brown Mediterranean soils

Geoderma 239–240 (2015) 156–167

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Mineralogical and chemical variability of mountain red/brown Mediterranean soils Amir Sandler a,⁎, Alain Meunier b, Bruce Velde c a b c

Geological Survey of Israel, 30 Malkhe Israel St., Jerusalem 95501, Israel IC2MP-HydrASA, Université, 40 Avenue Recteur Pineau, 86022 Poitiers, France Laboratoire de Géologie, CNRS URA 8538 Ecole Normale Supérieure, 24 rue Lhomond, 75231 Paris, France

a r t i c l e

i n f o

Article history: Received 20 July 2014 Received in revised form 7 October 2014 Accepted 11 October 2014 Available online xxxx Keywords: Terra rossa Clay minerals Pedogenic illite Mediterranean climate Israel

a b s t r a c t The purpose of the current study was to quantitatively explore the dynamics of clay pedogenesis in red-brown soils of the terra-rossa type and associated soils (mainly Rhodoxeralfs) of the Mediterranean region with respect to bedrock and dust composition, and the impact of the vegetation-translocated elements silicon and potassium. Fifteen non-cultivated, shallow and leached mountain pedons were sampled in Israel along three transects that represent various rainfall ranges and bedrock lithologies. Bulk and clay fraction chemical and mineralogical compositions, as well as grain size, were determined. Quantitative relative analysis of clay minerals displayed a large variability that was largely affected by leaching. Leaching was dependent on rainfall amount and bedrock/soil permeability. In general, if the better soils were drained and leached the higher were the amounts of both kaolinite and illitic phases. Soil clay assemblages lay between two mineralogical end members: one with dominance of smectite and the second with high values of kaolinite and illitization parameters. Potassium and silicon were mobile in soils relative to aluminum, but under high rainfall or certain vegetation, translocation by vegetation reached a degree that promoted illite or smectite formation. Local combinations of rainfall, bedrock/soil permeability, and vegetation translocation of potassium and silicon dictated a large variability in clays, including illite formation, within a restricted group of red-brown Mediterranean soils. © 2014 Elsevier B.V. All rights reserved.

1. Introduction Mountain soils developed under Mediterranean climate on hard limestone and dolostone are generally leached of carbonate minerals and are of brown to red colors. The typical soil in these environments is terra rossa, a term which is still in use in soil nomenclature of several Mediterranean countries, and mostly corresponds to Rhodoxeralfs. The genesis of terra rossa has been debated for decades (see reviews in Durn, 2003; Merino and Benerjee, 2008; Lucke et al., 2012). The profile of shallow terra rossa soils in Israel is generally AR, whereas in deeper soils the B horizon might be recognized but textural differences between the A and B horizons are faint. The contact between the A or B to R horizon is generally sharp (Ravikovitch, 1969; Dan et al., 2007). An associated soil, which forms under similar conditions but whose genesis is rather disputed as well, is Mediterranean brown forest soil. It has been suggested that it developed above soft limestone or exposed calcrete (nari) with rich vegetation (Ravikovitch, 1969). Dan et al. (2007) noted that Mediterranean brown soil developed on dolomite whereas brown forest soil developed on exhumed calcrete. However, Rabinovitch-Wein (1986), based on her observations in the Galilee, ⁎ Corresponding author. E-mail address: [email protected] (A. Sandler).

http://dx.doi.org/10.1016/j.geoderma.2014.10.008 0016-7061/© 2014 Elsevier B.V. All rights reserved.

suggested that brown forest soils developed either on limestone or dolostone with alternations of soft clayey beds. Furthermore, brown forest soil was not included in the classification of the soils of Israel (Dan et al., 1976), as it was considered brown rendzina. Early studies in Israel considered terra rossa as a residual product of the bedrock (Reifenberg, 1947; Ravikovitch, 1969; Gal et al., 1974), while other studies recognized and emphasized the significant contribution of dust (Yaalon et al., 1966; Singer, 1967; Yaalon and Dan, 1967; Yaalon and Ganor, 1973; Dan, 1990). Based on clay-to-silt ratios in dust and on the insoluble fraction of limestone in soils, the dust contribution to terra rossa soils was estimated as one-third to two-fifths (Yaalon and Ganor, 1973). By comparing rates of dust accretion to hard limestone dissolution it was estimated that eolian material makes up to 50% of the fine fraction in soils on hard limestone rocks (Yaalon, 1997). Imprints of endolithic lichens on dolostone surfaces below terra rossa led Danin et al. (1983) to conclude that eolian dust is the main source of fine soil particles. Hence, both bedrock insoluble residue and eolian material contribute to terra rossa formation. Most studies on red Mediterranean soils focused on two main processes that apparently contribute to their unique appearance and characteristics: rubefacation by iron minerals and clay illuviation (review by Fedoroff and Courty, 2013). Many of those studies reported on clay minerals, mainly illite and kaolinite (e.g. Boero et al., 1992), which were generally

A. Sandler et al. / Geoderma 239–240 (2015) 156–167

considered as inherited from parent material. However, kaolinite has been considered as a pedogenic product, at least in part (e.g. Durn et al., 1999). Only very few studies referred to the formation and stability of other clay minerals in red Mediterranean soils (e.g. Torrent and Cabedo, 1986). In terra rossa soils in Israel either kaolinite or montmorillonite (smectite) has been considered as the dominant clay. Koyumdjiski (quoted in Singer, 2007) suggested that terra rossa on dolomite was dominated by montmorillonite due to the continuous supply of magnesium, whereas on limestone it was dominated by kaolinite. A contradictory observation (Rabinovitch-Wein, 1986) indicated that the bedrock period, and not its lithology, is the key factor: montmorillonite dominated terra rossa on Cretaceous limestone and dolostone, whereas kaolinite dominated terra rossa developed on Eocene limestone. Singer (2007 p. 98) concluded that in red brown terra rossa soils smectite is the major clay mineral, with secondary kaolinite and traces of illite whereas in red terra rossa soils kaolinite is the major clay, accompanied by illite, smectite, and smectite/illite. Though illite has been reported reaching ~30% of the clay fraction, as secondary to kaolinite (70%), in a red terra rossa soil on Eocene limestone (Koyumdjisky et al., 1988) its origin has been overlooked. Exceptional high illite content in terra rossa has been reported also in two near-by locations with higher precipitation. In the Anti-Lebanon and Lebanon mountains illite origin was overlooked as well (Darwish and Zurayk, 1997), whereas at Mt. Hermon (Singer, 1978) it was attributed to inheritance from bedrock. In a recent study that examined and compared the mineralogical composition of the clay fraction in soils throughout the various climate regimes of Israel, it has been suggested that clays of Mediterranean soils are largely pedogenic (Sandler, 2013). Smectite forms in low-permeable but leached clayey soils, and reaches optimum evolution with ~ 90% smectitic IS in certain grumosols (vertisols). This evolutionary trend was also recognized in poorly-drained soils of catenary chains that are related to other soil types such as terra rossa, hamra and pale rendzina (Israeli soil classification). Kaolinite forms in well-drained red brown sandy (hamra) and terra rossa soils, developed on coastal sand or hard carbonate, respectively, to the detriment of smectitic IS phases of parent materials. Kaolinite becomes the principal mineral in evolved terra rossa and hamra soils and illite is often significantly enriched. The aim of the current study was to quantitatively explore the dynamics of clay pedogenesis in red-brown soils of the terra-rossa type and associated soils. The soils studied were based on hard limestone and dolostone and under Mediterranean climatic zones (see details below). Clay pedogenesis was evaluated with respect to the mineralogical composition of bedrock and dust and the chemical composition of bulk soils and their clay fraction. The role of the vegetation-translocated elements silicon and potassium, which are crucial for clay formation of illite–smectite phases of the smectite-to-illite series, was evaluated as well. The different soil compositions were related to current environmental conditions in order to decipher the factors that govern soil variability within this relatively limited range of soil types. 2. Study area and sampling strategy A mountain range of sedimentary rocks located in central Israel between the Mediterranean coastal plain and the Dead Sea rift, which includes the Judea and Samaria anticlines, Mt. Carmel and the Galilee mountains, exposes mainly Late Cretaceous to Eocene carbonate rocks. Terra rossa forms on hard carbonate rocks, pale rendzina forms on chalk and marl rocks, whereas brown rendzina and brown forest soils are intermediate types developed on various bedrocks (Singer, 2007 and references therein). Fifteen soil pedons were sampled in three transects (Fig. 1): a. Judea Mts., from Jerusalem westward (16 samples); b. Eocene transect, from the Lower to the Upper Galilee, on the same limestone formation (6 samples); c. Galilee mountains, from the highest summit

157

Fig. 1. Location map. Lower gray patch frames the Judea transect (a) of soils based on Cretaceous limestone and dolostone. Upper gray patch frames the Eocene limestone transect (b) and the Galilee transect (c) of soils based on Cretaceous limestone and dolostone.

(Mt. Meron) westward (14 samples). The a and c transects are on various Cretaceous rocks (Fig. 1, Table 1). The elevation ranges of the three transects are 325–835, 400–685, and 80–1150 m above sea level (asl), and the mean annual precipitation (referred to as “rainfall”) ranges are 475–700, 550–675, and 625–950 mm, respectively. The rainfall values were extracted from a rainfall map for 1950–1980, prepared by the Ministry of Agriculture. Most sampled locations are under the Mediterranean Csa regime, and a few are under the colder Csb regime according to the Köppen climatic classification in which C indicates temperate climate, s indicates dry simmer, a indicates warmest month average temperature above 22 °C with at least four months averaging above 10 °C, and b indicates warmest month averaging below 22 °C, but with at least four months averaging above 10 °C (Goldreich, 2003). All soils sampled overlie hard limestone or dolostone bedrock, except two soft marly dolostones on Mt. Meron. The selected pedons were of shallow soil depths and away from apparent colluvium contribution. Depth sampling of soil profiles included mostly topsoil and subsoil samples in order to represent potential contrasts between organic-rich and latest dust contribution at the top and the main body of soil solum below. Topsoil samples were generally of 0–10 cm depth, or 0–5 cm, in cases where dark organic matter was distinct at the upper few centimeters. A few relatively thick profiles were sampled in addition to a greater depth. Settled dust was collected in plastic trays filled with glass marbles, washed in the laboratory with distilled water and dried up.

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Table 1 General information on soils sampled: sample number, depth of the soil profile and of each sample, location name, soil classification, coordinates, mean annual precipitation, elevation above sea level, bedrock geological formation name and lithology, geomorphological slope calculated from DTM map for 100 × 100 m around sampling point, basic vegetation information. Sample

Depth cm

a. Judea transect 561 70 a 0–5 b 5–15 c 20–40 562 30 a 0–10 b 15–30 563 12 a 0–10 565 55 a 0–10 b 15–25 c 30–45 566 150 a 0–20 b 40–60 c 100–120 567 45 a 0–10 b 15–30 568 35 a 0–10 b 15–30 b. Eocene transect 570 N70 a 0–10 b 15–30 571 N60 a 0–10 b 15–30 572 40 a 0–10 b 15–30 c. Galilee transect 574 30 a 0–10 b 15–30 575 N80 a 0–10 b 15–30 576 55 a 0–5 b 15–25 c 30–45 577 50 a 0–10 b 15–30 578 35 a 0–10 579 45 a 0–5 b 5–15 580 N40 a 0–5 b 10–25

Location, soil classification

Coordinates Israel grid

Rainfall mm

Elev. m asl

Formation

Lithology

Slope °

Remarks

M. HaHamisha Typic Haploxerol

211597/635985

700

795

G. Ye'arim

Dolomite

12

Stony, maquis Organic-rich

Q. ‘Anavim Typic Rhodoxeralf

211994/635383

650–700

705

Kesalon

Nebbi Samuel Typic Rhodoxeralf Illan stream Typic Rhodoxeralf

217149/637476

600–650

835

Bet-Meir

203070/636000

600

325

Ramat Shelomo Typic Rhodoxeralf

220369/635147

550

Cesar road Typic Rhodoxeralf

201693/623706

G. HaMatos Typic Rhodoxeralf

6

With calcrete debris Local depression, mixed vegetation Few mm gray organic matter at top

Dolomite

8

Very thin, dwarf-shrub steppe and pines

B'ina

Limestone

18

Karstic depression, mixed vegetation

720

Veradim

Dolomite

11

Pine woods

500

410

B'ina

Limestone

13

Above laminar calcrete Terraced, pine woods

219915/626942

450–500

800

B'ina

Limestone

3

Dwarf-shrub steppe and olives

G. HaMoreh Typic Rhodoxeralf

233961/724868

550

500

Bar Kokhba

Limestone

5

Rocky summit plateau, dwarf-shrub steppe

Ami'ad Typic Rhodoxeralf

249615/759281

600–650

400

Bar Kokhba

Limestone

15

Rocky, dwarf-shrub steppe and birch

Malkiyya Typic Rhodoxeralf

247880/778515

650–700

685

Bar Kokhba

Limestone

11

Graminae, formerly cultivated

Bar-Yohay Mt. 1 Typic Haploxerol

239382/766079

950

1150

Deir-Hanna

Soft dolomite

5

Bar-Yohay Mt. 2 Typic Rhodoxeralf

238916/766326

950

1120

Deir-Hanna

Soft dolomite

15

Hill slope, maquis

Miron roads fork Lithic Haploxerol

238681/766843

900–950

1130

Deir-Hanna

Dolomite

22

Maqis Organic-rich

Tsiv'on Typic Rhodoxeralf

239159/770247

850

750

Sakhnin

Dolomite

14

Above C horizon with dolomite debris Maqis

Gish Typic Rhodoxeralf Me'iliyya Lithic Rhodoxeralf

240771/769276

800–850

730

B'ina

Limestone

10

above C horizon with dolomite debris Dwarf-shrub steppe and oaks

222956/770071

750–800

485

B'ina

Limestone Limestone

15

214141/769175

600–650

80

B'ina

Limestone

8

Kabri Lithic vertic Rhodoxeralf

Summit plateau, maquis Organic-rich

Local depression, mixed vegetation Stony Stony Graminae, formerly cultivated Stony Stony

Abbreviations: M. = Ma'ale, Q. = Qiryat, G. = Giv'at.

3. Methods Samples of 1–1.5 kg were dried at 50 °C for several days. A dry sample was manually crushed and split to retain ~0.5 kg sample, which was further disaggregated, cleaned of visible organic particles and sieved through a 2 mm sieve (bulk sample). A portion of ~50 g was mechanically ground for bulk mineralogical and chemical composition. A portion of ~10 g was shaken with distilled water for 3 h, treated with ultrasonic rod centrifuged to separate the b2 μm fraction (clay fraction). The clay fraction suspension was saturated with Ca and pipetted on glass slides; the rest was dried and powdered for chemical analysis.

Randomly oriented powders of bulk samples and oriented preparations of the clay fraction were recorded using a Philips (PANanalytical) Xpert Pro diffractometer, generator PW3040, goniometer PW 3050/60 (theta/theta), equipped with an Xcelerator detector (Ni-filtered Cu-Kα radiation, 40 kV, 40 mA and step size of 0.017°2θ). The oriented-clay slides were analyzed in air-dried and ethylene glycol (EG) solvated states. In order to evaluate the dynamics of clay transformations in the soils studied the air-dried patterns were decomposed by the DECOMPXR program (Lanson, 1997) following the principles outlined by Velde et al. (2003) and Barré et al. (2007). Convergence of the fitting routine to 99% concordance was used for identifying the peak position and

A. Sandler et al. / Geoderma 239–240 (2015) 156–167

intensity. The principle peaks used for starting a program run were: SI (randomly ordered smectitic illite–smectite mixed layer minerals), IS (ordered illitic illite–smectite mixed layer minerals), chlorite, poorlycrystallized illite (PCI), well-crystallized illite (WCI) and kaolinite. A limiting factor used was that peak width at half maximum intensity (WHM) should not exceed 2°2θ. Indeed, being dependent on the coherent scattering domain size (CSDS), the larger WHM is, the lower is the CSDS. The 2°2θ is a threshold beyond which the diffraction intensity results from the contribution of several mineral phases. This constraint resulted in two SI phases in most samples. Kaolinite could have been decomposed into two peaks for most samples, but in cases that KaS was observed a third low-angle peak was needed to achieve 99% concordance. Therefore, for consistency, kaolinite was decomposed to three phases in all samples. The absolute area, as well as the percentage of the total area of each of the decomposed peaks, was calculated. Bulk samples and clay fractions were analyzed for their major element composition, and also for Sr, Ba and Zr, by ICP-OES (Perkin Elmer, OPTIMA 3300) after lithium metaborate (LiBO2) fusion. Each analysis run included repeated determinations of three of the international standards SO-3, BE-N, BHVO-1, SCo-1, NIM-L, and NIM-G. Continuous grain-size distribution was measured by laser diffraction (Malvern Mastersizer MS 2000) after dispersion using sodium hexametaphosphate, dissolution of organic matter by H2O2 and ultrasonic treatment (details in Crouvi et al., 2008).

159

eolian material in central Israel: unimodal dust samples collected at the Dead Sea (Singer et al., 2003), the fine silt mode of northern Negev loess (Haliva-Cohen et al., 2012), and is somewhat finer than the main mode of dust collected in Jerusalem (Haliva-Cohen et al., 2012). Such grain-size modes of b 20 μm in dust are assumed to derive mainly from remote sources. The 3–6 μm fine mode should be considered as a clay size, since laser-determined grains of b~ 5 μm are equivalent to the clay fraction determined by sedimentation (Pieri et al., 2006; Crouvi et al., 2008). Accordingly, the amounts of clay and silt fractions are fairly comparable. The small differences in grain size between topsoil and subsoil are not consistent, indicating ample soil turbation. The grain-size distribution of the Eocene and Galilee soils is rather similar but the Judea soils have a somewhat higher proportion of N20 μm population. As the areas studied are quite remote from the silt sources in the south the nature of the small population of fine sand fraction had to be evaluated. The material remaining on a 63 μm sieve of five samples was repeatedly treated with H2O2 and ultrasonically to increase disaggregation and sieved. After five such cycles the N 63 μm grains were examined under a scanning electron microscope and were found to be either organic debris or resistant clay aggregates. No individual mineral of N 63 μm grains could be detected.

4.2. Bulk mineralogical and chemical compositions 4. Results 4.1. Grain size Of the 27 samples that were analyzed for grain size (Fig. 2) twothirds were bimodal with modes at 3–6 μm and 10–35 μm and the other third was unimodal as the coarse population was low and appeared as a shoulder. An additional mode at around 0.7 μm appears as a shoulder in many samples. The coarse mode is similar to that of

561A

3-6 mode

6

% Volume

563C

15-40 mode

5

562B 563

3

565A

2

565B 565C

1

B

567A 0.1

1

Grain size µm

10

576A 576B

4

576C 577A

3

578A

2

579A 580A

1 0

0.1

1

Grain size µm

10

100

571A 571B

0.7 mode

3

568B

574B 10-26 mode

570B 11-25 mode

4

568A

574A

570A

3-5 mode

6 5

% Volume

100

3-5 mode

6 5

562A

4

0

C

561B

% Volume

A

All samples are predominated by similar amounts of quartz and phyllosilicates (actually clays, as micas are rather rare), with trace amounts of K-feldspar and plagioclase (Suppl. 1). This is reflected as well in the comparable silt- (mainly quartz) and clay- (mainly phyllosilicates) size fractions. Minor or trace amounts of calcite, dolomite and goethite occur in many samples. Trace amounts of dolomite in profiles based on limestone originate from dust (e.g. SA 568, 578), either remote or adjacent. Trace amounts of calcite in profiles based

572A 572B

2 1 0 0.1

1

Grain size µm

10

100

Fig. 2. Grain-size distribution of soil samples. A. Judea transect. The main mode is at 3–6 μm and a secondary mode is at 15–40 μm. B. Eocene transect. The main mode is at 3–5 μm and a secondary mode is at 11–23 μm. C. Galilee transect. The main mode is at 3–5 μm and the secondary mode, which is mostly a shoulder, is at 10–26 μm, except one deep subsoil sample with a coarser mode. Most samples have a small finer mode at ~0.7 μm.

160

A. Sandler et al. / Geoderma 239–240 (2015) 156–167

on dolomite originate either from dust, or from limited pedogenic calcite, commonly formed around roots (see Figs. 4.2.3-4 in Singer, 2007). The bulk chemical composition of soils (Suppl. 2) with a dominance of silica (SiO2) and alumina (Al2O3) confirms the predominance of quartz and phyllosilicates. The CaO concentration, which is on average 3.5% for soil samples and 20% in dust (Suppl. 2), indicates significant carbonate leaching. Accordingly, CaO is negatively correlated (R2 = 0.73) with alumina (Fig. 3). The plots of elements/oxides versus Al2O3 were grouped by three rainfall categories: 450–600, 625–750 and 775–950 mm (Fig. 3). However, no distinct relations to rainfall could be recognized. Iron oxide displayed the best correlation to alumina (R2 = 0.91) and a single linear array for dust and soil samples, indicating a common source for both oxides in dust and stability in the soils studied. Titanium and manganese oxides are also positively correlated with alumina but with lower coefficients (R2 = 0.66 and 0.62 for soil samples, not shown). Silica is only vaguely correlated with alumina as a result of several factors: variable quartz to phyllosilicate ratios in dust (higher in Judea soils), solubility and leaching of silica (10–20 mg/l in groundwater), and pedogenic processes outlined below. Potassium oxide is totally scattered relative to alumina and in many soil samples is depleted relative to dust (Fig. 3). However, in some samples of the high rainfall category potassium depletion is not recognized, indicating complex behavior in soils, related to plant activity. Zirconium and barium are also scattered relative to alumina. Relative to dust zirconium is enriched in most soils whereas barium is depleted. Barium depletion could be related to calcite dissolution, but as no correlation between these two elements is observed (not shown) it seems that their behavior in soils is independent. This is also true for strontium, which displays the largest depletion in soils relative to dust among all analyzed elements. However, Sr/CaO and Ba/CaO ratios are positively well correlated with alumina (Fig. 3), indicating that both

strontium and barium are enriched in soils, relative to calcium, along with soil leaching and enrichment of alumina. The Ba/Sr ratio, which is an indicator for leaching in soils (e.g. Sheldon and Tabor, 2009), is 2.5 to 4.5 times higher in soils than in dust. The ratios of elements/oxides to alumina for topsoil and subsoil samples are plotted against rainfall, the main agent of soil leaching (Fig. 4, deep samples are excluded). The CaO/Al2O3 ratio displays similar values of in topsoil and subsoil of a profile and is rather uniform along the rainfall range, reflecting intensive soil leaching for most samples. The Fe2O3/ Al2O3 ratio is also uniform along the rainfall range and so are the ratios of magnesium, manganese and titanium oxides to alumina (not shown). A slight decrease at highest rainfall is observed for the latter ratio, indicating that in the soils studied titanium is less stable than aluminum. The zirconium to alumina ratio decreases three-fold with rainfall indicating that zirconium, like titanium, is less immobile than aluminum in these soils. Barium and strontium ratios to alumina display a similar trend of a two-fold decrease as rainfall increases, though at the highest rainfall a slight increase is observed. Silica and potassium oxide ratios to alumina, however, display an exceptional trend of decrease along rainfall up to 850 mm, but a marked increase above 900 mm apparently related to intensive plant activity. 4.3. Clay fraction chemical composition Some of the oxides/elements are depleted in the clay fraction relative to the bulk chemical composition, both having similar loss on ignition (LOI, Suppl. 2). This is mainly due to diminishing of minerals that typically prevail in the silt fraction. Thus SiO2, CaO, Na2O, TiO2, and Zr are depleted due to reduction in quartz, calcite, plagioclase and heavy minerals, respectively. Alumina is consequently enriched in the clay

25

15 10

8

SiO2

Fe2O3

CaO

60

10

20

50 40

6 5

30

4

0 7

12

17

7

22

12

17

7

22

12

17

22

Al2O3

Al2O3

Al2O3

380

2.2

310

1.7

Ba

440 Zr

K2O

590

1.2

240

290

170

140

0.7 7

12

17

7

22

12

17

Sr

350 200

26

9

21

7

16 11

1 17 Al2O3

22

22

17

22

5 3

6 50

17 Al2O3

Sr/Ca*103

Ba/Ca*103

500

12

12

Al2O3

Al2O3

7

7

22

1 7

12

17 Al2O3

22

7

12 Al2O3

Fig. 3. Bulk chemical composition of soils: plots of selected major oxides and trace elements versus alumina. Ba/Ca and Sr/Ca ratios versus alumina are presented as well. The data are grouped by three rainfall ranges that display large overlaps. Five dust samples are presented for comparison.

A. Sandler et al. / Geoderma 239–240 (2015) 156–167

161

1.2 0.8 0.4 0

0.5

550 650 750 850 950 annual precipitation mm

450

450

950

0.08

22

12

550 650 750 850 950 annual precipitation mm

550 650 750 850 annual precipitation mm

950

550 650 750 850 annual precipitation mm

950

13 Sr/Al2O3*103

Ba/Al2O3*103

0.12

0.04 450

550 650 750 850 annual precipitation mm

32

0.16

3.5

2.5

0.45 450

K2O/Al2O3

4.5 0.55

SiO2/Al2O3

Fe2O3/Al2O3

CaO/Al2O3

1.6

8

3 450

550 650 750 850 annual precipitation mm

950

450

Zr/Al2O3*103

40

25

10 450

550 650 750 850 950 annual precipitation mm

Fig. 4. Bulk chemical composition of soils: plots of selected major oxides and trace elements to alumina ratios versus rainfall. CaO and Fe2O3 to alumina ratios display similar values along the rainfall range. K2O and SiO2 display gradual decrease with rainfall but a sharp increase at N800 mm rainfall. A small peak for K2O at 675–700 mm is of the highest rainfall in the Judea (a) and Eocene (b) transects. Ba and Sr display moderate decreases with rainfall whereas Zr displays a prominent decrease.

fraction, whereas some elements do not display significant differences relative to bulk composition. The relations to alumina of selected oxides/elements are presented again for the clay fraction (Fig. 5). The plots display weak correlations with alumina of most oxides/elements, though iron and titanium oxides (the latter is not presented) are positively correlated. These correlations indicate that iron and titanium are preserved in the soil along with alumina, regardless of the clay mineral types. Calcium and magnesium oxides (the latter is not presented) decrease relative to alumina due to decrease in carbonate minerals and increase in clay minerals. Both are negatively correlated with alumina, though in a less pronounced manner than in the bulk composition. Strontium and Ba, however, do not follow calcium and magnesium oxides, as their absorption to clays is more significant in the clay fraction. Rainfall has a weak or no effect on the relations of most oxides/elements to alumina (Fig. 5) and on the concentrations normalized to alumina (not shown). However, it has a unique effect again for silica and potassium oxides that slightly decrease between 650–850 mm but sharply increase at N900 mm (Fig. 6). This indicates that the behavior of silica and potassium oxide in the bulk composition is dictated by the clay fraction, which responds to K and Si availability. 4.4. Clay fraction mineralogical composition Diffractograms of glycolated specimens indicate that IS and kaolinite predominate in all samples, IS being the major mineral in most samples. Illite is a minor mineral but its content varies considerably, whereas chlorite is a trace mineral in most samples. Small amounts of mixedlayered kaolinite–smectite (KaS), goethite and quartz are frequently observed. Three selected parameters were derived from the values calculated by the decomposition program and used to follow clay pedogenic

processes: a. the relative peak area (to percentage of the total area of clay phases) of kaolinite (including KaS); b. the center of gravity (cg) parameter (Barré et al., 2007), which estimates the total state of illitization of all 2:1 phases; and c. “IS”, the sum of the relative areas of WCI, PCI and IS. The parameters were calculated for 28 soil samples, three dust samples and one loess sample (Table 2). The ranges of values for the three parameters were 4–58 for kaolinite, 5.2–7.5 for cg, and 0–59 for “IS”. The decomposition principles are first demonstrated on a dust sample from Jerusalem (Fig. 7). The PCI peak includes also palygorskite, known to be present in dust and loess, but not in leached soils (Velde, 1985 p. 152; Sandler, 2013). Compared to dust, soil parameters display the following trends (Figs. 8, 9, 10, 11, Table 2): a. kaolinite decreases under rainfall below 600 mm but strongly increases above 625 mm in the Eocene transect and above 800 mm in the Galilee transect. The decrease is related to the relative increase in SI phases, which are formed under limited drainage due to either low rainfall or to less karstic bedrocks as the Turonian limestone in the Galilee transect; b. SI formation is also reflected in the cg parameter, which represents total illitization, and displays a similar trend to that of kaolinite: it decreases under low rainfall and on soils above Turonian limestone, and rapidly increases with rainfall in the Eocene and Galilee transects, above 600 and 800 mm, respectively. Illitization processes may result in pedogenic formation of IS and illite phases, represented by the “IS” parameter. It reaches highest values in the soil under the highest rainfall (675 mm) of the Eocene transect and in the soil under the highest rainfall of the Galilee transect (950 mm). Yet, a nearby soil does not display the same degree of illitization as it is much thicker. Hence, bedrock lithology, rainfall and local conditions control soil formation of kaolinite, SI and illitic phases. This is best demonstrated for soils developed on Turonian limestone, which are more smectitic than those developed on Cenomanian dolomite or Eocene limestone. Still, the three Turonian

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12

2.5

1.0

48 SiO2

Fe2O3

CaO

4.0

10

8 17

20

23

17

20

42

23

17

20

Al2O3

Al2O3

2.5

330

250

1.9

23 Al2O3

280 Ba

Zr

K2O

45

200 1.4

230 150

0.8 17

20

180 17

23

20

23

17

20

Al2O3

Al2O3

23 Al2O3

Sr

105

80

55 17

20

23 Al2O3

Fig. 5. Clay fraction chemical composition of soils: plots of selected major oxides and trace elements versus alumina. The data are grouped by three rainfall ranges that display large overlaps.

soils of the Galilee transect (# 578, 579, 580) display the same trends of parameter decrease with rainfall increase. 5. Discussion 5.1. Implications for parent materials Previous studies argued either for the insoluble residue of carbonate bedrocks, or for both dust and insoluble residue, as the parent material of terra rossa and related soils, as mentioned in the Introduction. The second option is accepted today by soil scientists of the Mediterranean region, but the exact contribution of each source has not yet been quantified. The grain size and chemical composition results of the current

study confirm the importance of dust, apparently its dominance, as a parent material. The coarse mode of the grain size population, namely 11–40 μm, is similar to that of published analyses of settled dust in central Israel. The coarser mode of the Judea soils and their higher quartz content are in accord with proximity to the desert belt, from where dust originates. Moreover, the preliminary examination of the sand fraction indicates absence of sand quartz grains, though such grains often occur in Cretaceous carbonate rocks. The chemical composition, as reflected in relations between various oxides and discussed below, also indicates that the soils studied are well-related to dust composition. Clay fraction mineralogical composition of dust in central Israel is rather uniform due to homogenization of various desert and local sources (Sandler, 2013).

0.15

K2O/Al2O3

SiO2/Al2O3

2.6

2.3

0.11

0.07

2.0

1.7

0.03 450

550

650

750

850

annual precipitation mm

950

450

550

650

750

850

950

annual precipitation mm

Fig. 6. Clay fraction chemical composition of soils: plots of K2O and SiO2 to alumina ratios versus rainfall. Like in the bulk composition (Fig. 4) a sharp increase is observed at N800 mm rainfall.

A. Sandler et al. / Geoderma 239–240 (2015) 156–167 Table 2 Three quantitative parameters that were calculated from decomposed X-ray patterns for soil, dust and loess samples. The parameters are presented along sample number, depth, rainfall and elevation as in Table 1. Sample 2* 3* 4* Loess 561a b c 562a b 563a 565a b c 567a b 568a b 570a b 571a b 572a b 574a b 576a b c 577a 578a 579a 580a

Depth

0–5 5–15 20–40 0–10 15–30 0–10 0–10 15–25 30–45 0–10 15–30 0–10 15–30 0–10 15–30 0–10 15–30 0–10 15–30 0–10 15–30 0–5 10–25 30–45 0–10 0–10 0–5 0–5

Rainfall mm

Elevation asl m

700 700 700 675 675 625 600 600 600 500 500 475 475 550 550 625 625 675 675 950 950 925 925 925 850 825 775 625

795

705 835 325

410 800 500 400 685 1150 1120

750 730 485 80

% Ka

cg

% “IS”

19 17 19 13 26 20 15 15 13 15 7.6 7.5 6.8 5.4 5.3 11 12 9.4 8.8 46 44 58 49 42 40 20 17 23 41 19 7.3 4.0

6.11 6.01 6.05 6.64 5.64 5.56 5.50 5.18 5.20 5.41 5.30 5.42 5.48 5.46 5.56 5.45 5.43 5.35 5.36 5.63 5.53 7.49 6.91 6.93 6.29 5.75 5.79 5.86 6.19 5.65 5.57 5.58

37.4 30.2 24.3 42.1 4.4 2.1 2.1 3.4 0.7 1.15 0.5 0.2 0.1 0.1 0 1.7 1.4 1.1 0.8 8.7 5.2 58.6 46.1 39.3 12.4 4.9 3.8 7.4 19.0 0.5 0.4 0

A large variability of mineralogical composition is demonstrated, however, for the clay fraction of the soils studied; all that have been developed under Mediterranean climate regimes, are shallow and based on karstic carbonate rocks. Two mineralogical end member assemblages are observed: one with dominance of SI and low values of kaolinite, cg, and “IS” parameters; and the second with a low SI amount and high values of the three parameters. The most suspicious trend is the positive relation between the kaolinite and the “IS” parameters (Fig. 11). Variable assemblages in between the two end members reflect a complex situation with several factors acting simultaneously. Yet, this variability could not be related either to bedrock or dust composition, which are both parent materials of the fine fraction. Ka 2 SI 2 IS

SI 1

Ka 1

Ch PCI + P WCI

KaS

Fig. 7. Decomposed X-ray patterns of the oriented air-dried clay fraction of dust in Jerusalem. The original diffractogram is in blue, the primary phases are in gray and their combined plot is in violet. This diffractogram exhibits all clay phases that were considered in the soils studied. The red star marks the position of the center of gravity (cg). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

163

The difference between the soils' mineralogical composition and that of the bedrocks is best demonstrated for the soils based on the Eocene limestone (Bar Kokhba Fm.). The insoluble fraction of this formation is dominated by smectite and clinoptilolite, minor amounts of palygorskite and no kaolinite or illite (Nathan and Flexer, 1977; Sandler, unpublished data), whereas the overlying soils contain considerable kaolinite and illite, hardly any smectite, and no clinoptilolite or palygorskite. Similarly, soils developed on Turonian limestone are almost devoid of illite and illitic IS phases, which are common in this limestone (Sandler and Harlavan, 2006; Sandler and Saar, 2007). The difference between soil and dust clays is generally more subtle, but could already be recognized in soils under rainfall of 475 mm, the lowest in the current study. Under the rainfall range of 475 to around 675 mm formation of pedogenic SI phases is the dominant change, in accord with the common knowledge on the formation of pedogenic smectite (Weaver, 1989 p. 153; Wislon, 1999). The relative values of kaolinite, cg and “IS” parameters (Fig. 10) decrease correspondingly. However, above 625 mm for the Eocene and Judea transects and above 800 mm for the Galilee transect all three parameters increase, reach dust values or exceed them. On the intensively karstified Eocene limestone the increase in these three parameters ends up with a clay assemblage that is totally different from dust as well. Hence, it turns out that the principal agent for clay pedogenesis is leaching intensity, which basically depends on rainfall and soil/bedrock permeability. On such hard carbonate rocks permeability is dictated by bedrock karstification progression and also by soil thickness. In general, the better soils are drained, the higher the amounts of both kaolinite and illitic phases. This is well demonstrated by the main difference between the Turonian and the Eocene limestone formations under similar rainfall. The soils on Eocene limestone with the highest karst progression, as observed on the surface and vertical road-cuts, have higher values of the three clay parameters than corresponding soils on the Turonian limestone. Moreover, where permeability is high, the effect of rainfall is enhanced, as can be seen from the steep upsurge of the three parameters where rainfall on the Eocene soils increases by only ~ 60 mm (# 572 versus 571). A similar relation to permeability has also been observed in permeable sandy soils in the central coastal plain of Israel (Sandler, 2013) where a slight increase in rainfall significantly changed the clay mineralogy. 5.2. Effect of pedogenesis on the chemical composition Formation of soil illite and illitic IS along with kaolinite requires that parent 2:1 clay minerals are dissolved to supply Si, Al, and some K, Fe and Mg; the latter is also supplied by dissolution of calcite and dolomite. The main reason for the instability of the detrital 2:1 clays, which are either derived from desert dust or from the insoluble residue of the bedrock marine carbonate rocks, is apparently reduced pH from N8 in sources to ~ 7.5 or less in soil environments (Suppl. 1). Aluminum seems to be rather immobile in the soils studied, as demonstrated by iron oxide to alumina ratios in bulk samples (0.46–0.57), clay fraction (0.46–0.57), and dust (0.48–0.56). The limited range for the three groups suggests that pedogenically formed clays utilize most or all available alumina and iron. Iron released during clay transformation is precipitated as clay-size oxyhydroxides and hematite, the main coloration agent in red-brown Mediterranean soils and is in part incorporated into the soil clay structure. Thus, weathering of clays is a significant agent in hematite formation in red-brown Mediterranean soils, as has been suggested before for soils in the Province of Cordoba, Spain (Torrent and Cabedo, 1986). This is supported here by the very low concentrations of iron and aluminum in the underlying mountain aquifers, which are b 5 and b10 μg/l, respectively, compared to silica, which is in a range of 10–20 mg/l. Aluminum immobility is also evidenced by the relative depletion of Zr (Fig. 4), which is generally considered as an immobile element in soils and weathering profiles. In certain cases mobility of zirconium in

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B

A

2

4

6

8

10

12

14

4

6

8

10

12

14

2

4

6

8

10

12

14

C

2

Fig. 8. Decomposed X-ray patterns of oriented air-dried clay fraction of the Eocene transect soils: A. topsoil under highest rainfall (#572a); B. topsoil under intermediate rainfall (#571a); C. topsoil under lowest rainfall (#570a). Note disappearance of SI phases and increase of IS phases relative to kaolinite with rainfall.

soils was observed and attributed to organic complexes, along with Al (e.g. Kurtz et al., 2000), or to the preferential dissolution of radiationdamaged domains (e.g. Balan et al., 2001). In the soils studied here Zr is considerably more mobile than Al, and the organic matter is rather low (Suppl. 1) compared to tropical soils. Hence, radiation damage rather than organic complexation is the crucial factor for Zr mobility. This behavior is not directly related to clay neoformation in the soils studied but should be taken in account while estimating weathering processes in soils affected by Saharan dust, the main source of these zircons.

Potassium is retarded in the soil environment due to its affinity to clays. It is rather mobile in soils, relative to Al and Fe, but much less than Na, Ca and Mg. However, total, or soluble, or exchangeable potassium are known to be enriched in the upper part of various soil types in Israel (e.g. Dan and Yaalon, 1982; Lado et al., 2012) and elsewhere (e.g. Salem and Hole, 1973; Jobbagy and Jackson, 2001; Simonsson et al., 2007). Both potassium and silicon are intensively circulated by vegetation from subsoil to topsoil (Derry et al., 2005; Barré et al., 2009) and hence their mutual increase under high rainfall (Figs. 4, 6)

A

B

C

D

Fig. 9. Decomposed X-ray patterns of oriented air-dried clay fraction of the Galilee transect soils: A. topsoil of a thin soil under the highest rainfall (#574a); B. topsoil of a thick soil under the highest rainfall (#576a); C. topsoil based on well-developed karstic bedrock under high rainfall (#577a); D. topsoil under low rainfall based on Turonian limestone (#579a). Note disappearance of SI and increase of IS phases in A and C and maximum SI formation and IS disappearance in D.

A. Sandler et al. / Geoderma 239–240 (2015) 156–167

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7.5

60

7 6.5

cg

Ka

40

20

6 5.5

0

5 400

500

600

700

800

900

annual precipitation mm

400

500

600

700

800

900

annual precipitation mm

55

"IS"

35 15 -5 400

500

600

700

800

900

annual precipitation mm Fig. 10. Plots of clay parameters, quantitatively calculated from decomposed XRD patterns of the oriented air-dried clay fraction, versus rainfall. Dust data for comparison are arbitrarily presented at 425 mm. Note general increase in the three parameters with rainfall, especially the sharp increase in the Eocene transect soils.

suggests that vegetation productivity reached a point where potassium and silicon addition overcame their leaching. 5.3. Pedogenic illite Formation of illite has been suggested before in arid and semi-arid soils in places such as Australia (Nettleton et al., 1973), Israel (Singer, 1989) and China (Huang et al., 2011). However, those occurrences have been questioned: the soil illite in Australia was suggested to be derived from the lacustrine bedrock (Wislon, 1999), loess in Israel displayed some illitization of IS phases, but discrete illite was lower than in dust (Sandler, 2013) and illite in Chinese loessial soils has been argued to be detrital (Ji et al., 1999; Zhao et al., 2005). Direct evidence for illite precipitation on K-feldspar (Eggleton and Buseck, 1980; Velde and Meunier, 2008), or even on plagioclase (Bétard et al., 2009) in weathered crystalline rocks, has been documented only on a microscopic scale. In both microenvironments leaching was limited and potassium was locally enriched.

55

"IS"

35

15

-5 0

15

30

45

60

Ka Fig. 11. A plot of the kaolinite versus “IS” parameter. Note positive correlation along all values with two increments: a gentle slope at low values and a steep slope at high values.

Topsoils of temperate climate regions in the USA were also documented to be enriched in illite, suggested to have formed by potassium and ammonium fixation into vermiculites and mica-vermiculates. The nutrient enrichment was caused by translocation by vegetation (e.g. Velde, 2001; Graham and O'Geen, 2010). Only very few studies have indicated considerable illite formation in leached soils of the temperate or tropical regimes. Two studies demonstrated it in basaltic soils (Juang and Uehera, 1968; He et al., 2008) and one in terra rossa (Torrent and Cabedo, 1986). In addition, a few studies reported specific plants that promoted some illite formation in soils (e.g. Madsen and Nornberg, 1995; Tice et al., 1996). A simple biogeochemical model linked plant primary productivity and clay mineral thermodynamic stability and explained how plant activity might stabilize clay minerals in the upper part of temperate grassland soils. The outputs of the model suggested that plant activity under certain conditions, including higher potassium translocation, may well stabilize illitic clay in surface soils, whereas higher silicon translocation may well stabilize Ca-montmorillonite (Barré et al., 2009). In the current study maximum evolution of illite along with kaolinite and accompanied by diminishing SI phases was manifested in two different environments: a thin soil (# 574) based on faintly karstified bedrock under the highest rainfall, and a soil based on intensively karstified limestone under moderate rainfall (# 572). Both soils also display the largest mineralogical differences between topsoil and subsoil (Table 2, Fig. 10), indicating that kaolinite and illitic phases form mainly at the surface. It is well acknowledged that pedogenic formation of kaolinite is favored over smectite under higher rainfall and improved leaching (Weaver, 1989 pp. 152–153; Wislon, 1999). However, the mutual increase in illitic phases along with kaolinite (Fig. 11) is in contrast to the common knowledge on the occurrence of illite in soils (Weaver, 1989 pp. 176–181; Wislon, 1999). Theoretically, this soil illite could be inherited from dust and only relatively increase along with the dissolution of SI and IS phases of parent materials. However, the following arguments negate inheritance: 1) pedogenic dissolution of original SI and IS phases would equally increase the amounts of illite and kaolinite in soils, or kaolinite could even increase more due to neoformation. However, it has recently been shown that the amount of kaolinite in Israeli leached soils was

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equal to, or somewhat higher than in dust, whereas illite could reach twice to three-fold higher amounts (Sandler, 2013); 2) the quantitative relative values of kaolinite, cg and “IS” parameters decrease from dust to Mediterranean soils under medium rainfall and increase again in higher rainfall and leaching. If more rainfall and leaching would preferentially preserve illite and kaolinite in soils then their relation to rainfall would produce a linear curve rather than a U-shape curve (Fig. 10); 3) As illite and illitic IS vanish under reduced leaching, especially in soils based on Turonian limestone (#567b, #581) where pedogenic SI is formed, then intensive leaching should definitely reduce their stability. Under the highest rainfall (Mt. Meron summit, #574) the formation of illitic phases along with kaolinite could be related to the impact of rich vegetation. The region vegetation is characterized by dense population of both trees, such as oak and pistachio species, and dense seasonal weeds and grasses, prevailing longer here due to the highest rainfall and coolest weather. The vegetation provides K and Si to the soil solution to form a favorable environment for illitization of SI phases and neoformation of soil illite. This may continue as long as carbonate dissolution keeps pH rather high. Indeed, all soil samples yielded pH (1:1) values in the range of 6.9 to 7.8. This situation resembles the unique case in Hawaii, where soils, based on basaltic rocks containing no micaceous phases originally, produce more illite at higher elevation and rainfall (Juang and Uehera, 1968). However, high rainfall and rich vegetation are not sufficient for soil illitization, as the nearby soil (#575) with the same rainfall and vegetation is much less illitized. That soil has a much thicker profile that restricts drainage, hence SI formation is favorable (Fig. 10). It might be that here, like in soils based on the Turonian limestone, Si concentration dominates the soil solution and smectite formation is favored over illite, as has been theoretically calculated in the model of Barré et al. (2009). A remarkable situation of illite formation is evident in the soil based on the intensively karstified Eocene formation, where rainfall is medium and vegetation is not rich. Documentation shows that this soil has the highest non-exchangeable potassium and ammonium among 31 soil profiles in Israel (Feigin and Yaalon, 1974). Vegetation on this formation is dominated by Bulbous Barley (Hordeum bulbosum) throughout the Galilee (Rabinovitch-Wein, 1986), but substantial illite formation occurs only at maximum rainfall (on this formation). This suggests enhanced translocation of potassium and ammonium by that specific vegetation. Indeed, it has been found that grasses are generally a more efficient agent of nutrient translocations than bushes and trees, and certain grasses are more efficient than others (Barré and Velde, unpublished results). Here, however, further study of soil solution composition in key sites is needed to better constrain the parameters involved in illite formation in terra rossa and related soils. 5.4. Implications for duration of pedogenic processes The relations between chemical constituents, besides those related to carbonate minerals (Fig. 3), as well as grain-size distribution suggest dust as a major parent material of the studied soils. Though mineralogical changes occur mainly in top soils, the similar mineralogical and chemical compositions of topsoil and subsoil along most profiles indicate that the continuously added airborne material is transformed all along the profile, as these soils are rather shallow. As local leaching conditions dictate mineralogical and chemical variations between soils, even between closely-spaced sites, both mineralogical and chemical compositions reflect the impact of current environmental conditions. Soils were formed on the current landscape long ago, probably intensively during glacial periods (Horowitz, 1979). Since its formation the soil material was partially preserved, partially eroded, and partially locally resedimented. However, the mineralogical and chemical variability indicates that each site reflects steady-state equilibrium between dust accretion, climatic conditions and landscape. As all these parameters are known to have been rather stable since the Holocene, and as

short-term climatic variations within the Holocene were too short to leave their impact on soils, it can be presumed that this apparent steady-state equilibrium was achieved within the last few thousand years. 6. Conclusions The relations between chemical elements, as well as grain-size distribution, suggest dust as a major parent material of the studied terra rossa and related shallow mountain soils. Quantitative relative analysis of clay minerals reveals large variability. Clay formation in the soil environment is a dominant process that conceals inheritance from bedrock or dust. Consequently, clay neoformation releases iron and promotes hematite production, the main coloration agent in red-brown Mediterranean soils. The principal agent for clays pedogenesis is leaching intensity, which basically depends on rainfall and soil/bedrock permeability. Permeability is dictated by bedrock karstification progression and also by soil thickness. In general, the better soils are drained the higher is the amounts of both kaolinite and illitic phases. Soil clay assemblages lay between two mineralogical end members: one with dominance of SI and low values of kaolinite and illitization parameters, and the second with low SI amounts and high values of kaolinite and illitization parameters. Potassium and silicon are rather mobile relative to aluminum, but at the highest rainfall they are enriched due to vegetation translocation. Enrichment of potassium occurs also in highly leached soils with certain vegetation and promotes illite formation at the expense of original IS phases. However, the exact environmental conditions and the specific plants that shift soil solution towards stability of illite are still to be determined. The response of the clay mineral assemblages to the current environmental conditions suggests that the apparent steady-state equilibrium with the environment was achieved within the last few thousand years. Acknowledgments Thanks are due to C. Fontaine (HydrASA) for the assistance in XRD analyses, to D. Stiber, O. Yoffe and R. Binstock (GSI) for the chemical analyses, to O. Crouvi (GSI) for the grain-size analysis, and to R. Calvo and C. Netzer for the slope calculation and preparation of maps. Two anonymous reviewers are thanked for useful comments. Appendix A. Supplementary data Sampling sites: red points = a section; yellow points = b section; light blue = c section; black pins = dust and loess (the southernmost site). Supplementary data associated with this article can be found in the online version, at http://dx.doi.org/10.1016/j.geoderma.2014.10.008. These data include Google map of the most important areas described in this article. References Balan, E., Neuville, D.R., Trocellier, P., Fritsch, E., Muller, J.-P., Calas, G., 2001. Metamictization and chemical durability of detrital zircon. Am. Mineral. 86, 1025–1033. Barré, P., Velde, B., Catel, N., Abbadie, L., 2007. Soil–plant potassium transfer: impact of plant activity on clay minerals as seen from X-ray diffraction. Plant and Soil 292, 137–146. Barré, P., Berger, G., Velde, B., 2009. How element translocation by plants may stabilize illitic clays in the surface of temperate soils. Geoderma 151, 22–30. Bétard, F., Caner, L., Gunnell, Y., Bourgeon, G., 2009. Illite neoformation in plagioclase during weathering: evidence from semi-arid Northeast Brazil. Geoderma 152, 3–62. Boero, W., Premoli, A., Melis, P., Barberis, E., Arduino, E., 1992. Influence of climate on the iron oxide mineralogy of terra rossa. Clay Clay Miner. 40, 8–13. Crouvi, O., Amit, R., Enzel, Y., Porat, N., Sandler, A., 2008. Sand dunes as a major proximal dust source for late Pleistocene loess in the Negev Desert, Israel. Quat. Res. 70, 275–282. Dan, J., 1990. Effect of dust deposition on the land of Israel. Quat. Int. 107–111.

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