Characterization of clay minerals in a dystrochrept

N. Jb. Miner. Mh.

Jg. 2002 (12)

551–576

Stuttgart, Dezember 2002

Characterization of clay minerals in a dystrochrept developed in a gneiss (Dora-Maira Massif, Western Alps). Complementarity between X-ray diffraction and transmission electron microscopy Bruno Alessandria, Dino Aquilano, Marc Amouric and Olivier Grauby With 13 figures and 3 tables

Alessandria, B., Aquilano, D., Amouric, M. & Grauby, O. (2002): Characterization of clay minerals in a dystrochrept developed in a gneiss (Dora-Maira Massif, Western Alps). Complementarity between X-ray diffraction and transmission electron microscopy. – N. Jb. Miner. Mh. 2002 (12): 551– 576; Stuttgart. Abstract: A single soil (dystrochrept) profile developed from weathered gneisses (Gambasca, Po valley, Dora-Maira Massif, SW Alps), was characterized by petrographic observations, X-ray diffraction (XRD) and high resolution transmission electron microscopy (HRTEM) coupled with microchemical analyses. Di- and trioctahedral micas, corresponding to the white mica and to the biotite relics found in petrographic sections, are phengite and more or less K-altered biotite observed by HRTEM, while a third 10 Å phase, with illitic chemical composition, corresponds to the material surrounding primary micas. Kaolinite and halloysite were commonly observed by HRTEM: free or lying in contact with biotite and phengite (kaolinite) and as hollow tubes or spirals (halloysite). A few chlorite particles were found intimately mixed with biotite (Mg – chlorite) and phengite (Al – , Mg – chlorite). XRD diagram decomposition shows disordered and ordered (1 : 1) interlayered mica – vermiculite whilst HRTEM data established different vermiculitization stages of biotite. Key words: HRTEM, phyllosilicates, soil, weathering, XRD.

DOI: 10.1127/0028-3649/2002/2002-0551

0028-3649/ 02/2002-0551 $ 6.50

ã 2002 E. Schweizerbart’sche Verlagsbuchhandlung, D-70176 Stuttgart

552

B. Alessandria et al.

Introduction The characterization of soil phyllosilicates, and especially of those of the clay fraction, is generally not a simple task and requires more than one analytical approach. In fact one has often to deal with complex mixtures of clay mineral phases, including mixed-layered clay minerals, showing different coherent scattering domain size (CSDS). Also, among these phases, several can exhibit similar main structural characteristics, e.g. the same d(001), making ambiguous the identification. The resulting complex X-ray diffraction (XRD) patterns of these mixtures can be analyzed by means of profile decomposition programs (DiffracAt by Siemens 1991, Decompxr by Lanson, 1992) which greatly improve the identification of different clay phases (Lanson 1997). Some sound attempts have been already made on clay assemblages to validate the decomposition procedure, which is a mathematical treatment of the XRD pattern: – Righi & Meunier (1991) characterized the clays of an acid brown soil (dystrochrept); – Lanson & Besson (1992), Lanson & Velde (1992) characterized the smectite to illite transformation for a diagenetic series with illite – smectite (IS) interlayered clays; – Righi et al. (1993) studied hydroxy-interlayered vermiculite (HIV) and IS interlayered minerals of a cryorthod; – a more complex case study was recently made by Righi & Elsass (1996) who performed the decomposition of XRD patterns obtained on soil clay fraction. High resolution transmission electron microscopy (HRTEM), along with analytical electron microscopy (AEM) analysis, is the powerful technique useful to characterize clay minerals, as proved by Banfield & Eggleton (1990), Elsass & Robert (1992), Romero et al. (1992), Aoudjit (1993), Aoudjit et al. (1995) and Banfield & Barker (1998). Righi & Elsass (1996) compared, from decomposed XRD patterns and HRTEM observations, the information obtained on fine soil clay fraction and concluded that the coupling of the two methods gave a realistic identification of the clay phases present in a complex soil sample, as well as the direct measurement of limited number of clay particles on HRTEM images. In this paper, one soil was chosen as a typical example of acid soil developed from a colluvium formed by alteration of both a in situ gneiss (substrate) and fallen milonitic gneisses (upper level rocks).

Characterization of clay minerals

553

The phyllosilicates of this soil constitute a complex system which can be schematically represented by: – relics of primary di- and trioctahedral micas, – weathering products derived from the transformation of primary micas, – neoformed clay minerals. Our aim was at identifying all phyllosilicate phases by means of two complementary techniques, XRD and HRTEM, very scarcely coupled in soil science until now. The former approach gives an averaged statistical information and the latter, especially when applied to undisturbed whole soil samples, allows the solution of texture problems (as are the topological relationships in the solid-state transformations) and provides information about certain details, impossible only by XRD study.

Soil material The soil, a dystrochrept (U.S. Department of Agriculture – USDA – 1986) has been sampled in Gambasca (Po Valley, Western Alps, Italy) at an elevation of 520 m. It developed from a colluvium (with a slope of 20 %) and is representative of all profiles lying on the right side of the valley. The substrate rock is an alteritic gneiss while the upper level rocks (at an elevation of 800 m) are dark milonitic gneisses. All soil horizons, except the A1, were sampled at two levels (Table 1). Each sample was fractionated by sedimentation into four grain size classes: coarse sand (0.2 – 2 mm), fine sand (0.02 – 0.2 mm), silt (0.002 – 0.02 mm) and clays ( £ 2 mm), as shown in Fig. 1. The main characteristics of the soil horizons are listed in Table 1; a short description of the profile is given below. 0 – 20 cm: A1, dark yellowish brown (10 YR 3/5, Munsell Soil Color Charts, Table 1. Characteristics of the soil samples. Depth (cm)

Horizon

pH (H2O)

15 35 55 85 115

A1 B B C C

4.3 4.3 4.4 4.8 5.0

Clay fraction (g Kg – 1)

Organic carbon (g Kg – 1)

CEC (cmolc Kg – 1)

10.1 21.2 27.3 22.3 17.2

3.83 1.05 0.32 0.21 0.12

10.7 11.7 13.1 10.3 9.7

554


Fig. 1. Percentages of the grain size fractions as a function of depth (cm) in the soil profile. Coarse sand (0.2 – 2 mm), fine sand (0.02 – 0.2 mm), silt (0.002 – 0.02 mm) and clays ( £ 2 mm).

1994), loam, weak fine crumb structure (by organic matter), friable (most). Less than 10 % gravel by volume. Common fine and medium roots and pores. Clear smooth boundary. 20 – 65 cm: B, yellowish brown (10 YR 5/5), loam (to sandy-loam), weak fine and medium subangular blocky structure. Friable (most). Few fine pores. Thin coatings in pad face (dark yellowish brown, 10 YR 4/4). Few streaks, less than 5 % (reddish yellow, 7.5 YR 6/8). Common subvertical medium roots. Gradual boundary. 65 –130 cm: C, 40 % fine and medium gravely materials. Massive wet.


555

Methods and experimental procedures Petrographic observations Five thin sections were prepared from undisturbed whole soil samples and one from substrate rock. Moreover, four thin sections were prepared from the upper level rocks. Chemical analyses Bulk chemical analyses were performed on the upper level rocks, soil and substrate rock. Chemical elements were analyzed by X-ray fluorescence (XRF) and loss on ignition (LOI) at 1100 Ê C was measured by thermogravimetric analysis. FeO was measured by titration. X-Ray Diffractometric (XRD) study XRD analyses were performed on soil and powdered rocks samples with a Siemens D-5000 diffractometer (Bragg-Brentano geometry, CuKa radiation, 40 kV, 30 mA). XRD patterns were examined in the interval 2.5 – 33Ê 2q, to determine the contribution of each mineral phase. We used the “Profile fitting” program of the Diffrac-At software (1991). The decomposition procedure was carried out, according to the criteria of Lanson (1992), Lanson & Besson (1992) and Lanson (1997), by progressively increasing the number of gaussian shaped peaks to obtain the best fit with the smallest number of elementary curves. Three series of diffractograms were recorded: – survey XRD on bulk soil samples ( < 2 mm) in order to establish their overall mineralogical composition; – XRD on clay fractions ( < 2 mm) obtained from the soil samples by sedimentation after disintegration of cements by ultrasonic dispersion; – XRD on oriented clay fractions of B horizon. The oriented aggregates previously were Mg and K saturated; then they underwent glycerol solvation and thermal treatments (2 hr at 330 and 2 hr at 500 Ê C). In addition, the Tamura test (Tamura 1958) was applied to the oriented clay fraction, in order to extract Al and Fe hydroxides which hinder the complete collapse, after thermal treatment, of the interlayer sheet of phyllosilicates.

556


High resolution transmission electron microscopy (HRTEM) study For HRTEM study, slices of undisturbed soil fragments, containing optically selected mica-rich zones, were removed from petrographic sections and ion-thinned with argon beams following the procedures detailed by Amouric & Parron (1985). Observations were made with a JEOL 2000 FX instrument (200 kV accelerating voltage, 2.3 mm spherical aberration, 50 mm objective aperture). Imaging with optimum defocus for micas (from 800 Å to 1200 Å, after Amouric et al. 1981) was selected. To avoid possible electron-beam damage to the specimen, tilt procedures were not used so that only one-dimensional images were recorded. However most of these images showed characteristic 00 l lattice fringes of the dominant mineral phases present in the samples. Simultaneously, energy dispersive X-ray (EDX) analyses were carried on the same samples so that microchemical analyses were systematically and directly coupled with each selected structure image from HRTEM. A Tracor Northern system (TN 5502, series II) was used as AEM and beam probes 50 to 150 Å in diameter were adopted throughout the study for a better correlation between structure and chemical analyses.

Results Optical petrography In all the soil horizons are observed: – Fe – oxyhydroxides, which are responsible of the yellowish-brown colour; – well-preserved white micas; – kaolinite, generally embedded with Fe – oxyhydroxides; – a few biotite, frequently associated to the white mica that constitute, with quartz and Na – K-feldspars, the primary mineral assemblage. In the substrate rock (gneiss) only white micas occur, with quartz and some Na- and K-feldspar relics. The major components of the upper level rocks are biotite, phengite, quartz, albite and K-feldspar; the minor components are chlorite, amphibole and, more rarely, titanite, garnet, apatite and allanite.


557

Table 2. Bulk analyses (%, 105 Ê C dried sample) of the upper level rocks, soil horizons and substrate rock. Precision range: 0.01 %. Sample

SiO2

Al2O3 Fe2O3 FeO TiO 2 MnO

MgO CaO Na2 O K2O LOI1

S

upper level rock 72.46 13.27

2.76

2.02 0.40

0.05

0.22

1.17 3.08

5.62

0.35 101.40


2.59

1.91 0.38

0.05

0.48

0.87 2.94

5.81

0.49 100.18


2.32

1.34 0.31

0.05

0.29

0.90 2.73

4.50

0.78 100.53


1.87

1.00 0.25

0.06

0.27

1.11 2.89

4.80

0.92

A1 horizon

65.87 15.24

3.79

0.70 0.43

0.02

1.09

0.12 1.16

4.02

7.73 100.17

B horizon

55.11 22.32

5.48

0.36 0.68

0.03

1.23

0.06 1.52

2.59 10.84 100.22

B horizon

51.48 24.36

6.28

0.24 0.74

0.03

1.27

0.04 1.54

2.53 11.88 100.39

C horizon

52.38 23.75

5.89

0.26 0.74

0.04

1.12

0.12 1.46

2.62 11.08

C horizon

53.44 23.88

5.73

0.29 0.74

0.04

1.35

0.07 1.62

2.63 10.77 100.56

substrate rock

78.12 11.81

1.83

0.35 0.16

0.005 0.44

0.09 0.69

4.05

1

2.23

99.64

99.46 99.77

LOI: Loss on ignition at 1100 Ê C.

Chemical analyses Chemical analyses on bulk samples are given in Table 2. Chemical differences between the upper level rocks, the A1 horizon, the other soil horizons and the substrate rock are important. X-ray diffraction (XRD). Oriented clay fraction (B horizon) The patterns recorded from different horizons are quite similar, differing only by the relative intensity of diffraction peaks. Because of this similarity, detailed examination was confined to the samples from the B horizon, the horizon richest in clay (Fig. 1). The comparison between XRD patterns of Mg-saturated and Mg-saturated glycerol-solvated clay fractions shows that no change occurs after glycerol solvation, indicating that the samples do not contain expandable phases such as smectites. The decomposition of the 3 –10Ê 2q range of the XRD diagram for Mgsaturated oriented clay aggregates gave three basic curves with maximum intensities at 14.33, 12.51 and 9.98 Å (Fig. 2 a). In the 10 –14Ê 2q range two elementary curves are obtained at 7.53 and 7.25 Å respectively. In the 17– 33Ê 2q range the maximum intensities appear at 4.87, 3.74, 3.57, 3.43 and 3.33 Å (not shown). The peak at 14.33 Å was attributed to vermiculite (and/ or chlorite).

558


Fig. 2. a, b.

According to simulations with the NEWMOD program (Reynolds Jr. & Reynolds 1996) the peak at 12.51 Å, associated with those at 4.87 and 3.43 Å, was attributed to interstratified mica – vermiculite (or chlorite), with 60 % mica. The peaks at 9.98 and 3.33 Å indicate mica layers. The pair of peaks at 7.53 and 3.74 Å, along with the pair at 7.25 and 3.57 Å, were attributed to halloysite and kaolinite, respectively.


559

Fig. 2. Decomposed XRD pattern of oriented clay fraction of B horizon (at 55 cm depth). a) Mg-saturated, air-dried sample, b) Mg-saturated sample heated to 330 Ê C, 2 hr; c) Mg-saturated sample heated to 500 Ê C, 2 hr; d) Tamura test: Na – citrate equilibration, K-saturation and thermal treatment (500 Ê C, 2 hr). CuKa radiation, d-spacing in Å, —— experimental curve, - - - - elementary computed curve, best fit computed curve.

560


Thermal Treatment and Tamura Test The samples Mg-saturated, air dried (Fig. 2 a) were thermally treated at 330 Ê C for 2 hr (Fig. 2 b). The intensity of the curves at 14.33 and 12.51 Å is reduced and their peaks are displaced to 14.27 and 12.11 Å respectively, while the intensity of the 9.98 Å peak increases and is resolved into two curves at 10.00 and 10.26 Å. The 7 Å peak becomes more symmetric: the curve at 7.53 Å shifts towards 7.31 Å and the 7.25 slightly moves to 7.21 Å, with an increased intensity (Fig. 2 b). For samples were maintained for 2 hr at 500 Ê C (Fig. 2 c). The 7 Å peak disappeared, whereas the collapse of the 14 Å layers was largely incomplete, as indicated by the curves at 14.11, 12.13, 10.52 and 10.13 Å. According to Tamura (1958) the citric ion extracts the Al – hydroxypolycations from the vermiculite interlayers; this induces, in turn, a more marked effect on the successive K-saturation and thermal treatment. Hence the sample was equilibrated with Na – citrate solution, K-saturated and, finally, thermally treated at 500 Ê C for 2 hr (Fig. 2 d). After the Tamura test, the experimental profile becomes more symmetric and shifts towards the 10 Å spacing. Therefore, the 14.33 Å spacing (on Fig. 2 a) indicates the presence of both HIV and Mg-vermiculite. Non-oriented samples of the bulk soil The XRD diagrams taken on the coarse sand fraction of the B horizon (Fig. 3 a, b) yield further information on the minerals forming the soil. The decomposition of the experimental XRD profile gave the following results: – 3 –10Ê 2q range, six elementary curves were obtained centred at 24.4, 14.30, 12.23 (broad), 12.19, 10.00 and 9.97 Å; – 10 –14Ê 2q range, two curves were obtained at 7.73 and 7.28 Å; – 17–19Ê 2q range, three curves were found centred at 4.98 (low intensity), 4.97 (high intensity) and 4.89 Å (broad) respectively. As for the clay fraction, no change occurred after glycerol solvation, indicating the absence of expandable phases. Furthermore, after thermal treatment at 500 Ê C, the peak at 14.30 Å collapsed completely to 10 Å. This shift could be attributed to vermiculite. From the NEWMOD simulation (Reynolds Jr. & Reynolds 1996), the peaks at 24.4, 12.19 and 4.89 Å, considering as well those at 3.49 and 3.04 Å (not shown), may be attributed to fully ordered interstratified mica – vermi-


561

Fig. 3. Decomposed XRD pattern of coarse sand fraction of B horizon (at 55 cm depth). a) 3 –14Ê 2q range; b) 17–19Ê 2q range. CuKa radiation, d-spacing in Å, —— experimental curve, - - - - elementary computed curve, best fit computed curve.

562


Fig. 4. Decomposed XRD pattern of silt fraction of A1 horizon (at 15 cm depth). CuK a radiation, d-spacing in Å, —— experimental curve, - - - - elementary computed curve, best fit computed curve.

culite (Reynolds 1980). The broad curve centred at 12.23 Å, together with those at 4.89 and 3.44 Å (not shown), indicates the presence of disordered interlayer mica – vermiculite, with 65 % mica. The very strong sharp peak at 9.97 Å and that at 4.97 Å, as well as those at 3.32 and 1.499 Å (not shown), denote the presence of dioctahedral mica. The peak at 10.00 Å, weaker and broader than that at 9.97 Å, when considered together with the weak peak at 4.98 Å and with the broad one at 3.33 Å (not shown), could be attributed to a different mica. The pair of curves at 7.73 and 7.28 Å, taken together with that at 3.76 and that at 3.65 Å respectively (not shown), are assigned to halloysite and kaolinite. Two peaks (at 6.40 and 5.92 Å, not shown) found in the patterns of all the horizons reveal the presence of albite. The XRD pattern of the samples from the A1 horizon (silt fraction) shows, besides the typical halloysite and kaolinite peaks (7.32 Å and 7.16 Å), a sharp peak at 7.04 Å (Fig. 4). The differences in the basal spacings of halloysite (7.73 and 7.32 Å) and kaolinite (7.28 and 7.16 Å) found in the samples from B and A1 horizons, can be attributed to defect structures and different hydration degrees (Giese 1991). The latter peak at 7.04 Å, along with other two weaker peaks at 14.02 and 4.72 Å (not shown), may be identified as the main basal reflection of a chlorite.


563

Fig. 5. Decomposed XRD pattern of substrate rock (at 140 cm depth). CuK a radiation, d-spacing in Å, —— experimental curve, - - - - elementary computed curve, best fit computed curve.

Samples of substrate and upper level rocks The XRD patterns of samples of substrate rocks allow to infer further information on the nature of primary micas occurring in the soil. Figure 5 shows the only peak appearing within the 3 –14Ê 2q range. The decomposition yields three curves at 9.96, 9.95 and 9.93 Å, respectively. The peak at 9.95 Å with the full width at half maximum intensity (FWHM) of the 0.15Ê 2q, associated with those at 4.97, 3.32 and 1.540 Å (not shown) indicate a trioctahedral mica, whereas the broad curve at 9.96 Å represents a larger coherent scattering domain size (CSDS) of the same crystallographic phase. On the same figure the curve at 9.93 Å, along with those at 4.97, 3.31, and 1.510 Å (not shown) indicates a dioctahedral mica. In the substrate rocks the peak intensity of the dioctahedral mica always prevails over the trioctahedral one. The XRD patterns (Fig. 6) of the sample of upper level rocks (at an elevation of 800 m) established that the dioctahedral mica has one size distribution (9.93 Å spacing, FWHM = 0.08Ê 2q), while the trioctahedral one displays two distributions: 10.06 Å, FWHM = 0.10Ê 2q, and another much larger (10.06 Å, FWHM = 0.30Ê 2q).

564


Fig. 6. Decomposed XRD pattern of upper level rocks (elevation of 800 m). CuK a radiation, d-spacing in Å, —— experimental curve, - - - - elementary computed curve, best fit computed curve.

High resolution transmission electron microscopy (HRTEM) and analytical electron microscopy (AEM) results The general texture Low magnification images ( £100,000 ´) show the respective morphologies as the intimate and characteristic organization of the main phases of the soil horizons. Such images are useful, for example, to demonstrate that the relics of primary rock, and particularly the micas, undoubtedly suffered a deformation stage during an early metamorphic event (Fig. 7) or to reveal a true mosaic texture (of mainly neo formed phases) in the soil (Fig. 8). Two groups of minerals, related to the primary micas, may be distinguished. First the “layered phases” which appear as more or less large laths (from 0.01 to few micrometers, depending on the phase) are shown in Fig. 7, where micas and kaolinite were identified. Secondly, several intermixed “non-layered phases” are frequently observed, as the granular and more-orless structurally organized “proto goethite” close to the Fe – oxyhydroxides (Figs. 7 and 8) or anatase crystallites. In addition an inhomogeneous Fe – Si rich material with a very weak contrast and resembling a gel may also exist in between the mica layers (Fig. 9). The 10 Å phyllosilicates Mainly a white and sometimes a black mica are seen in thin sections in the different samples. According to the depth in the


565

Fig. 7. Relic of primary rock in B horizon of the soil. The low magnification TEM image showing a mica deformation suffered during an early metamorphic event.

soil profile these primary micas are more or less abundant. HRTEM and AEM analyses specified their respective type. Structurally the two micas are characterized by their basal spacing of about 10 Å (Figs. 10 and 11). According to Amouric et al. (1981) and Amouric & Baronnet (1983), polytypic sequences and corresponding microdiffraction patterns allow their differentiation, showing predominant 2 M1 (or 1 M) and 1 M (or 1 Md) stacking sequence for the dioctahedral and the trioctahedral variety, respectively. The chemical analyses indicate that these are more or less K-depleted phengite (analysis 1, Table 3) and biotite (analysis 2, Table 3). The loss of K is usual for such a weathered material (Banfield & Eggleton 1988).

566


Fig. 8. Low magnification TEM image from a weathered zone of the B horizon of the soil. (H: halloysite, K: kaolinite, phengite and Fe – oxyhydroxides).


567

Fig. 9. Low magnification TEM image showing interstratified “non-layered phases” as Fe – Si rich gel-like materials in between the opened and leached phengite layers (B horizon).

Fig. 10. HRTEM image of kaolinite layers grown on leached phengite (phen. l.). The arrow indicates a probable lateral transformation from phengite to kaolinite (B horizon).

In both micas evidently disaggregated or dissolved areas are observed in which amorphous or poorly crystalline Fe – Si material frequently occur (Fig. 9).

568


Fig. 11. HRTEM image of random and, locally, ordered 1 : 3 interstratified between vermiculite (small arrows) and leached biotite (biot. l.) revealing a stage of vermiculitization in B horizon.

A third type of 10 Å phase is also observed, located in the matrix of mainly neo formed minerals, around the primary micas. Such a phase (Fig. 12) generally consists of small laths, with a less organized structure compared to micas. This phase suffered from beam damage, and is compatible with illitic chemical composition (analysis 3, Table 3). The 14 Å phyllosilicates Associated with the original micas, two groups of layered 14 Å phases are detected. For the first group, HRTEM images (Fig. 13) demonstrate that they consist of 9-Å (talc-like) and 5-Å (brucitelike) units alternating along c*; this is characteristic of a chlorite type structure (Olives & Amouric 1984, Olives 1985, Amouric et al. 1988, De Parseval et al. 1994). The corresponding microchemical analyses were not easy to obtain because these [(TOT)O] structures were only infrequently in the specimen. They are generally poorly developed, as relics, and intimately mixed with other phases. Nevertheless AEM data indicate it could be Mg – chlorite interstratified with biotite (analysis 4, Table 3) and Al – , Mg – chlorite when surrounded by a poorly crystalline phase (Fig. 13; analysis 5, Table 3).


569

Table 3. AEM data of the main layered phases from B horizon (iron is assumed to be completely Fe3 + ). Ana- Specimen zone lyses 1 2 3 4 5

6

7 8

Phengite 1 Weathered biotite 1 Illite1 Biotite and Mgchlorite mixture 2 Al-, Mg-chlorite surrounded by a poorly crystalline phase 2 Biotite and Mg, Fe-vermiculite mixture 1 Biotite and HIV mixture 1 Kaolinite 3

Si

IV

3.15 2.76 3.44 3.48

0.85 1.24 0.56 0.52

Al

VI

Al Fe

1.47 0.39 1.65 1.26

0.23 0.98 0.27 1.37

Mg

Ti

S Oct K

Na

Ca

S Int

0.34 1.09 0.20 1.85

0.03 0.08 0.01 0.09

2.07 2.54 2.13 4.57

b.d. 0.01 0.04 0.03

b.d. b.d. b.d. 0.01

0.94 0.64 0.37 0.56

0.94 0.63 0.33 0.52

3.87 0.13 3.31 0.36 0.52 b.d. 4.19

0.03 0.04 0.01 0.07

2.85 1.15 0.36 0.90 1.32 0.08 2.66

0.30 0.07 b.d. 0.38

3.31 0.69 1.27 0.48 0.52 0.05 2.32

0.18 b.d. 0.01 0.18

2.06 b.d. 1.78 0.10 0.06 b.d. 1.94

b.d. b.d. b.d. 0.00

1

data normalized to O10(OH)2. data normalized to O10(OH)8. 3 data normalized to O5(OH)4. b.d. Below detection. 2

Fig. 12. Low magnification TEM image of small laths of illite located in the matrix of neo formed minerals (B horizon). See also the hollow tubes of halloysite.

570


Fig. 13. HRTEM image of alternating units (9 Å + 5 Å) characteristic of chlorite type structure (B horizon).

The second group is much more representative. The 14 Å phases are generally observed in interstratified configuration and only with biotite. Details of HRTEM images are generally different from those for chlorites. For example, in Fig. 11 several 14 Å single layers are visible, mixed with 10 Å biotite layers. Thick white fringes, located at the interlayer levels (bright contrast), characterize this group of 14 Å phases, which easily permit the location of corresponding single layers in the images. Visible expansion and increase in brightness of the interlayer region in biotite sequences, combined with the corresponding microchemical analyses (analysis 6, Table 3) suggest a replacement of K by hydrated interlayer cations, revealing vermiculitization of biotite, as reported in equivalent weathering conditions by Banfield & Eggleton (1988) and Banfield & Barker (1998). The biotite – vermiculite interstratification is observed in details, giving either disordered or more or less ordered configurations (Fig. 11). The regularity and the number of vermiculite layers in such sequences are probably related to the amount of weathering (Pozzuoli et al. 1992). Sometimes, and always in biotite crystallites, when these sequences are observed slightly out of the best conditions, HRTEM images can directly reveal 24 Å


571

(10 + 14) spacing as a periodicity along c*. Such arrangements suggest that possible regular 1 : 1 biotite – vermiculite interlayer (which suffer extensive beam damage) may exist. Structurally detailed images showing extensive zones of pure vermiculite were, apparently, not obtained. Lastly, the microchemical analyses systematically carried out on each type of the observed sequence, globally demonstrate and specify that this second group of 14 Å phases can be classified as biotite and Mg– , Fe – vermiculite mixture (analysis 6, Table 3) and as biotite and HIV mixture (analysis 7, Table 3). The 7 Å phyllosilicates These minerals occur as two distinct phases. One is mainly observed intimately associated with both types of micas and the other is more ubiquitous, either near the micas or in the global matrix of the soil. The first type of phase is easily identified as kaolinite, due to its characteristic 7 Å basal spacing (Fig. 10; analysis 8, Table 3). A very fast alteration of the image contrast and an apparent destruction of the structure under the electron beam are also characteristics of this phyllosilicate. Kaolinite was observed either alone or lying in a parallel contact with biotite but it was never observed interstratified with it. The relationships with phengite are more complex. The most frequently observed structural relation is illustrated in Fig. 10, showing kaolinite clearly developed on phengite with its d(00 l) planes parallel (or quasi parallel) to those of mica. Also in Fig. 10 a rare lateral transformation from phengite to kaolinite is shown. However, in a few ambiguous cases, a mixture of the respective basal planes (about 7 and 10 Å spacing) could be distinguished with difficulty, in phengite, over a thin zone, and always near the interface with kaolinite, as reported by Robert son & Eggleton (1991). The second category of 7 Å phyllosilicates is recognized as halloysite from its typical morphology (Figs. 8 and 12), its very fast alteration under the electron beam, and its chemical composition similar to that of kaolinite. It is seen in two possible configurations: 1) very near and/or in structural continuity with kaolinite and 2) free, in voids around micas or in matrix.

Discussion The chemical analyses show that there is no homogeneity between the upper level rocks, substrate rocks and the soil. Within the soil, the composition of the A1 horizon is definitely differentiated from that of the B and C

572


horizons. Thus, the soil formed not only from the weathering of residuum, but it also received input of colluvial material generated from up-slope sources. As regards the XRD investigation of micas, to identify di- and trioctahedral micas we used the ratio between the peak intensities of the 5 Å and 10 Å spacings which, in oriented patterns, amounts to about 1/3 for the dioctahedral micas, being instead very weak for the trioctahedral micas (Wilson 1987). The 060 reflections were not use to distinguish di- trioctahedral micas because the quartz peak at 1.542 Å overlap the 060 reflections of trioctahedral micas. As is amply pointed out in the literature, trioctahedral micas are more easily altered, in a pedogenetic environment, than dioctahedral micas. Our results confirm this in the more marked broadening of the principal basal reflection of the trioctahedral micas, pointing to lower crystallinity. The AEM analyses confirm that K + is much more easily leached from the trioctahedral micas. The oxidation of iron, Fe2 + ® Fe3 + likely to occur in the soil, and/or the possible presence of H3O + in place of K + in the interlayer spaces, may explain the necessary equilibrium charge and appear to be likely mechanisms for the initial weathering processes in these minerals. Biotite crystallites exhibited only small packets of apparently regular micalayers, and many zones are interstratified with Mg– , Fe – vermiculite and HIV (Fig. 11), as previously found by Banfield & Eggleton (1988), Bain et al. (1990) and Harris et al. (1992). The areas of interstratified, fully ordered, mica – vermiculite, which originates the 24.4 Å and higher order peaks in XRD patterns, actually were not observed in HRTEM, whereas it was often possible to localize regions of 14 Å disordered layers interlayered with altered mica (Fig. 11). This discrepancy may be attributed to the damage that the vermiculite layers underwent, first during the standard ion-thinning of the specimen, then under the electron beam, especially when they occur as independent layers, whilst their collapse or amorphization seems to be hindered when “protected” by a large enough amount of interlayered mica. According to Olives & Amouric (1984) and Amouric et al. (1988) an octahedral layer may replace an interlayer K + and give origin to chlorite, which is more stable than vermiculite under the electronic beam and therefore shows a better resolution in HRTEM (Fig. 13). As concerns the kaolinite – phengite topology, two considerations should be made. First, the observations of the kaolinite developing on phengite


573

layers along with the more rarely observed lateral transformation from phengite to kaolinite (Fig. 10), suggest both epi- and topotactic textures. Secondly, referring to the parameters of dioctahedral 2 M1 mica (Bailey 1987) and of kaolinite (Giese 1991), the best fit on the (001) plane results to be: (aph – ak)/ak = 0.06 % and (bph – bk)/bk = 0.06 %, Dg = 0.177Ê (where ph = phengite, k = kaolinite). This is, of course, only necessary but not sufficient condition to promote epi- and/or topotaxy. Nevertheless, both observations and geometrical conditions cooperate in favour of both the hypotheses. TEM observations demonstrate that, in cross section, halloysite can generally occur with two main and common morphologies: in spirals (when issued from kaolinite) and in hollow tubes (Figs. 8 and 12). Commonly, the presence of halloysite indicates a late stage of weathering in a soil and is generally accompanied by a global hydration of the primary material and phases of this soil (Murray 1991).

Conclusion XRD diagram decomposition and HRTEM observations were used to identify the layered minerals forming soil and gave consistent and complementary information both for the extracted clay fractions and for the bulk samples. In particular, the ion-thinning technique for HRTEM specimens preserves the grain-matrix textural relationships, which are very important for the study of mica transformation in a soil. In short one can conclude that: – a white mica phase corresponds to a dioctahedral mica (phengite) still fairly well preserved showing predominant 2 M1 polytypic sequences; – the biotite relics correspond to the trioctahedral mica showing, on XRD patterns, broad CSDS distribution around 10.0 Å; this mica is a more or less K-depleted biotite observed by HRTEM, with frequent 1 Md stacking sequences and poorly conserved structures; – the third 10 Å phase, with low CSDS distribution and XRD patterns varying in 2q position and peaks intensity with soil horizons and grain size classes, surrounds primary micas and shows illitic composition; – both kaolinite and halloysites were commonly observed by HRTEM; – the few chlorite layers, found in the silt fraction from A1 horizon, are rare probable relics, as Mg – chlorite when intimately mixed with biotite, and as Al– , Mg – chlorite when surrounded from a poorly crystalline phase.

574


Two coexisting vermiculites are determined: Mg – vermiculite and HIV. Pure vermiculite was not clearly observed, whilst different vermiculitization stages of biotite were identified everywhere, either as disordered or more or less ordered interlayer. The same applies to the regular 1 : 1 interlayer biotite – vermiculite, found by XRD and poorly imaged with HRTEM.

Acknowledgements We wish to thank professor M. Calleri who kindly read the manuscript. Thanks are also due to Ministero per l’Università e la Ricerca Scientifica (MURST – Roma, funds ex-40 %) and to Centro Nazionale delle Ricerche (CNR – Roma) for financial support.

References Amouric, M. & Baronnet, A. (1983): Effect of early nucleation conditions on synthetic muscovite polytypism as seen by high resolution transmission electron microscopy. – Physics Chem. Miner. 9: 146 –159. Amouric, M., Giannet to, I. & Proust, D. (1988): 7–10 and 14 Å mixed-layer phyllosilicates studied structurally by TEM in pelitic rocks of the Piemontese zone (Venezuela). – Bull. Minér. 111: 29 – 37. Amouric, M., Mercuriot, G. & Baronnet, A. (1981): On computed and observed HRTEM images of perfect mica polytypes. – Bull. Minér. 104: 298 – 313. Amouric, M. & Parron, C. (1985): Structure and growth mechanism of glauconite as seen by high-resolution transmission electron microscopy. – Clays Clay Miner. 33, 6: 473 – 482. Aoudjit, H. (1993): Genèse, cristallochimie, organisation et propriétés physico-chimiques des argiles des sols et des arènes développées sur roches cristallines (France). – PhD thesis, Univ. Paris 6, France. Aoudjit, H., Robert, M., Elsass, F. & Curmi, P. (1995): Detailed study of smectite genesis in granite saprolites by analytical electron microscopy. – Clay Miner. 30: 135 –148. Bailey, S. W. (1987): Appendix: X-ray powder patterns of micas. – In: Bailey , S. W. (ed.): Micas. Reviews in Mineralogy, Mineral. Soc. of America, Washington D.C. 13: 574. Bain, D. C., Mellor, A. & Wilson, M. S. (1990): Nature and origin of an aluminous vermiculitic weathering product in acid soil from upland catchments in Scotland. – Clay Miner. 25: 467– 475. Banfield, J. F. & Egglet on, R. A. (1988): Transmission electron microscope study of biotite weathering. – Clays Clay Miner. 36, 1: 47– 60.


575

– – (1990): Analytical transmission electron microscope studies of plagioclase, muscovite, and K-feldspar weathering. – Clays Clay Miner. 38, 1: 77– 89. Banfield, J. F. & Barker, W. W. (1998): Low-temperature alteration in tuffs from Yucca Mountain, Nevada. – Clays Clay Miner. 46, 1: 27– 37. De Parseval, P., Amouric, M., Baronnet, A., Fortune, J. P., Moine, B. & Ferret, J. (1994): HRTEM study of the chloritization of micas in the talc – chlorite deposit at Trimouns (Pyrenees, France). – Eur. J. Mineral. 6: 123 –132. DIFFRAC – AT (1991): Software package. – Siemens. 3.0. Elsass, F. & Robert, M. (1992): Application of high resolution electron microscopy to soil clay origin and organisation in a temperate climate. – Geologica Carpathica Service Clays 2: 55 – 61. Giese, R. F. (1991): Kaolin minerals, structures and stabilities. – In: Bailey, S. W. (ed.): Hydrous Phyllosilicates (Exclusive of Micas). Rev. in Miner., Miner. Soc. of America, Washington D.C. 19: 29 – 66. Harris, W. G., Morrone, A. A. & Coleman, S. E. (1992): Occluded mica in hydrozy-interlayered vermiculite grains from a highly-weathered soil. – Clays Clay Miner. 40, 1: 32 – 39. Lanson, B. (1992): Application de la décomposition des diffractogrammes de rayons X à l’identification des minéraux argileux. – Siemens (ed.): Comptesrendus du Colloque rayons X Siemens, Paris 2: 62 –72. – (1997): Decomposition of experimental X-ray diffraction patterns (profile fitting): a convenient way to study clay minerals. – Clays Clay Miner. 45, 2: 132 – 146. Lanson, B. & Besson, G. (1992): Characterization of the end of the smectite-toillite transformation: Decomposition of X-ray patterns. – Clays Clay Miner. 40: 40 – 52. Lanson, B. & Velde, B. (1992): Decomposition of X-ray diffraction patterns: a convenient way to describe complex diagenetics smectite-to-illite evolution. – Clays Clay Miner. 40: 629 – 643. Munsell Soil Color Charts (1994): Revised edition. – Munsell Color, Gretag Macbeth, New York. Murray, H. H. (1991): Kaolin minerals: their genesis and occurrences. – In: Bailey, S. W. (ed.): Hydrous Phyllosilicates (Exclusive of Micas). Reviews in Mineralogy, Miner. Soc. of America, Washington D.C. 19: 67– 89. Olives, J. (1985): Biotites and chlorites as interlayered biotite – chlorite crystals. – Bull. Minér. 108: 635 – 641. Olives, J. & Amouric, M. (1984): Biotite chloritization by interlayer brucitization as seen by HRTEM. – Amer. Miner. 69: 869 – 871. Pozzuoli, A., Vila, E., Franco, E., Ruiz-Amil, A. & Calle, C. de la (1992): Weathering of biotite to vermiculite in quaternary lahars from Monti Ernici, Central Italy. – Clay Miner. 27: 175 –184.

576


Reynolds, R. C. (1980): Interstratified Clay Minerals. – In: Brindley, G. W. & Brown, G. (eds.): Crystals structures of clay minerals and their X-ray identification. Miner. Soc., London: 249 – 303. Reynolds, R. C. Jr. & Reynolds, R. C. III (1996): Description of program NEWMOD for Windows. The calculation of one-dimensional X-ray diffraction pattern of mixed-layered clay minerals. – Hanover, New Hampshire, 8 Brook Road, 25 pp. Righi, D. & Elsass, F. (1996): Characterization of soil clay minerals: decomposition of X-ray diffraction diagrams and high-resolution electron microscopy. – Clays Clay Miner. 44, 6: 791– 800. Righi, D. & Meunier, A. (1991): Characterization and genetic interpretation of clays in an acid brown soil (dystrochrept) developed in a granitic saprolite. – Clays Clay Miner. 39, 5: 519 – 530. Righi, D., Petit, S. & Bouchet, A. (1993): Characterization of hydroxy-interlayered vermiculite and illite/smectite interstratified minerals from the weathering of chlorite in a cryorthod. – Clays Clay Miner. 41, 4: 484 – 495. Robertson Ian, D. M. & Egglet on, R. A. (1991): Weathering of granitic muscovite to kaolinite and halloysite and of plagioclase-derived kaolinite to halloysite. – Clays Clay Miner. 39, 2: 113 –126. Romero, R., Robert, M., Elsass, F. & Garcia, C. (1992): Evidence by transmission electron microscopy of weathering microsystems in soils developed from crystalline rocks. – Clay Miner. 27: 21– 23. Tamura, T. (1958): Identification of clay minerals from acid soils. – J. Soil Sci. 9: 141–147. USDA – U.S. Department of Agriculture (1986): Clés de la taxonomie des sols. – Monographie technique, Cornell University, Ithaca, New York 13: 347. Wilson, M. J. (1987): X-ray powder diffraction methods. – In: Wilson, M. J. (ed.): A handbook of determinative methods in clay mineralogy. – Chapman and Hall, New York: 48 – 52. Received: January 29, 2002. Authors’ addresses: Bruno Alessandria, Dino Aquilano, Dipartimento di Scienze Mineralogiche e Petrologiche dell’ Università di Torino, via Valperga Caluso 35, 10125 Torino, Italy. Marc Amouric, Oliv ier Grauby, Centre de Recherche sur les Mécanismes de la Croissance Cristalline, CNRS, Campus de Luminy, Case 913, 13288 Marseille Cedex 9, France. Corresponding author: Bruno Alessandrina, Dipartimento di Scienze Mineralogiche e Petrologiche dell’ Università di Torino, via Valperga Caluso 35, 10125 Torino, Italy. E-mail: [email protected]