1620–1600 1510–1535 1460–1440/2976–2937 1413–1333 1190–1127 1180–1050

V. Barbera et al. / Geoderma 143 (2008) 191–205

(e.g. unweathered parent material, bedrock), ρp being the bulk density (t/m3) of the protolith and Δz the weathered equivalent of the columnar height (m) of a representative elementary volume.

195

Table 4 Chemical characteristics of the investigated soils Site/horizons pH Total acidity CaCO3 BS EB (CaCl2) Na+ K+

CEC Ca++ Mg++

cmolc kg− 1 g kg− 1 % cmolc kg− 1

4. Results

Buccheri Radura A 5.6 Bw 5.8

44 52

5.0 3.8

61 0.26 0.21 22.12 9.65 52.9 68 0.66 0.13 27.57 17.56 67.6

Buccheri Querceto A 5.9

52

2.5

56 0.19 0.16 18.06 9.81 50.1

Monte Lauro 1 A 6.4

36

1.3

46 0.24 0.09 14.84 7.20 48.2

Monte Lauro 2 A 6.6 Bw 6.7 C 6.9

40 44 32

0.0 0.0 3.8

52 0.46 0.13 16.77 8.70 49.7 56 0.53 0.08 15.60 10.93 48.1 66 0.85 0.08 26.98 21.49 75.4

Monte Lauro 3 A 5.1 Bss1 5.8 Bss2 6.5 Bss3 6.7 Bss4 6.3 Bss5 6.4

48 44 36 40 40 28

2.5 2.5 1.3 1.3 2.5 3.8

62 59 58 64 63 64

0.78 0.78 0.56 1.16 1.24 1.75

0.22 0.11 0.09 0.13 0.11 0.14

26.05 25.37 24.05 25.62 25.14 24.48

8.51 9.58 9.55 10.01 9.83 8.46

57.7 60.8 59.3 57.9 57.6 54.7

Bulk Skeleton % Clay Silt Sand Munsell density g kg− 1 g kg− 1 g kg− 1 color g cm− 3 (moist)

Radura A (0–25) Bw (25–65) R N 65

Monte Lauro 4 A 7.4 Bss1 7.3 Bss2 7.4 Bss3 7.4

40 32 28 36

18.8 13.8 6.3 12.5

87 92 73 80

0.12 0.45 0.26 0.34

0.10 0.07 0.09 0.11

45.73 52.31 46.11 52.50

5.20 5.93 6.72 8.11

58.8 63.6 72.8 76.4

1.30 1.40 1.80

25 60

321 401

324 256

355 343

10 YR 3/2 2.5 Y 4/3

EB = exchangeable base cations, CEC = cation exchange capacity, BS = base saturation.

Querceto A (0–30/50) R N 50

1.43 1.80

10

296

322

382

10 YR 4/3

M. Lauro 1 A (0–35) R N 35

1.17 1.80

25

460

282

258

10 YR 3/1

M. Lauro 2 A (0–35) B (35–65) C (65–85) R N 85

1.06 1.24 1.11 1.80

5 30 60

422 407 509

277 279 215

301 314 276

10 Y/R 4/2 10 YR 4/4 10 YR 4/2

M. Lauro 3 A (0–25) Bss1 (25–45) Bss2 (45–70) Bss3 (70–100) Bss4 (100–120) Bss5 (120–140) R N 140

0.99 1.36 1.21 1.20 1.30 1.30 1.80

5 2 5 2 2 2

443 471 464 450 487 523

263 266 270 275 266 257

294 263 266 275 247 220

10 YR 10 YR 10 YR 10 YR 10 YR 10 YR

4/1 4/1 4/1 4/1 5/1 4/1

M. Lauro 4 A (0–20) Bss1 (20–50) Bss2 (50–90) Bss3 (90–110) R N 110

1.04 1.13 1.01 0.94 2.00

0 5 5 5

389 406 387 407

285 276 257 255

326 318 356 338

10 YR 10 YR 10 YR 10 YR

2/1 2/1 3/1 1/1

4.1. Soil physical and chemical aspects The particle size distribution (Table 3) of the selected profiles shows a high percentage of clay in all samples, ranging from 296 g kg− 1 up to 523 g kg− 1. The clay content along the soil profile does not vary greatly. There might be a tendency of higher contents with increasing soil depth (illuviation). A relatively high skeleton proportion was measured in the Bw horizon of the sites Radura and M. Lauro 2. In the soil profiles of M. Lauro 3 and 4 soil skeleton was almost absent. Except in the soil profile M. Lauro 4, the calcium carbonate content was close to the detection limit (Table 4). Most soils are weakly acidic with a pH between 5.1 and 6.9. Base saturation ranges between 47 to 68% for weakly acid soils and between 73

Table 3 Physical characteristics of the investigated soils Site/Horizon (cm)

and 92% in the soil profile M. Lauro 4. The CEC-values vary within a relatively small range of 46.1 to 76.4 cmolc kg− 1. The exchangeable cations indicate an increase with depth of Mg2+ and Ca2+ at the sites Radura, M. Lauro 2 and M. Lauro 4. K+ usually shows the highest values in the topsoil. The differences between M. Lauro 4 and the other soils might be due to the topographic position of this soil (in a flood plain) that is subjected to receive material from adjacent soils. The profiles Radura and Querceto are close to each other. They have, however, a different vegetation giving rise to some differences in soil chemistry (e.g. differences in exchangeable bases, base saturation and CEC). The sites M. Lauro 1, 2 and 3 represent a kind of a toposequence, with the last one in a flat position. Due to the lateral movement of water, soil erosion may have occurred leading to a change in soil thickness, clay illuviation and cations fluxes on the slope. There is a tendency of higher clay contents, more exchangeable bases, a higher base saturation and CEC from the sites M. Lauro 1 to 3 (A horizon). M. Lauro 3 and 4 are typical Vertisols. They have a different position in the landscape resulting in a higher accumulation of bases and CaCO3 (M. Lauro 4; Table 4). The main geochemical data of the bulk material is given in Table 5. The composition of the substrates is relatively constant among all sites. The SiO2 content is in the range of 40 to 50%, reflecting the basaltic character of the material. Compared to the

196

V. Barbera et al. / Geoderma 143 (2008) 191–205

Table 5 Geochemical characteristics (total analysis of the bulk material including soil skeleton and fine earth) of the investigated soils Site/horizon

Al2O3

SiO2

TiO2

CaO

MgO

K2 O

Na2O

Fe2O3

MnO2

LOI

g kg− 1 Buccheri Radura A Bw R

147.2 170.2 147.1

430.1 429.3 469.6

39.3 35.5 30.0

33.5 35.8 74.5

29.2 34.5 63.0

5.5 3.3 1.5

14.1 9.8 25.1

113.2 123.0 114.8

1.7 1.0 1.4

204.4 221.2 99.0

Buccheri Querceto A R

142.9 168.7

476.1 401.5

41.3 27.7

35.6 64.1

34.9 38.9

3.8 1.7

17.9 18.0

111.4 116.9

1.4 1.9

138.4 137.0

M. Lauro 1 A R

155.8 161.8

495.2 435.1

50.6 34.5

15.0 61.7

19.2 52.8

7.0 4.7

9.0 22.1

127.1 130.3

3.5 1.3

165.6 87.0

M. Lauro 2 A B C R

152.2 161.3 160.9 158.0

467.4 480.2 451.1 479.1

47.8 48.2 29.6 34.6

11.6 22.6 14.6 65.6

12.8 17.4 18.1 45.7

9.3 6.5 3.1 1.7

8.9 15.6 4.4 22.5

90.6 128.1 130.8 123.8

3.5 1.8 2.3 1.3

166.7 150.1 212.7 79.5

M. Lauro 3 A Bss1 Bss2 Bss3 Bss4 Bss5 R

132.3 140.9 135.6 138.1 135.1 137.2 163.8

529.6 483.3 460.8 465.9 455.7 451.8 485.1

36.7 38.8 38.1 37.4 37.5 35.4 34.8

9.9 12.3 10.6 12.4 12.6 14.0 70.7

10.2 9.1 10.7 11.8 11.7 12.9 46.4

9.3 9.8 9.8 10.1 9.7 9.8 3.6

6.8 8.4 8.5 9.9 9.9 10.2 24.3

83.4 87.6 86.0 86.8 86.1 86.9 132.8

3.4 2.1 2.7 2.7 3.2 3.4 2.1

174.6 157.7 161.1 152.4 157.6 150.4 17.8

M. Lauro 4 A Bss1 Bss2 Bss3 R

168.4 174.9 173.2 172.2 165.7

460.7 473.3 473.0 463.3 507.1

27.4 26.2 25.7 28.6 29.1

40.6 34.0 29.6 29.9 79.6

16.2 18.3 16.7 17.0 66.2

5.7 5.7 6.0 5.7 3.0

4.5 5.0 4.0 3.8 24.6

103.6 104.4 102.0 106.8 126.9

2.7 1.9 2.0 2.0 1.9

170.1 156.4 167.9 170.7 12.3

LOI = loss on ignition.

parent material, the soil is depleted in CaO, MgO and Na2O. Also the variability of the Al2O3 and Fe2O3 content between the soils (and within the soil profile) is rather small. 4.2. Soil and clay mineralogy All soils contained smectite, kaolinite and interstratified smectite–kaolinite in the clay fraction. Additionally, some mica, vermiculite and HIV could be detected in various proportions. In most cases, X-ray diffraction patterns of the Mg-saturated clay samples showed two fundamental reflections near 1.44 nm and a right-skewed peak at 0.70 nm. Ethylene glycol solvation gave rise to a shift of the peak at 1.44 nm to 1.70 nm (Fig. 2). Additional reflections were now measured near 0.85 nm, 0.77 nm and 0.71 nm (e.g. Radura 0–25 cm, Fig. 2). Smectite was identified by the pronounced peak near 1.70 nm after ethylene glycol solvation. The peak at 0.71 nm can fully be attributed to kaolinite (or halloysite) as no chlorite was detected (no peak in the Mg-saturated sample at 1.4 nm with a theoretical d(002) reflection at 0.71 nm; and also no peak near 1.4 nm after heating at 550 °C). The peak height at 0.71 nm correlated with the d(002) of kaolinite at 24.9° 2θ. As halloysite is fibrous, it

cannot be oriented. The 00l peaks would be at about 20 and 35° 2θ (Moore and Reynolds, 1997). No peaks or only “traces” of a peak were discernible at 20° 2θ. Halloysite played, therefore, a subordinate role. The peak at 0.78 nm (EG solvation) corresponds to an interstratified smectite–kaolinite (with a high proportion of kaolinite; approximately 70%; Moore and Reynolds, 1997) and the one at 0.85 nm to the higher order reflection d(002) of smectite. Interstratified smectite–kaolinite was found between 0.77 and 0.82 nm corresponding to 50 to N90% of kaolinite layers (Moore and Reynolds, 1997). The weak peak near 1.0 nm after K-saturation reflects some vermiculite. The increase of the peak at 1.0 nm after K-saturation was more pronounced in the soil of M. Lauro 2. Besides vermiculite also hydroxy-interlayered vermiculite (HIV) was present, indicated by the remaining peak at 1.40 nm after K-saturation and the peak at 1.13 nm after heating at 335 °C. Heating at 550° caused a full collapse of this peak. The C-horizon of M. Lauro 2 contained a high amount of smectite, a low concentration of kaolinite and traces of vermiculite. In the investigated soils, the concentration of smectite generally decreased towards the surface while kaolinite increased (Fig. 2).

V. Barbera et al. / Geoderma 143 (2008) 191–205

197

Fig. 2. X-ray patterns of soil clays (b2 μm) of the top- and subsoil of some typical profiles. The XRD-curves are smoothed and corrected for Lorentz and polarisation factors. d-spacings are given in nm. Mg = Mg saturation, EG = ethylene glycol solvation, K = K-saturation and corresponding heating treatment.

X-ray measurements in the region 2−30° 2θ evidenced that almost no quartz and no feldspars were present (data not shown). Some soils had traces of Fe-oxyhydroxides. This means that the clay fraction of the investigated soils predominantly consists of clay minerals. All soils contained a considerable amount of both crystalline and poorly crystalline Al- and Fe-phases. A large part of Fe was measured in the crystalline form (Table 6). The Alo and Sio contents were in all soils similar and only, if ever, weak trends in the profile were measured. The molar ratio of (Alo − Alp)/Sio was – with the exception of the site Lauro 4 – much lower than 1. This ratio should be close to 1 for allophanes and allophane-like minerals

(2SiO2·Al2O3·2.5H2O; Lorenzoni et al., 1995) and close to 2 for imogolite-like minerals (SiO2·Al2O3·2.5H2O). The likelihood of ITM is very low if the molar ratio is b0.75 or N 2.4. According to this ratio, almost no ITM was present in the investigated soils. Only in the soil profile Lauro 4, this ratio varied between 1.7 and 2.1. Due to the vertic characteristics of this soil, no trend along the profile could be discerned. 4.3. Mass balance calculations Element leaching from the soil profile varied among the different sites (Table 7). The deeper the whole soil profile, the

198

V. Barbera et al. / Geoderma 143 (2008) 191–205

Table 6 Fe-, Al- and Si-phases in the soils (with o = oxalate-extractable content, p = pyrophosphate-extractable content, d = dithionite-extractable content) Alo

Ald

Alp

Feo

Fed

Fep

Sio

Sip

ITM

Ferrihydrite

Fed − Feo

Feo − Fep

g kg− 1

g kg− 1

g kg− 1

g kg− 1

g kg− 1

g kg− 1

g kg− 1

g kg− 1

g kg− 1

g kg− 1

g kg− 1

g kg− 1

Buccheri Radura A 2.03 Bw 2.56

1.66 1.72

1.79 1.78

7.53 3.54

19.74 10.96

2.36 2.06

0.91 0.97

3.76 5.00

6.46 6.86

12.79 6.02

12.22 7.42

5.17 1.49

0.26 0.80

Buccheri Querceto A 1.51

1.16

1.54

3.42

12.73

2.13

0.61

4.35

4.36

5.81

9.32

1.28

− 0.04

Monte Lauro 1 A 2.05

3.24

3.28

5.82

30.84

3.09

0.55

5.06

3.94

9.90

25.02

2.73

− 2.21

Monte Lauro 2 A 2.01 Bw 1.82 C 1.33

3.46 2.27 0.99

2.18 1.63 1.53

4.51 2.80 1.13

36.94 27.80 8.30

1.83 1.55 1.58

0.63 0.85 0.74

3.60 3.60 4.80

4.46 6.04 5.24

7.66 4.77 1.91

32.43 25.00 7.17

2.67 1.25 − 0.46

− 0.28 0.22 − 0.27

Monte Lauro 3 A 1.60 Bss1 1.47 Bss2 1.56 Bss3 1.46 Bss4 1.41 Bss5 1.40

1.43 1.34 1.42 1.26 1.23 1.21

2.92 3.02 3.23 1.85 1.57 1.10

4.57 3.58 4.42 3.10 3.16 2.56

15.33 13.05 14.82 13.53 13.05 12.90

2.52 2.84 2.93 1.76 1.54 1.08

0.89 0.82 0.83 0.84 0.84 0.89

6.42 7.15 7.84 4.35 3.87 2.76

6.32 5.79 5.91 5.93 5.97 6.32

7.76 6.09 7.51 5.21 5.37 4.35

10.76 9.47 10.41 10.47 9.90 10.34

2.05 0.74 1.49 1.31 1.62 1.48

− 1.48 − 1.90 − 2.08 − 0.47 − 0.20 0.33

Monte Lauro 4 A 2.18 Bss1 2.23 Bss2 2.34 Bss3 2.30

2.09 2.37 2.48 2.61

0.16 0.23 0.21 0.24

4.46 4.26 3.60 2.93

18.48 18.88 18.13 18.46

0.24 0.30 0.31 0.35

1.15 1.17 1.06 0.98

0.54 0.57 0.71 0.64

8.17 8.32 7.52 6.98

7.59 7.25 6.11 4.98

14.02 14.62 14.56 15.54

4.22 3.97 3.28 2.58

1.76 1.71 2.01 2.09

Site/horizon

higher was generally element leaching. A very good relationship between element losses and soil profile depth exists for Mg (R2 = 0.88; p b 0.01) and Na (R2 = 0.88; p b 0.01); the deeper the soil, the higher the overall losses. No correlation between soil depth and losses of Si, Al or Fe was found. As the parent material and the age of the soils are similar, differences in weathering can be attributed to erosion and accumulation processes. The open-system mass transport function τ gave negative values and, thus, losses of elements especially for the elements Na, Ca and Mg. The values for Ca were between − 46% and − 87%, for Mg between − 44% and − 82% and for Na between − 38% and − 84%. This demonstrates that a substantial part of the original amount has been leached during pedogenesis. Losses could be measured for all elements except for K. K seems to be retained very strongly in the interlayers of clay minerals and, furthermore, to be strongly integrated in the element-cycling between plants and soil. The ecosystem-internal cycling of K is often very pronounced (Egli, 1998). Some K might also have been introduced into the soil by manure and/or loess (or deposition of sahalien dust (?): generally, the ticker the soil column, the greater are the gains in K (R2 = 78; p b 0.01)). The input of loess or sahalien dust, however, must have been very low (due to the low quartz content). The soils M. Lauro 1–3 are situated along a small slope (with M. Lauro 1 at the top, M. Lauro 2 in the middle and M. Lauro 3 at the footslope). Classically, the thickest soil profile was found

(Alo − Alp)/Sio

at the footslope. Weathered soil material was eroded from the area uphill and deposited in the flatter area. The measured element losses reflect this kind of effect. The element losses per normalised horizon (0–25 m) are higher at M. Lauro 1 and 2 than at M. Lauro 3 (where chemical weathering is overshadowed by inputs due to accumulation). The site M. Lauro 4 shows a slight accumulation of Al, K and partially of Si in the soil. This could be either due to some accumulation (from erosion processes). 4.4. Soil organic matter Organic C concentration varied between 14.9 and 42.1 g kg− 1 in the topsoil (Table 8). The self-mulching effect in the soils with vertic properties led to an only weakly decreasing tendency of organic carbon with soil depth (M. Lauro 3 and 4). The stored amount of organic carbon in the whole soil profile is considerable and reaches values of up to 28 kg/m2 (M. Lauro 4). The soils M. Lauro 2–4 generally had a higher C/N ratio. This might be primarily due to the vegetation (which is different at the sites Radura and Querceto). Carbon concentration after oxidation by H2O2 ranged from 0.29–5.36 g kg− 1. This corresponded to 1–28% of total soil organic carbon. The highest ratios were measured at the sites M. Lauro 3 and 4. In some soils, the oxidation efficiency decreased with soil depth. The concentration of N after oxidation was generally low. Compared to the untreated soils, the H2O2

V. Barbera et al. / Geoderma 143 (2008) 191–205

199

Table 7 Losses and gains of elements in different soil depths (horizons), the whole soil profile and standardised to the initial columnar height of 25 cm (according to Eq. (3)) Site

Soil depth (cm)

Al

Si

Ca

Mg

K

Na

Fe

Mn

Sum

− 30.35 − 30.86 − 61.21 − 29.76 − 43.91 − 43.91 − 21.66 − 36.90 − 36.90 − 20.46 − 33.96 − 32.75 5.63 − 61.08 − 29.76 3.11 − 7.35 − 11.75 − 9.30 − 8.13 − 4.98 − 38.4 1.46 − 1.71 2.65 4.72 − 3.11 2.53 0.92

− 16.04 − 19.48 − 35.52 − 15.73 − 27.22 − 27.22 − 13.43 − 29.88 − 29.88 − 16.56 − 20.99 − 18.33 − 1.97 − 41.29 − 18.39 − 16.75 − 12.95 − 17.07 − 16.38 − 11.83 − 10.75 − 85.74 − 19.65 − 5.13 − 9.14 − 11.77 − 6.51 − 32.55 − 14.23

− 11.25 − 12.57 − 23.83 − 11.04 − 9.36 − 9.36 − 4.62 − 19.41 − 19.41 − 10.76 − 11.28 − 10.41 − 1.11 − 22.81 − 9.89 − 8.47 − 7.00 − 8.64 − 8.27 − 6.01 − 5.37 − 43.76 − 10.03 − 5.81 − 8.44 − 10.18 − 5.45 − 29.88 − 14.21

1.03 0.66 1.70 1.01 0.54 0.54 0.27 0.08 0.08 0.04 2.14 1.30 2.51 5.94 1.88 1.67 1.31 1.76 1.87 1.26 1.34 9.21 1.96 0.50 0.85 1.12 0.42 2.90 1.35

− 4.89 − 7.71 − 12.6 − 4.8 − 4.61 − 4.61 − 2.27 − 9.63 − 9.63 − 5.34 − 6.12 − 4.35 − 1.55 − 12.02 − 5.36 − 5.06 − 3.77 − 4.81 − 4.33 − 3.14 − 2.79 − 23.91 − 5.90 − 2.88 − 4.32 − 5.30 − 2.84 − 15.35 − 7.18

− 9.11 − 4.71 − 13.82 − 8.93 − 29.98 − 29.98 − 14.79 − 24.74 − 24.74 − 13.72 − 13.16 − 11.61 3.79 − 20.97 − 11.53 − 14.33 − 11.52 − 14.85 − 14.08 − 10.35 − 8.73 − 73.86 − 16.91 − 2.35 − 2.34 − 2.86 − 2.36 − 9.91 − 4.67

− 0.03 − 0.24 − 0.27 − 0.03 − 0.73 − 0.73 − 0.36 0.01 0.01 0.01 0.47 − 0.02 0.66 1.10 0.41 0.34 − 0.04 0.12 0.14 0.21 0.25 1.02 0.33 0.15 0.05 0.12 0.03 0.35 0.20

− 79.1 − 76.0 − 155.1 − 77.6 − 153.4 − 153.4 − 75.7 − 144.3 − 144.3 − 80.0 − 96.0 − 87.9 11.6 − 172.2 − 84.1 − 47.2 − 47.8 − 63.5 − 57.6 − 43.7 − 35.1 − 294.9 − 57.9 − 15.9 − 16.1 − 18.4 − 18.9 − 69.3 − 31.9

kg m− 2 Radura

Querceto

M. Lauro 1

M. Lauro 2

M. Lauro 3

M. Lauro 4

0–25 25–65 ∑ whole profile ∑ 0–25 cm (normalised) 0–40 ∑ whole profile ∑ 0–25 (normalised) 0–35 ∑ whole profile ∑ 0–25 (normalised) 0–35 35–65 65–85 ∑ whole profile ∑ 0–25 (normalised) 0–25 25–45 45–70 70–100 100–120 120–140 ∑ whole profile ∑ 0–25 (normalised) 0–20 20–50 50–90 90–110 ∑ whole profile ∑ 0–25 (normalised)

− 8.44 − 1.10 − 9.54 − 8.28 − 38.18 − 38.18 − 18.83 − 23.83 − 23.83 − 13.21 − 13.07 − 11.66 3.66 − 21.07 − 11.46 − 7.73 − 6.49 − 8.28 − 7.22 − 5.70 − 4.02 − 39.43 − 9.18 1.33 4.61 5.73 0.92 12.59 5.92

Losses are expressed by negative values and gains by positive values.

treatment led to a relative enrichment of nitrogen (lower C/N ratios after the treatment). The following main signals with respect to soil organic matter could be detected in the FT-IR spectra: at 1510 cm− 1, 1630 cm− 1 and at 2920 cm− 1. The band at 2920 cm− 1 represents aliphatic C–H stretching (aliphatic methyl and methylene groups). The band at 1630 cm− 1 can be assigned to C_O vibrations of carboxylates and aromatic C_C and the peak at 1510 cm− 1 primarily to aromatic rings and partially to other components such as amide II vibrations (Hesse et al., 1995). As confirmed by the broad band at 1630 cm− 1, the most prevalent functional groups are characterised by C_O of carboxylic and amide groups and by C_C of aromatic ones, that amount to 55% to 94% of the total. Values of aliphatic C–H stretching (2920 cm− 1) range from 4% to 41% while aromatic rings and amide II range from 0.7% to 4%. The H2O2 treatment caused a significant change in the composition of organic matter (Fig. 3). The SOM after the treatment showed higher relative peak intensities at 2920 cm− 1 and, thus, an enrichment (relative to the other functional groups) in aliphatic C–H compounds. The relative peak intensities at 1630 cm− 1 generally decreased after the treatment while no significant changes could be observed for the relative peak intensities at 1510 cm− 1. The ratio of the peak intensities 1630 cm− 1/2920 cm− 1 (data not shown) gave usually the highest values in the A horizon. This indicates that the aromaticity decreases with soil depth.

5. Discussion 5.1. Mineralogy The clay fraction of the soils was dominated by kaolinite and smectite, while mica and vermiculite were less important. This was supported by both clay quantification methods. The CEC method gave, however, distinctly higher kaolinite concentrations (mean = 45%) than the peak-area method (mean = 13%). The CEC method has several limitations: it considers only kaolinite and smectite (although in some soils other minerals were present) and it assumes that oxyhydroxides and OM do not much contribute to the CEC. The correlation analyses of soil properties with kaolinite and smectite using the CEC method did not always give identical values to the area method (data not shown), but the general conclusion was the same. The area method most probably slightly underestimated the kaolinite and overestimated the smectite content. This method has less stringent limitations and seems, therefore, to produce more reliable results. The clay minerals in soils developed from basaltic rocks are assumed to be newly formed products which derived from weathering of rock minerals, because basalt commonly does not contain phyllosilicates as rock-forming minerals. Amphiboles, mica, volcanic glass or ashes may act as a source for smectite formation and plagioclase of volcanic glass as a source of

200

V. Barbera et al. / Geoderma 143 (2008) 191–205

Table 8 Carbon and nitrogen concentrations in the soil before and after the H2O2 treatment, stocks of C and N, proportion of stable organic carbon (SOC) on total organic carbon (TOC) and concentration of N and C in the easily oxidisable (e.o.) fraction Site/horizons

TOC −1

TOC

N −2

N −1

C/N −2

SOC −1

SOC/TOC

N after H2O2 −1

g kg

kg m

g kg

kg m

Buccheri Radura A 42.3 Bw 6.4 Sum

14.81 3.31 18.12

3.81 0.57

1.33 0.30 1.66

11.1 10.7

0.46 0.29

1.1 4.6

0.11 0.23

Buccheri Querceto A 17.1 Sum

9.76 9.76

1.18

0.67 0.67

14.2

0.61

3.6

Monte Lauro 1 A 19.8 Sum

8.1 8.1

1.41

0.58 0.58

14.1

1.16

Monte Lauro 2 A 14.9 Bw 3.5 C 2.3 Sum

5.52 1.32 0.51 7.35

0.93 0.13 0.08

0.35 0.05 0.02 0.42

16.6 26.9 28.8

Monte Lauro 3 A 20.7 Bss1 9.8 Bss2 12.4 Bss3 8.1 Bss4 9.1 Bss5 7.3 Sum

7.49 2.67 4.44 2.91 2.36 1.89 21.76

1.65 0.55 0.72 0.43 0.58 0.34

0.6 0.15 0.26 0.15 0.15 0.09 1.4

Monte Lauro 4 A Bss1 Bss2 Bss3 Sum

5.33 7.21 7.94 3.65 27.78

2.09 1.17 1.14 0.88

0.44 0.4 0.46 0.17 1.65

25.5 21.3 19.7 19.4

g kg

g kg

C (e.o.) −1

N (e.o.) −1

C/N (e.o.)

g kg

g kg

4.2 1.3

41.8 6.1

3.70 0.34

11.3 18.0

0.10

6.1

16.5

1.08

15.3

5.9

0.21

5.5

18.6

1.20

15.5

0 0.32 0.64

8.8 9.0 28.1

0.36 0.07 0.08

3.6 4.6 8.0

13.6 3.2 1.7

0.57 0.06

23.8 53.0

12.1 17.8 17.2 18.8 15.7 21.4

0.67 0.72 0.87 1.08 1.13 0.95

3.2 7.3 7.0 13.4 12.5 13.1

0.28 0.08 0.22 0.34 0.28 0.20

2.4 9.0 4.0 3.2 4.0 4.8

20.0 9.1 11.5 7.0 8.0 6.4

1.37 0.47 0.50 0.09 0.30 0.14

14.6 19.3 23.1 78.0 (?) 26.6 45.4

12.2 18.2 17.3 22.0

5.36 3.29 2.98 1.88

21.0 15.4 15.2 9.7

0.44 0.26 0.10 0.30

12.2 12.7 29.8 6.3

20.1 18.0 16.7 17.5

1.65 0.91 1.04 0.58

12.2 19.8 16.1 30.2

kaolinite (Egli et al., 2003). The formation of smectites may also be due to hydrothermal alteration of glass particles during or immediately after the emplacement of the pyroclastic flow (Mirabella et al., 2005). A close relationship exists between

%

C/N after H2O2

smectite and kaolinite (Fig. 4) showing that smectite is actively transformed in the soil into kaolinite. A transitory product in this weathering pathway was interstratified smectite–kaolinite. In the three profiles there is a clear trend with an increasing

Fig. 3. Percentile distribution of the relative proportion of the bands (and corresponding functional groups) at 1510 cm− 1 at 1630 cm− 1 and at 2920 cm− 1 before and after the H2O2 treatment.

V. Barbera et al. / Geoderma 143 (2008) 191–205

201

for kaolinite formation was detectable in the clay fraction of the soils, this type of transformation has been the dominant process. The alteration sequence must therefore have been: amphibole, mica (which were all originally present in the basaltic rock), volcanic glass or ashes to smectite, interstratified smectite– kaolinite and then kaolinite. Element losses are considerable when compared to other areas with available weathering data (Egli et al., 2001). Weathering in Mediterranean areas is usually quite slow due to the subtropical and rather dry climate. High element losses, therefore, indicate that weathering has occurred over a very long time period. Fig. 4. Relationship between the relative concentrations of smectite and kaolinite in the clay fraction (calculated with the area method).

kaolinite content in the clay fraction from the sub- to the topsoil. Especially in profile M. Lauro 2, the smectite to kaolinite transformation could be well discerned involving kaolinite– smectite interstratifications with increasing proportions of kaolinite. Weathering of basic igneous rocks in tropical and subtropical areas leads to neoformation of smectite clays which may transform into kaolinite through mixed-layer kaolinite– smectite (Viangini et al., 2004). The smectite-to-kaolinite transformation involves losses of Fe from the smectite structure with subsequent precipitation of low Fe-kaolinite and secondary Fe-phases. The transformation of the clay types during soil evolution means that soil chemical and physical properties change and consequently also the interaction of inorganic phases with organic matter.

5.3. Soil organic matter The stocks of SOM were in some soils surprisingly high and varied between 7 and 28 kg/m2. This amount is clearly above the expected values for “interrupted” woodland that should be in the order of 4 to 10 kg C/m2 (Post et al., 1982). Mediterranean soils in the investigated region obviously are capable to store a relatively high amount of organic C. Organic matter showed a good correlation with the fine sand fraction of the soils (Fig. 5). This led us to the hypothesis that stable mineral–SOM complexes can be found to a certain extent in this fraction. ITM is known to stabilise soil organic matter. ITM was, however, detectable only in one soil and consequently no correlation

5.2. Mass balances, weathering The comparison among mass losses, kaolinite and smectite concentrations gives a good idea of chemical transformation involved during the pedogenetic process. The higher the Al, Si, Ca, Mg and Na losses, the lower was the measured smectite content (Table 9). Kaolinite shows an opposite trend: the higher the losses of Al, Ca, Mg and Na, the higher is the concentration of kaolinite in the soil. More intense weathering conditions, therefore, lead to a continuous transformation of smectite into kaolinite. Because no other mineral acting as a potential source

Table 9 Correlation coefficients R between smectite and kaolinite and element losses Losses of

Absolute concentration

Relative concentration

Smectite

Kaolinite

Smectite

Kaolinite

Al Si Ca Mg K Na Fe Mn

− 0.76⁎⁎ − 0.49⁎ − 0.79⁎⁎ − 0.89⁎⁎ − 0.71⁎ − 0.87⁎⁎ − 0.75⁎⁎ − 0.48

0.56⁎ 0.48 0.68⁎⁎ 0.50⁎ 0.07 0.76⁎⁎ 0.52⁎ 0.18

− 0.43 − 0.45 − 0.72⁎⁎ − 0.62⁎ − 0.66⁎ − 0.69⁎⁎ − 0.41 − 0.29

0.65⁎⁎ 0.60⁎ 0.69⁎⁎ 0.65⁎⁎ 0.17 0.65⁎⁎ 0.56⁎ 0.36

Absolute concentration = clay fraction × concentration in clay fraction, relative concentration: the sum of clay minerals is 100%. ⁎Error probability p b 0.05, ⁎⁎error probability p b 0.01.

Fig. 5. Correlations of (upper) fine sand and (lower) clay with total organic carbon.

202

V. Barbera et al. / Geoderma 143 (2008) 191–205

correlation between the amount of clays and SOM was calculated (Fig. 5). If clays were associated with SOM then a positive correlation should be expected. The clay content was therefore not a good predictor for organic matter. In literature, one does not find an unequivocal evidence that clays play a role in stabilising organic matter in soils. The amount of clayassociated organic matter is rather influenced by the type of clays. In the investigated soils, complexing reactions between Fe, Al, kaolinite and SOM contribute to the accumulation and stabilisation of organic matter while smectite had no effect on the SOM preservation. To estimate the refractory soil organic matter pool, wet oxidation with H2O2 was used (Eusterhues et al., 2005). Total organic carbon in the soils was related to the concentration of stable organic carbon. The correlation was in the topsoil lower than in the subsoil (Fig. 7). The relationship in the topsoil might be overshadowed by the continuous input of fresh organic carbon while this is not the case in the subsoil. The oxidationresistant organic matter fraction was enriched in aliphatic components (Fig. 3) and in N when compared to the untreated samples (Table 8). This is in agreement with Leifeld and KögelKnabner (2001), Cuypers et al. (2002) and Eusterhues et al. (2005) who found oxidation-resistant organic matter enriched in aliphatic C. Baldock et al. (2004) found that carbohydrates distinctly decreased during decomposition of organic matter. Proteins, lignins and partially lipids relatively increased and

Fig. 6. Relationships of (upper) oxalate-extractable Fe and (lower) kaolinite (calculated with the area method) with total organic carbon.

with organic matter was measured. Close relationships exist between organic C and the oxalate-extractable Fe- and Alfractions (cf. Fig. 6). Kaolinite displays, furthermore, a weak correlation with SOM (Fig. 6). Similar to the findings of Wiseman and Püttmann (2006), a good correlation between kaolinite and Feo or Alo was detected. This suggests that kaolinite interacts with oxides. Kaolinite complexes with oxyhydroxides are likely to play an important role in the preservation of organic carbon. The existence of statistical correlations between Feo, Alo and kaolinite does not mean that the binding relationships between oxides and kaolinite are straightforward; that is, organic material may be bound to oxides which are, in turn, bound to clay minerals (cf. Wiseman and Püttmann, 2006). The curvilinear shape of the correlation between kaolinite and organic C suggests that a high kaolinite content does not offer attractive surface sites to organic matter. When the concentration of kaolinite increases, poorly crystalline Al- and Fe-phases, which usually readily bind to OM, are also attracted to kaolinite surfaces and offer consequently increasingly less area to OM. According to Denef and Six (2004) soils with kaolinite can easily form (macro)aggregates independent of biological processes due to physical or electrostatic interactions between the 1:1 clay minerals and oxides. No correlation between pyrophosphate-extractable Fe and Al and organic C was found. Surprisingly, a negative

Fig. 7. Relationships between organic C content before and after the H2O2 treatment plotted as a function of (upper) all samples and (lower) only the subsoil samples.

V. Barbera et al. / Geoderma 143 (2008) 191–205

were more protected from decay in forest and agricultural soils. The protection from decay was explained by an immobilisation of proteins (due to inorganic N; cf. Kleber et al., 2007) and by the greater biochemical recalcitrance of lipids and lignins. The H2O2 treatment preferentially destroyed C_O of carboxylic and amide groups and C_C of aromatic ones. Kaolinite is likely to be associated with components of the bands at 1510 cm− 1 (components which are more easily oxidised by H2O2), because a negative correlation between kaolinite and the ratio of C after the H2O2 treatment to the total organic C exists (Fig. 8). Kaolinite correlated with the bands at 1510 cm− 1 (before the H2O2 treatment) with RSpearman = 0.57 (p = 0.02). Aromatic rings have a specific affinity to kaolinite. Because kaolinite has a very low charge, the interaction between this type of organic matter is not predominantly of electrostatic nature but driven by van-der-Waals forces. Hydrophobic substances may strongly associate with non-charged mineral surfaces (Kleber et al., 2007). The high C/N ratio of the easily oxidisable fraction (Table 8) and the presence of aromatic compounds suggest furthermore that a part of this fraction derives from charcoal compounds (and thus showing the influence of fire). Fire-derived compounds are aromatic-rich (charcoal) and should be found in the more easily oxidisable fraction (Eusterhues et al., 2005). Kaolinite probably protects to a certain degree these components from biodegradation (but not from H2O2 oxidation). This, however, would not fully agree with findings of Wattel-Koekkoek et al. (2001) who measured a kaolinite-associated organic matter rich in polysaccharide products (and smectite-associated SOM rich in aromatic compounds). A weak correlation also exists between Alp and Fep and the CH2O2/Ctot ratio. pyrophosphate-extractable Fe and Al contents, normally used as a proxy for Al and Fe in organic complexes, had no relation with total organic C, but were negatively correlated to the peroxidation resistant fraction and positively to the easily oxidisable one. Unlike to the total organic carbon, the stable organic C pool is not correlated to the oxalateextractable contents of Al and Fe. The only significant relationship between CH2O2 and oxyhydroxides exists with the crystalline Fe fraction (Fed − Feo; RSpearman = 0.52, p = 0.037). A part of the stable organic matter seems to form complexes with crystalline,

Fig. 8. Relationship of kaolinite (in % of the clay fraction; area method) with the ratio of stable organic carbon to the total organic carbon.

203

secondary Fe-phases. As the stable organic C pool is enriched in N, we assume that also proteinaceous materials adsorb on crystalline Fe (cf. Baldock et al., 2004; Kleber et al., 2007. CH2O2 is furthermore positively correlated with the coarse silt fraction (RSpearman = 0.58; p = 0.021). The absence of any correlations of CH2O2 with poorly crystalline phases shows the recalcitrant nature of a significant part of this C fraction. After the treatment with H2O2 the remaining functional groups did not correlate with Al- or Fe-phases. CH2O2 is predominantly explained by the components in the FT-IR spectrum at 1630 cm− 1 and 2920 cm− 1. In contrast, the H2O2 oxidisable content (Ctot − CH2O2) significantly correlates with the oxalate-extractable contents (Feo: RSpearman = 0.80; p b 0.01; Alo: RSpearman = 0.49; p = 0.04). This kind of OM fraction was obviously preferentially protected by interaction with poorly crystalline minerals in aggregates of the fine sand (RSpearman = 0.65; p b 0.01) and coarse silt (RSpearman = 0.54, p = 0.02) fraction. 6. Conclusions Mediterranean soils which were transitions between Phaeozems and Vertisols were analysed with respect to weathering conditions, mineral formation and transformation and stabilisation mechanisms of organic matter. The storage capacity of organic matter was in these soils relatively high. During the long soil formation period, the soil material was strongly weathered. The main mineral weathering reaction consists of a continuous transformation of smectite into kaolinite. The formation of kaolinite is bound to elevated element losses. Kaolinite is hypothesised to form (macro)aggregates due to physical or electrostatic interactions between the 1:1 clay minerals and oxides. Oxyhydroxides and kaolinite seem to stabilise organic matter in the soil as they are closely related to soil organic matter. The refractory proportion of organic C was between 1 and 28% of total organic carbon. This fraction was enriched in aliphatic compounds and N and did not greatly interact with the kaolinite or poorly crystalline Fe or Al phases. A part of this fraction (most probably proteins) is bound to crystalline Fe-oxides. The easily oxidisable (with H2O2) fraction correlated well with oxyhydroxides and kaolinite, was enriched in C_O of carboxylic, amide groups, in C_C of aromatic components (charcoal?) and had partially a high C/N ratio. Although kaolinite was related to SOM, the clay content was a bad predictor for organic matter in the investigated soils. The amount of clay-associated organic matter is rather influenced by the type of clays: Kaolinite obviously played a more important role than smectite. The results confirm that stabilisation of soil organic matter is predominantly due to its interaction with mineral phases. Interaction occurs, similarly to the proposed conceptual model of organo-mineral interactions and mechanisms of Kleber et al. (2007), with polar organic functional groups of amphiphiles via ligand exchange with singly coordinated mineral hydroxyls forming stable innersphere complexes, proteinaceous materials adsorbed on charged surfaces (“contact” zone; Kleber et al., 2007) and association of hydrophobic substances with non-charged mineral surfaces (“zone of hydrophobic interactions”; Kleber et al., 2007).

204

V. Barbera et al. / Geoderma 143 (2008) 191–205

Acknowledgements This research was supported by a grant from the University of Palermo. We would like to express our appreciation to B. Kägi for his assistance in the laboratory. We are, furthermore, indebted to Zhifei Liu and to Jock Churchman for their helpful comments on an earlier version of the manuscript. References Baldock, J.A., Masiello, C.A., Gélinas, Y., Hedges, J.I., 2004. Cycling and composition of organic matter in terrestrial and marine ecosystems. Marine Chemistry 92, 39–64. Bosatta, E., Ågren, G., 1997. Theoretical analyses of soil texture effects on organic matter dynamics. Soil Biology and Biochemistry 29, 1633–1638. Brice, W.C., 1978. The Environmental History of the Near and Middle East since the Last Ice Age. Academic Press, London. Brimhall, G.H., Dietrich, W.E., 1987. Constitutive mass balance relations between chemical composition, volume, density, porosity, and strain in metasomatic hydrochemical systems: Results on weathering and pedogenesis. Geochimica et Cosmochimica Acta 51, 567–587. Brueckner, H., 1986. Man's impact on the evolution of the physical environment in the mediterranean region in historical times. GeoJoumal 13, 7–17. Chadwick, O.A., Brimhall, G.H., Hendricks, D.M., 1990. From a black to a gray box — a mass balance interpretation of pedogenesis. Geomorphology 3, 369–390. Cuypers, C., Grotenhuis, T., Nierop, K.G.J., Maniero Franco, E., de Jager, A., Rulkens, W., 2002. Amorphous and condensed organic matter domains: the effect of persulfate oxidation on the composition of soil/sediment organic matter. Chemosphere 48, 919–931. Denef, K., Six, J., 2004. Clay mineralogy determines the importance of biological versus abiotic processes for macroaggregate formation and stabilization. European Journal of Soil Science 56, 469–479. Egli, M., 1998. Estimates of current and historical fluxes of N, Ca, and K in the forest ecosystem Möhlin. GeoJournal 44, 73–89. Egli, M., Fitze, P., 2000. Formulation of pedologic mass balance based on immobile elements: a revision. Soil Science 165, 437–443. Egli, M., Fitze, P., Mirabella, A., 2001. Weathering and evolution of soils formed on granitic, glacial deposits: results from chronosequences of Swiss alpine environments. Catena 45, 19–47. Egli, M., Mirabella, A., Sartori, G., Fitze, P., 2003. Weathering rates as a function of climate: results from a climosequence of the Val Genova (Trentino, Italian Alps). Geoderma 111, 99–121. Eusterhues, K., Rumpel, C., Kögel-Knabner, I., 2005. Stabilization of soil organic matter isolated via oxidative degradation. Organic Geochemistry 36, 1567–1575. Fierotti, G., Dazzi, C., Raimondi, S., 1988. Carta dei Suoli della Sicilia. Palermo. Regione Siciliana, Assessorato Territorio e Ambiente & Università degli Studi di Palermo, Istituto di Agronomia Generale, Cattedra di Pedologia. Palermo. Gjems, O., 1967. Studies on clay minerals and clay mineral formation in soil profiles in Scandinavia. Norske Skogfersøksvesen 81, 301–415. González-Pérez, J.A., González-Vila, F.J., Almendros, G., Knicker, H., 2004. The effect of fire on soil organic matter — a review. Environment International 30, 855–870. Hesse, M., Meier, H., Zeeh, B., 1995. Spektroskopische Methoden in der organischen Chemie. Georg Thieme Verlag, Stuttgart. Hossner, C.R., 1996. Dissolution for total elemental analysis. In: Sparks, D.L. (ed.), Methods of Soil Analysis, Part 3 Chemical Methods, Soil Science Society of America Inc. and American Society of Agronomy Inc., Madison, Wisconsin, pp. 49–64. Kahle, M., Kleber, M., Jahn, R., 2002. Review of XRD-based quantitative analyses of clay minerals in soils: the suitability of mineral intensity factors. Geoderma 109, 191–205. Kaiser, K., Zech, W., 1999. Release of natural organic matter sorbed to oxides and a subsoil. Soil Science Society of America Journal 63, 1157–1166.

Kleber, M., Sollins, P., Sutton, R., 2007. A conceptual model of organo-mineral interactions in soils: self-assembly of organic molecular fragments into zonal structures on mineral surfaces. Biogeochemistry 85 (1), 9–24. Knicker, H., Skjemstad, J.O., 2000. Nature of organic carbon and nitrogen in physically protected organic matter of some Australian soils as revealed by solid-state 13C and 15N NMR spectroscopy. Australian Journal of Soil Research 38, 113–127. Lanson, B., 1997. Decomposition of experimental X-ray diffraction patterns (profile fitting): a convenient way to study clay minerals. Clays and Clay Minerals 45, 132–146. Laves, D., Jahn, G., 1972. Zur quantitaiven röntgenographischen BodentonMineralanalyse. Archiv für Acker- und Pflanzenbau und Bodenkunde 16, 735–739. Leifeld, J., Kögel-Knabner, I., 2001. Organic carbon and nitrogen in fine soil fraction after treatment with hydrogen peroxide. Soil Biology and Biochemistry 33, 2155–2158. Lentini, F., Carbone, S., Catalano, S., Grasso, M., Monaco, C., 1990. Principali elementi strutturali del Thrust Belt Appenninico- Maghrebide in Sicilia centro-orientale. Memorie della Società Geologica Italiana 45, 495–502. Lorenzoni, P., Mirabella, A., Bidini, C., Lulli, L., 1995. Soil genesis on trachytic and leucitic lavas of Cimini volcanic complex (Latium, Italy). Geoderma 68, 79–99. Lützow, M.v., Kögel-Knabner, I., Ekschmitt, K., Matzner, E., Guggenberger, G., Marschner, B., Flessa, H., 2006. Stabilization of organic matter in temperate soils: mechanisms and their relevance under different soil conditions — a review. European Journal of Soil Science 57, 426–445. McKeague, J.A., Brydon, J.E., Miles, N.M., 1971. Differentiation of forms of extractable iron and aluminium in soils. Soil Science Society of America Proceedings 35, 33–38. Mirabella, A., Egli, M., Raimondi, S., Giaccai, D., 2005. Origin of clay minerals on pyroclastic deposits in the island of Lipari (Italy). Clays and Clay Minerals 53, 409–421. Moore, D.M., Reynolds, R.C., 1997. X-ray Diffraction and the Identification and Analysis of Clay Minerals, 2nd edition. Oxford University Press, New York. Naveh, Z., 1990. Fire in the Mediterranean — a landscape ecological perspective. In: Goldammer, J.G., Jenkins, M.J. (Eds.), Fire in Ecosystems Dynamics: Mediterranean and Northern Perspective. SPB Academic Publishing, The Hague, pp. 1–20. Niederbudde, E.A., Kussmaul, H., 1978. Tonmineraleigenschaften und– Umwandlungen in Parabraunerde-Profilpaaren unter Acker und Wald in Süddeutschland. Geoderma 20, 239–255. Olson, J.S., Watts, J.A., Allison, L.J., 1983. Carbon in Live Vegetation of Major World Ecosystems. Report ORNL-5862. Oak Ridge National Laboratory, Oak Ridge, Tennessee, USA. Parfitt, R.L., Hemni, T., 1982. Comparison of an oxalate-extraction method and an infrared spectroscopic method for determining allophane in soil clays. Soil Science and Plant Nutrition 28, 183–190. Post, W.M., Emanuel, W.R., Zinke, P.J., Stangenberger, A.G., 1982. Soil carbon pools and world life zones. Nature 298, 156–159. Rasmussen, C., Torn, M.S., Southard, R.J., 2005. Mineral assemblage and aggregates control carbon dynamic in a California conifer forest. Soil Science Society of America Journal 69, 1711–1721. Righi, D., Dinel, H., Shulten, H.R., Schnitzer, M., 1995. Characterization of clay–organic matter complexes to oxidation by peroxide. European Journal of Soil Science 46, 423–429. Saggar, S., Parshotam, A., Sparling, G.P., Feltham, C.W., Hart, P.B.S., 1996. 14 C-labelled ryegrass turnover and residence times in soils varying in clay content and mineralogy. Soil Biology and Biochemistry 28, 1677–1686. Schwertmann, U., Niederbudde, E.A., 1993. Tonmineralbestimmung in Böden. In: Jasmund, K., Lagaly, G. (Eds.), Tonminerale und Tone. Struktur, Eigenschaften, Anwendung und Einsatz in Industrie und Umwelt. Steinkopff-Verlag, Darmstadt, pp. 255–265. Senesi, N., D'Orazio, V., Ricca, G., 2003. Humic acids in the first generation of EUROSOILS. Geoderma 116, 325–344. Shoji, S., Fujiwara, Y., 1984. Active aluminium and iron in the humus horizons of Andosols from northeastern Japan: their forms, properties, and significance in clay weathering. Soil Science 137, 216–226. Stevenson, F.J., 1994. Humus Chemistry: Genesis, Composition, Reactions, 2nd ed. Wiley, New York.

V. Barbera et al. / Geoderma 143 (2008) 191–205 Tan, K.H., 2003. Humic matter in soil and the environment. Marcel Dekker, Inc., New York. Theng, B.K.G., Churchman, G.J., Newman, R.H., 1986. The occurrence of interlayer clay–organic complexes in two New Zeland soils. Soil Science 142, 262–266. Viangini, S., Righi, D., Petit, S., Terribile, F., 2004. Mixed-layer kaolinite– smectite minerals in red-black soil sequence from basalt in Sardinia (Italy). Clays and Clay Minerals 52, 473–483.

205

Wattel-Koekkoek, E.J.W., van Genuchten, P.P.L., Buurman, P., van Lagen, B., 2001. Amount and composition of clay-associated soil organic matter in a range of kaolinitic and smectitic soils. Geoderma 99, 27–49. Wiseman, C.L.S., Püttmann, W., 2006. Interactions between mineral phases in the preservation of soil organic matter. Geoderma 134, 109–118. Yaalon, D.H., 1997. Soils in the Mediterranean region: what makes them different? Catena 28, 157–169.