Quantitative Mineralogy of a Saprolite-Soil Sequence on a Talc and Chlorite Bearing Substrate, Upper Katanga, D.R

FACULTEIT WETENSCHAPPEN Opleiding Master of Science in de geologie

Quantitative mineralogy of a saprolite-soil sequence on a talc and chlorite bearing substrate, Upper Katanga, D.R. Congo.

Els Timmermans

Academiejaar 2014–2015

Scriptie voorgelegd tot het behalen van de graad Van Master of Science in de geologie

Promotor: Prof. Dr. J. De Grave, Prof. Dr. E. Van Ranst Begeleider: Drs. M. Dumon Leescommissie: Prof. Dr. G. Baert, Prof. Dr. S. Bertrand

ACKNOWLEDGEMENTS

For any geology student, the graduating year is an exciting chapter in life. The master thesis is an important part of it. Luckily I could count on great support of my thesis promotor Prof. Dr. Eric Van Ranst. I would like to thank him for giving me the opportunity to do my master thesis in particular and studying in general at the Laboratory of Soil Science of the Geology and Soil Science Department. Without his knowledge, perseverance and dedicated time I would have a much tougher time in completing this document.

I would like to express my sincere gratitude to my tutor Drs. Mathijs Dumon for answering my questions and for the help and guidance through the practical part of this master thesis.

Veerle Vandenhende guided me in the Laboratory of Soil Science of the Geology and Soil Science Department. She explained and gave me insight into the lab analyses. It was a great pleasure to work with her.

Furthermore, I would like to thank all people of the Soil Science Research Unit of the Ghent University for the company and the help.

Last but not least, my parents and family may not be forgotten. They provided the support I needed to devote so much of my time in my education in general and this document in particular. They cared my every need and bore the considerable financial strain students impose upon parents all for their happiness and fruitful ending of their education.

Els Timmermans

TABLE OF CONTENTS

ACKNOWLEDGEMENTS ...... i LIST OF FIGURES ...... iv LIST OF TABLES ...... vi SAMENVATTING ...... viii 1. INTRODUCTION ...... 1 1.1 Background and research significance ...... 1 1.2 Objectives ...... 2 2. LITERATURE REVIEW ...... 3 2.1 Geology ...... 3 2.2 Weathering and formation of clay minerals ...... 8 2.2.1 Weathering of talc ...... 8 2.2.2 Weathering of chlorite ...... 10 2.2.3 Formation of kaolinite ...... 12 3. MATERIAL AND METHODS ...... 15 3.1 Environmental setting ...... 15 3.2 Sampling strategy ...... 16 3.3 Physico-chemical analyses...... 16 3.3.1 Texture, clay separation and saturation ...... 16 3.3.2 Soil acidity ...... 17 3.3.3 Cation exchange capacity and exchangeable base cations ...... 17 3.3.4 Selective chemical extraction ...... 17 3.3.5 Total elemental analysis ...... 18 3.4 Quantitative X-ray diffraction (QXRD) ...... 18 3.4.1 Random powder samples ...... 21 3.4.2 Oriented samples ...... 22 3.4.3 SiroQuant ...... 23 3.4.4 Python X-ray Diffraction (PyXRD) ...... 24 3.4.5 Calculation of the mineral composition ...... 26 3.5 High resolution transmission electron microscope (HR-TEM) ...... 26 3.6 Micro X-ray fluorescence (XRF) ...... 26 3.7 Micromorphological analysis ...... 27

4. RESULTS ...... 28 4.1 Morphological properties ...... 28 4.2 Physico-chemical properties ...... 28 4.3 Total elemental composition ...... 32 4.4 Weathering indices and molar ratios ...... 34 4.5 Total elemental composition of the silicate fraction ...... 36 4.6 Trace elements ...... 38 4.7 Quantitative X-ray diffraction (QXRD) ...... 39 4.8 High resolution transmission electron microscope (HR-TEM) ...... 47 4.9 Micro X-ray fluorescence (µXRF) ...... 48 4.10 Micromorphological properties ...... 53 5. DISCUSSION ...... 59 5.1 Evolution of the physico-chemical and mineralogical properties of the two selected intervals ...... 59 5.1.1 Interval 1 ...... 59 5.1.2 Interval 2 ...... 61 5.1.3 Weathering degree ...... 63 5.2 Weathering of talc ...... 64 5.3 Weathering of chlorite ...... 65 5.4 Formation of kaolinite ...... 66 5.5 Relief inversion hypothesis ...... 67 6. CONCLUSION ...... 71 6.1 FUTURE RESEARCH ...... 72 7. REFERENCES...... 73 8. APPENDIX 1 ...... 87 9. APPENDIX 2 ...... 89 10. APPENDIX 3 ...... 91 11. APPENDIX 4 ...... 93 12. APPENDIX 5 ...... 94 13. APPENDIX 6 ...... 95 14. APPENDIX 7 ...... 103 15. APPENDIX 8 ...... 111

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LIST OF FIGURES

Figure 1: Location of the sampling area (red dot) (left) (modified from Cailteux et al., 2007) and position of the Pan-African belts and localization of Lufilian arc on the Congo Craton (right) (modified from Jackson et al., 2003)...... 6 Figure 2: Representation of Gondwana assemblage with involved cratons (modified from Gray et al., 2007)...... 6 Figure 3: Representation of the three structural zones of the Lufilian arc (modified from Jackson et al., 2003)...... 7 Figure 4: Schematic projection of talc along the c-axis (modified from Van Ranst, 2013)...... 9 Figure 5: Schematic projection of trioctahedral chlorite along the c-axis (modified from Van Ranst, 2013)...... 11 Figure 6: Schematic projection of kaolinite along the c-axis (modified from Van Ranst, 2013). ... 12 Figure 7: The two reaction mechanisms of the solid-state reaction, smectite kaolinization. The interlayer Al-atoms indicate Al-hydroxy-interlayer complexes (modified from Ryan and Huertas, 2009)...... 13 Figure 8: Location of the D.R. Congo within Africa and of Lubumbashi (red) within D.R. Congo (modified from “marysrosaries” and “webokapi”) ...... 15 Figure 9: Illustration of the RTS method (modified from Zhang et al., 2003)...... 22 Figure 10: Preparation of a smear slide...... 23 Figure 11: USDA soil textural triangle, bleu represents the samples and numbers from interval 1 and red the samples and numbers from interval 2...... 30 Figure 12: Observed (black line) and modelled (red line) XRD patterns of the bulk sample 13/154 from interval 1...... 43 Figure 13: Observed (black line) and modelled (red line) XRD patterns of the bulk sample 13/160 from interval 1...... 44 Figure 14: Observed (black line) and modelled (red line) XRD patterns of the bulk sample 13/360 from interval 2...... 45 Figure 15: Observed (black line) and modelled (red line) XRD patterns of the bulk sample 13/369 from interval 2...... 46 Figure 16: HR-TEM images of (a) talc from 9.30 m to 9.40 m depth (sample 155), (b) kaolin from 9.30 m to 9.40 m depth (sample 155) and (c, d) talc from 34.90 m to 35.09 m depth (sample 363) (modified from Dumon et al., 2015)...... 47 Figure 17: X-ray fluorescence intensity maps of the composing elements of the studied area of sample 356. A photomicrograph of the study area is shown in the uppermost left corner...... 49 Figure 18: X-ray fluorescence intensity maps of the composing elements of the studied area of sample 358. A photomicrograph of the study area is shown in the uppermost left corner...... 50 Figure 19: X-ray fluorescence intensity maps of the composing elements of the studied area of sample 366. A photomicrograph of the study area is shown in the uppermost left corner. The D represents dolomite...... 51 Figure 20: Cumulative X-ray fluorescence spectrum from the studied area of sample 356...... 52

Figure 21: Cumulative X-ray fluorescence spectrum from the studied area of sample 358...... 52 Figure 22: Cumulative X-ray fluorescence spectrum from the studied area of sample 366...... 53 Figure 23: Microscopic characteristics of talc: (a) alignment of weathered talc (9.93-10.10 m); (b, c) altered talc (9.50-9.93 m); (d) talc in dolomite dominated material (36.13-36.23 m); (e, f) talc associated with dolomite (35.85-36.23 m). The D and Tc represent dolomite and talc, respectively...... 56 Figure 24: Micrographs of chlorite at depths between 34 and 36.5 m (interval 2). (a, b) radial shaped chlorite (PPL and XPL) and (c, d, e, f) weathered chlorite (PPL)...... 57 Figure 25: Micrographs of (a) (hydr)oxide nodule with quartz and 2:1 phyllosilicate inclusions (9.00-9.10 m); (b) reductomorphic feature (9.30-9.40 m); (c) cavity filled with (hydr)oxides (9.50-9.64 m); (d) illuvial clay (9.93-10.10 m); (e, f) hematite crystals (H) (34.80-34.90 m and 35.70-35.85 m)...... 58 Figure 26: Visualisation of interval 1 based on the physico-chemical properties, DCB and quantitative analyse. T: texture, QPA: quantitative phase analysis, DCB: dithionite- citrate-bicarbonate and CEC: cation exchange capacity (modified from Dumon et al., 2015)...... 60 Figure 27: Visualisation of interval 2 based on the physico-chemical properties, DCB and quantitative analyse. T: texture, QPA: quantitative phase analysis, DCB: dithionite- citrate-bicarbonate and CEC: cation exchange capacity (modified from Dumon et al., 2015)...... 62 Figure 28: Illustration of the relief inversion hypothesis: (a) peneplain, (b) river incision, (c) gravitational collapse and (d) fill up of the valley. Red: indications of the zones...... 68

LIST OF TABLES

Table 1: Lithostratigraphy and stratigraphic divisions of the Katangan Supergroup (modified from Cailteux et al., 2007) ...... 4 Table 2: Physico-chemical properties of selected samples of interval 1...... 29 Table 3: Physico-chemical properties of the selected samples of interval 2...... 31 Table 4: Total elemental composition (in wt%) of selected samples at different depths...... 33 Table 5: Weathering indices and molar ratios of selected samples at different depths...... 35 Table 6: The elemental composition of the silicate fraction (in wt%) of selected samples at different depths...... 37 Table 7: Concentration of trace elements (in ppm) in the selected samples at different depths. 38 Table 8: Mineral composition (in wt%) of selected bulk samples of the first interval...... 40 Table 9: Mineral composition (in wt%) of selected bulk samples of the second interval...... 41 Table 10: Total elemental composition of the chlorite and C/S phases (in wt%) of sample 13/365...... 66 Table 11: Mineralogical composition of the whole core (modified from Oostermeyer, 2014). .... 70

ABBREVIATIONS AND ACRONYMS

AD air-dry BS base saturation CEC cation exchange capacity CIA chemical index of alteration CSDS coherent scattering domain size DCB dithionite-citrate-bicarbonate D.R. Congo Democratic Republic of the Congo EG (ethylene-)glycolated Fm formation FWO Fonds Wetenschappelijk Onderzoek HR-TEM high resolution transmission electron microscopy ICP-OES inductively coupled plasma optical emission spectrometry LOI loss on ignition PPL plane polarized light R.A.T. Roches Argilo-Talqueuses RTS method Razor Tamped Surface method USDA United States Department of Agriculture.

SF ratio of SiO2 and Fe2O3 wt% weight percentage XANES X-ray absorption near-edge structure XRD X-ray diffraction XPL crossed polarized light µXRF micro X-Ray fluorescence

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SAMENVATTING

Dit onderzoek werd uitgevoerd binnen het kader van het FWO project G.0287.14 getiteld “Formation of kaolinite subgroup minerals in tropical soil-saprolite sequences”. Een 40 m diepe boring werd uitgevoerd in oktober 2011 in de streek rond Lubumbashi, als deel van het FWO project G.0011.10N getiteld “Impact of termites on the mineralogical, textural, molecular and organic composition of tropical soils”. De boorkern werd in stukken verdeeld en in oktober 2012 naar België overgebracht in houten kisten met aanduiding van de diepte. Hierna werd de kern herverdeeld en in kartonnen doosjes van 10 cm gestopt, met aanduiding van het diepte interval. Op basis van vorig onderzoek (Oostermeyer, 2014) werden twee intervallen geselecteerd en bestudeerd. De intervallen werden bestudeerd met het oog op de vorming van kaoliniet en het verweringsgedrag van chloriet en talk.

Het studiegebied waar de boorkern werd genomen, ligt op 12 km ten noordnoordoosten van Lubumbashi in Bumaki, Katanga provincie in D.R. Congo. Het studiegebied behoort tot het miombe ecosysteem (Frost et al., 1986) en wordt gekenmerkt door de frequente aanwezigheid van termietheuvels. Deze kunnen wel 8 m hoog en 15 m breed worden (Sys, 1957; Malaisse, 1974). Lubumbashi is gekarakteriseerd door een warm regenachtig gematigd klimaat met een droge winter en zomer, het regen- en droogzeisoen duren elk 5 maanden (Cws6 (Köppen, 1936)). De gemiddelde jaarneerslag en jaartemperatuur bedragen respectievelijk 1270 mm en 20 °C (Malaisse, 1974; Mbenza, 1990). De bodem werd geclassificeerd volgens de “World Reference Base for Soil Resources” (WRB) als een Ferralsol (Ngongo et al., 2009).

De Katanga Supergroep sedimenten bepalen de geologie van het studiegebied. Het Katangan bekken werd gevormd tussen 900 en 500 Ma ten gevolge van verschillende rift fases gedurende het Neoproterozoïcum. De supergroep is onderverdeeld in 3 groepen, namelijk de Roan, Nguba en Kundulungu Groep. De Roan groep werd afgezet in een continentaal riftbekken (Buffard, 1988), dat ontstond circa 880 tot 820 Ma (Johnson et al., 2007). Deze groep bestaat uit diagenetisch chloriet-dolomiet, quartz, chert, dolosteen, carbonaat- en dolomietrijke schalie en siliciklastisch gesteente (Cailteux et al., 2005b, 2007; Van Langendonck et al., 2013). De plaats waar de boorkern werd genomen, komt een breccia aan het oppervlak, bestaande uit R.A.T., Mines en Dipeta subgroepen. Deze subgroepen behoren tot de Roan groep. Circa 765 ± 5 Ma werd het zuidelijk gedeelte van het riftbekken opgeheven. Dit beëindigde de depositie van de Roan sedimenten en resulteerde in de vorming van de Nguba rift (Wendorff, 2005). De overgangszone bevat dropstenen in zwarte schalie (Wendorff and Key, 2009). De Nguba Groep bestaat onderaan uit een diamictiet (Wendorff and Key, 2009), gevolgd door kalksteen, siliciklastische gesteenten en carbonaatgesteenten (Kennedy et al., 1998). Deze riftfases, waarin zowel de Roan als de Nguba groep werden afgezet, werd gevold door een compressieve fase met bijhorende subductie. Dit leidde tot de vorming van een voorlandbekken dat vervolgens werd opgevuld met de Kundelungu Groep (Jackson et al., 2003). De groep wordt onderaan eveneens gekenmerkt door een diamictiet (Hoffman, 2005). Voor de rest bestaat de groep uit roze kalksteen en niet gemetamorfoseerde gesteenten (Batumike et al., 2007).

Verschillende stalen van beide intervallen werden onderworpen aan verscheidene fysico- chemische en mineralogische analyses, tevens werden X-stralendiffractie (XRD) patronen, micro X-stralen fluorescentie (µXRF) en transmissie electronenmicroscopie (TEM) beelden genomen. Deze werden uitgevoerd volgens de ‘Students Lab Manual’ (Students Lab Manual, 2013) in het Laboratorium voor Bodemkunde van de Universiteit Gent. viii

Interval 1 start op een diepte van 8.90 m en eindigt op 10.60 m, bevat delen van zones C en D. Het interval bestaat uit de volgende mineralen: quartz, hematiet, goethiet, andere oxides (rutiel en pyrolusiet), kaoliniet, mica en talk. De aanwezigheid van de Fe (hydr)oxides en andere oxides wordt bevestigd door de totaal analyse. De berekende silicaatfractie toont aan dat bijna al het ijzer aanwezig is in de Fe (hydr)oxides. Tevens vertonen de goethiet en hematiet gewichtsfracties een inverse relatie met elkaar. Kaoliniet is een typisch mineraal voor zwaar verweerde bodems en wordt beschouwd als een vergevorderd verweringsproduct. Mica en talk daarintegen zijn kenmerkend voor mindere verweerde condities. Op basis van analyses werd de aanwezigheid van dioctahedrische mica’s vastgesteld. Deze zijn kenmerkend voor meer verweerde bodems, en verweren minder snel dan trioctahedrische mica’s. Uiteindelijk zulllen de dioctahedrische mica’s traag verweren naar andere mineralen via gemengd gelaagde (mixed- layer) mineralen (Jackson et al., 1952). Talk is een vrij stabiel mineraal dat zelden voorkomt in bodems en werd voor het eerst herkend in zone C (Oostermeyer, 2014). Het voorkomen van talk en licht verweerde mica (2:1 fyllosilicaten) naast kaoliniet en Fe (hydr)oxides is vrij opmerkelijk. In een normale bodemevolutie zullen deze 2:1 fyllosilicaten al grotendeels afgebroken zijn. Slijpplaatjes van dit interval tonen ook de aanwezigheid van heterogeen moedermateriaal, namelijk een mengsel van grove en grote fragmenten (2:1 fyllosilicaten) ingebedt in een grondmassa. Deze grondmassa bestaat uit kaoliniet, 2:1 fyllosilicaten en Fe (hydr)oxides. De kleinere 2:1 fyllosilicaten wijzen op verwering en de grotere fragmenten vertonen nog geen desintegratie. Dit wijst op het samen voorkomen van verweerd en minder verweerd materiaal. Deze contradictie kan verklaard worden door een reliefinversie. De kationuitwisselingscapaciteit (CEC) waarden variëren tussen 4.12 cmol(+)/kg en 19.17 cmol(+)/kg. Er zijn 3 CEC piekwaarden die kunnen te wijten zijn aan primaire mineralen die iets meer verweerd zijn.

Interval 2 start op een diepte van 33.98 m en eindigt op 36.23 m, en bevat delen van zone G en H. Het interval bestaat uit de volgende mineralen: quartz, dolomiet, veldspaten, mica, smectiet, talk en chloriet en chloriet/smectiet (C/S). Zones G en H kunnen onderscheiden worden op basis van het dolomietgehalte. De veel grotere hoeveelheid dolomiet in zone H wordt bevestigd door een hoog magnesium en calcium gehalte. Het verschil in dolomietgehalte tussen de zones G en H kan te wijten zijn aan verschillend lithologisch moedermateriaal. Tevens bevat zone G een aanzienlijke hoeveelheid veldspaten, wat de voorafgaande hypothese bevestigt. Het moedermateriaal van zone G zal dus arkoses bevatten, dat van zone H chloriethoudende dolosteen. Zone G bevat een belangrijke hoeveelheid smectiet. Dit mineraal is enkel in sporen aanwezig in zone H. Wat erop wijst dat zone G meer verweerd is dan zone H. Talk en mica zijn voornamelijk aanwezig in zone G. Dit bewijst nogmaals het voorkomen van verschillend moedermateriaal binnenin het interval. Zone H is ook zandiger dan zone G. Chloriet en C/S zijn doorheen het volledige interval herkend. Het gaat hier om trioctahedrisch chloriet, dit wordt beaamd door een hoog magnesium en laag aluminium gehalte in de silicaatfractie. In dit interval zijn geen Fe of Mn oxides herkend op basis van de XRD patronen maar werden wel met een optische microscoop en via µXRF herkend. De CEC waarden variëren tussen 3.29 cmol(+)/kg en 58.09 cmol(+)/kg. De pieken in CEC worden verklaard door een grotere hoeveelheid van chloriet en C/S en/of smectiet.

Quartz komt voor in beide intervallen. Het mineraal is aanwezig in de silt- en zandfractie van interval 1 en 2, maar ook in de kleifractie van interval 1. Deze evolutie is te wijten aan het lage verweringspotentiaal van quartz (Goldich, 1938). Verscheidene analyses (bv. morfologie, quantitatieve mineralogie, textuur) tonen ook aan dat interval 1 meer verweerd is dan interval 2.

Chloriet zal verweren via een vaste fase (solid-state) overgangsmechanisme naar vermiculliet en/of smectiet (Ross et al., 1982; Banfield and Murakami, 1998; Meunier, 2005). Het verweringsproces start meestal met de oxidatie van Fe2+ (Banfield and Murakami, 1998) maar kan eveneens gedreven worden enkel door de vrijlating van Fe2+ (Proust et al., 1986). Dit gaat gepaard met de uitspoeling van Mg2+. Uiteindelijk, wordt het mixed-layer mineraal chloriet/vermiculiet of chloriet/smectiet gevormd. Zowel chloriet en C/S werden vastgesteld in interval 2 a.d.h.v. XRD patronen, daarnaast werd verweerde chloriet ook herkend in de slijpplaatjes. Vervolgens worden de mineralen chloriet/vermiculiet en chloriet/smectiet omgezet naar vermiculiet en/of smectiet (Ross et al., 1982; Meunier, 2005). Smectiet is duidelijk herkend in zone G en is enkel aanwezig in zone H als sporen. Uiteindelijk wordt de verweringssequentie van chloriet voltooid door de vorming van kaoliniet uit smectiet.

Twee mechanismen zijn mogelijk voor de vorming van kaoliniet, namelijk nieuw- en solid-state vorming. Bij nieuwvorming slaat kaoliniet neer uit de bodemopslossing, de resterende elementen worden uitgeloogd of slagen neer als (hydr)oxides (bv. Goethiet en hematiet) (Środoń, 1980; De Coninck et al., 1986). Bij de solid-state reactie wordt kaoliniet gevormd vanuit smectiet via smectiet/kaoliniet (Fisher and Ryan, 2006; Ryan and Huertas, 2009). Dit wordt gerealliseerd door 3 processen: [1] het strippen van een tetraëdrische laag uit de smectiet laag, [2] de inversie van een tetraëdrische laag in een smectiet laag en [3] de omzetting van de tussenlaag (interlayer) in een kaolin laag. Geleidelijk aan zullen de smectietlagen vervangen worden door kaolienlagen (Ryan and Huertas, 2009). Het mineraal is enkel herkend in interval 1 op basis van TEM beelden en XRD patronen. De aanwezigheid van Fe binnen kaoliniet zou wijzen op de vorming van het mineraal uit ijzerrijke smectiet dus m.a.w. solid-state transformatie. Uit berekeningen blijkt dat de kaoliniet nagenoeg geen ijzer bevat, d.w.z. dat kaoliniet eerder is gevormd door nieuwvorming. Daarnaast is er eveneens geen enkele indicatie van K/S in de XRD patronen.

Talk komt in beide intervallen voor. Interval 1 vertoont een dalende trend met de diepte. Zone G van interval 2 bevat 20 tot 35 % talk. In zone H daarintegen is talk afwezig of komt enkel voor in kleine hoeveelheden. Weinig literatuur is beschikbaar over de verwering van talk, mede doordat het mineraal zeer zelden voorkomt in bodems. Dit is te wijten aan hun instabiliteit voor fysische verwering (Pérez-Rodrigues et al., 1996) en aan het feit dat hun vorming door pedogenese onwaarschijnlijk is (Zelazny et al., 2002). Talk verweert meestal congruent maar er zijn enkele gevallen gekend waarbij het mineraal incongruent verweert via mixed-layer mineralen naar smectieten (solid-state mechanisme) (Veniale and Van der Marel, 1968; Guenot, 1970; Besnus et al., 1976). De pieken in de XRD patronen die talk aanduiden zijn zeer sherp en verschuiven niet bij glycolatie. Dit duidt op puur talk i.p.v. op talk/smectiet (T/S). Bovendien wordt dit ondersteund door TEM beelden. Dus talk verweert hier waarschijnlijk niet via solid-state transformation. Verder kan talk bewaard blijven in bodems als het wordt afgeschermd van de bodemoplossing door ijzer coatings (Harris et al., 1984; Pérez-Rodríguez et al., 1996). Dit is niet het geval in dit onderzoek, want in de slijpplaatjes zijn talk en Fe (hydr)oxides duidelijk van elkaar afgescheiden. De enige verklaring voor de verwering van talk is progressieve oplossing, waarbij geleidelijk kationen, eerst Mg2+ (Yang et al., 2006), uit het mineraal verwijderd worden, waardoor de structuur uiteindelijk zal uiteenvallen. Niettemin kan talk in theorie vrij resistent zijn tegen verwering, te wijten aan minimale tetrahedrische substitutie waardoor geen water wordt aangetrokken (Temuujin et al., 2003), het gebrek aan interlayer atomen en een ongespannen structuur (Yang et al., 2006). Het voorkomen van talk naast andere primaire

mineralen in interval 2 kan verklaard worden door de lagere verweringsgraad van het interval. Zoals reeds aangehaald is het voorkomen van talk en mica naast kaoliniet en Fe (hydr)oxides in interval 1 tegenstrijdig. Dit kan verklaard worden door een reliefinversie.

Stel dat het originele landoppervlak een peneplain was, bestaande uit moedermateriaal onderaan, gevolgd door saproliet en volledig verweerd materiaal vanboven. Vervolgens werd de peneplain ingesneden door een rivier, waardoor een vallei werd gevormd en saproliet en verweerd materiaal werd verwijderd. Het ontblotende materiaal was vooral minder verweerde saproliet, hiertoe behoren zone G en H. Het onderscheid tussen zone G en H is zoals eerder vermeld te wijten aan verschillend moedermateriaal. Zone G is ook meer verweerd dan zone H. Uiteindelijk werd de valleiwand instabiel en stortte in onder de zwaartekracht. Hierdoor werd er minder en meer verweerd materiaal gemengd en afgezet aan de voethelling. Dit mengsel werd dus afgezet bovenop zone G en vertegenwoordigd zone C en D. Dit wordt bevestigd door de aanwezigheid van een grindlaag in zone C (Oostermeyer, 2014). Deze grindlaag is verarmd aan fijne partikels en aangereikt aan grover materiaal. Dit is een aanwijzing voor een nieuw erosievlak, bovenop de nieuwe afzettingen. De rivier trok steeds verder en verder weg, waardoor er een droge vallei achterbleef. Deze vallei werd later opgevuld met sterk verweerd materiaal, komende van voornamelijk het bovenste deel van het voormalig peneplenatievlak. Dit materiaal vormt zones A en B. Een tweede grindlaag is vastgesteld onderaan zone A. Deze laag werd gevormd door termietactiviteit. Termieten zullen fijner materiaal naar boven brengen en gebruiken voor de opbouw voor hun termietheuvel. Dit leidt tot een accumulatie van grover materiaal onder de termietheuvel (Oostermeyer, 2014). Tenslotte verweerde de top van de afgezette zone A verder waardoor lateriet werd gevormd. Dit verhard materiaal is meer resistent tegen verwering dan het omliggende materiaal (saproliet). Dus het laatstgenoemde verweerde sneller en werd geleidelijk aan verwijderd. Uiteindelijk werd de voormalige vallei volledig opgevuld en evolueerde naar een kam.

1. INTRODUCTION

1.1 Background and research significance

This master dissertation is carried out within the framework of the FWO project G.0287.14 entitled “Formation of kaolinite subgroup minerals in tropical soil-saprolite sequences”. For this study two intervals in a 40 m soil-saprolite sequence under a Macrotermes mound near Lubumbashi in the Katanga province of the D.R. Congo, were analysed. The results from a previous FWO research project “Impact of termites on the mineralogical, textural, molecular and organic composition of tropical soils” (G.0011.10N) indicated mineralogically diverse parent materials are present, including talc- and chlorite-rich formations. As such this material is ideally suited for studying the influence of parent material composition on kaolinite formation pathways in a tropical setting.

Oostermeyer (2014) did an initial investigation within the framework of the FWO project G.0011.10N of the 40 m long core also used in this work. This resulted in a characterization and preliminary delineation of zones in the core. In addition, Mujinya et al. (2010, 2011, 2013, and 2014) have comprehensively examined the Macrotermes mounds in the Lubumbashi region of D.R. Congo. Soil materials of the mound up to 3 m below surface were investigated. One striking finding is these mound materials are enriched in 2:1 phyllosilicates such as primary micas, smectites and illite, while surrounding soils are composed of a typical quartz, kaolinite and iron-oxide mineralogy.

The occurrence of talc in soils is not often reported, due to their susceptibility to weathering (Pérez- Rodriguez et al., 1996). However, several authors have identified the mineral in a soil. The main excepted theory for their preservation is an occlusion of talc by iron oxides.

The humid tropics are characterized by high temperature, abundant rainfall and old geomorphic surfaces (due to limited glacial activity over the last 15000 years). Therefore kaolinite is formed as an (advanced) weathering product, due to (rapid) leaching of base cations and Si. Two formation pathways for kaolinite are suggested namely neoformation (dissolution-crystallisation) and solid- state transformation.

This research focuses on the weathering behaviour of talc and chlorite and the (related) formation of kaolinite throughout the 40 m long core. For this purpose 2 intervals were selected, based on the results from Oostermeyer (2014). These intervals are located at depths between 8.90-10.60 m (zones C and D) and 33.98-36.23 m (zones G and H), respectively. Based on previous work, the first interval mainly contains kaolinite and iron oxides, while the second one mostly comprises chlorite, mixed-layer minerals and dolomite. Since very limited literature is available on weathering of talc, this study can contribute to a better understanding of talc weathering. Research questions that can be addressed are: what are the changes in mineralogical composition in the transition of parent rock to soil material? Can we find evidence for solid-state formation, neoformation or both? If these modes are identified, are their differences in the chemical and structural properties of the formed minerals, and what is their relation to the formation mode?

1.2 Objectives

The general objectives of this study are: [1] a quantitative study of the clay mineralogy to search for indications of kaolinite formation and explain the faith of talc, [2] matching the obtained mineralogical data with physico-chemical information, [3] trying to explain the behaviour of talc and chlorite in these soil environments and [4] collecting more information on the formation of kaolinite. For this purpose a combination of several techniques is used. Quantitative XRD will be the main method, supported by several other analytical techniques namely selective extractions, total elemental composition, transmission electron microscopy (TEM), Micro X-ray fluorescence (µXRF) and micromorphology. This thesis can also contribute to improvements of quantitative XRD modelling of (tropical) soil materials.

2. LITERATURE REVIEW

2.1 Geology

The Proto-Congo craton was formed by the assembly of six Archaean nuclei ca. 2.1 Ga ago. These nuclei have been affected by several Paleoproterozoic events, which lead to the assembly of the Columbia (Nuna) supercontinent at low to moderate latitudes between 1.9-1.8 Ga (De Waele et al., 2008). Rogers and Santosh (2002) suggested that the fragmentation of the supercontinent started roughly at 1.6 Ga until 1.2 Ga. However, Pesonen et al. (2012) believe that the break-up occurred much later around 1.12 Ga. The Proto-Congo Craton stabilized since the late-Paleoproterozoic (Tack et al., 2008). During the Meso- and Neoproterozoic, several intra-cratonic tectonic events took place, but they never led to the breakup of the craton (Tack et al., 2008).

The core for this study is taken in the Katangan belt or Lufilian arc (Figure 1). It is well known for hosting the largest Cu-Co deposits in the world (Kampunzu and Cailteux, 1999). The lithostratigraphic and stratigraphic divisions of the Katangan Supergroup are represented in Table 1. The Supergroup consists of a ca. 10.000 m thick sedimentary succession. Between 900 and 500 Ma, the Katangan basin developed as follows: along the southern margin of the Congo Craton, several rifting events occurred during the Neoproterozoic. First the opening of a continental rift basin (the first Katangan rift basin), associated with localised volcanism (in the Lufilian and Zambezi belts) leads to the deposition of fluvial clastic and lacustrine type sediments, which later form the rocks of the Roan Group (Buffard, 1988). This rifting phase (Roan rift) occurred between ca. 880-820 Ma (Johnson et al., 2007) and is related to the (early Neoproterozoic) extension of the Rodinia Supercontinent (Wendorff and Key, 2009).

The Roan Group rests unconformably on the Pre-Katangan basement. The age of the base of the Katangan Supergroup is still under discussion. The occurrence of pebbles of ca. 980 Ma Kibaran tin granites (Madi, 1985) in the basal conglomerate and of detrital cassiterite in the lowest Roan sediments (Jedwab, 1997) suggested an age of 980 Ma. The Nchanga Granite is the youngest pre- Katangan intrusion in the Paleoproterozoic to Mesoproterozoic basement. Therefore its age can be regarded as the start of the Katangan Supergroup. The age of 883 ± 10 Ma was established by Armstrong et al. (2005).

Table 1: Lithostratigraphy and stratigraphic divisions of the Katangan Supergroup (modified from Cailteux et al., 2007)

Super- Group Subgroup Formation Lithology group Biano Arkoses, conglomerates, argillaceous sandstones Ku 3 Sampwe Dolomitic pelites, argillaceous to sandy siltstones Ngule Kundelungu Kiubo Dolomitic sandstones, siltstones and pelites (formely Ku 2 Upper Mongwe Dolomitic pelites, siltstones and sandstones Kundelungu) Lubudi Pink oolitic limestone and sandy carbonate beds Ku Gombela Kanianga Carbonate silstones and shales Ku 1 Lusele Pink to grey micritic dolomite Kyandamu Petit Conglomérat (glacial diamictite) Monwezi Dolomitic sandstones, siltstones and pelites Bunkeya Dolomitic sandstones, siltstones and shales in northern Ng 2 Nguba Katete areas; alternating shale and dolomite beds ("Série (formely Récurrente") in southern areas Lower Kipushi Dolomite with dolomitic shale beds in southern areas Kundelungu) Kakontwe Carbonates Muombe Ng Carbonate shales and siltstones; "Dolomie Tigrée" at the Ng 1 Kaponda base Mwale Grand Conglomérat (glacial diamictite) Mwashya Kanzadi Sandstones or alternating siltstones and shales

(formely Kafubu Carbonaceous shales Upper Dolomitic shales, siltstones, sandstones, including angan Mwashya) Kamoya

R 4 conglomeratic beds and cherts in variable position Kat Kansuki Dolomites including volcaniclastic beds R 3.4 Mofya R Dolomites, arenic dolomites, dolomitic siltstones Dipeta 3.3 R 3 Argillaceous dolomitic siltstones with interbedded R 3.2 sandstone or white dolomite; intrusive gabbros Argillaceous dolomitic siltstones ("Roches Gréso- R.G.S. R3.1 Schisteuses") Roan Stromatolitic, laminated, shaly or talcose dolomites; Kambove R locally sandstone at the base; interbedded siltstones in R 2.3 the upper part Dolomitic shales containing carbonaceous horizons; Dolomitic occasional dolomite or arkose Mines shales Arenitic dolomite at the top and dolomite shale at the R 2 R 2.2 base; pseudomorphs after evaporite nodules and concretions Stromatolitic dolomite (R.S.C.), silicified/arenic dolomites Kamoto (R.S.F./D.Strat.), grey argillaceous dolomitic siltstone at R 2.1 the base (Grey R.A.T.) pseudomorphs after evaporites at the contact with R.A.T. Red argillaceous dolomitic siltstones, sandstones and R.AT. pelites ("Roches Argilo-Talqueuses"), the base of the R 1 R.A.T. sequence is unknown

Within the Roan Group, four subgroups were deposited: the Roches Argilo-Talqueuses (R.A.T.), Mines, Dipeta and Mwashya. The exact place where the core was taken is indicated in Figure 1. This figure also shows the presence of breccia comprising R.A.T., Mines and Dipeta Subgroups. Roches Argilo-Talqueuses indicate the presence of talc, although talc only occurs as a secondary mineral in oxidized zones. The primary mineralogy of fresh R.A.T. is mainly composed of diagenetic chlorite– dolomite and detrital and diagenetic quartz and chert, representing >85% of R.A.T. (Cailteux et al., 2005b and references therein). Wendorf (2000) observed an abrupt lithological transition between the R.A.T. and overlying Mines Subgroup, suggesting two different tectono-sedimentary cycles with two types of sources. This hypothesis was disproved by Kampunzu et al. (2005). They used mineralogical and geochemical data to conclude that both subgroups have the same provenance, source and depositional environment. Their sediments were deposited in a confined evaporitic environment. The Mines Subgroup contains dolomitic shales and siltstones with relicts of algal reefs and stromatolites (Van Langendonck et al., 2013). It was deposited in shallow intratidal to supratidal environments and hypersaline lagoons (Cailteux et al., 2005b). The overlying Dipeta Subgroup consists at the base of argillaceous and siliciclastic beds and at the top of carbonate beds. The Mwashya Subgroup comprises carbonaceous and dolomitic shales and dolomites (Cailteux et al., 2007).

At 765 ± 5 Ma a major uplift in the southern part of the Roan rift basin, ended the deposition of the Roan Group, and gave rise to the Nguba rift in which the Nguba Group was deposited (Wendorff, 2005). The Nguba Group records the second rifting stage resulting from the early Neoproterozoic extension of Rodinia (Wendorff and Key, 2009). The transition zone between the Mwashya Subgroup and the overlying Nguba Group contains dropstones in black shales, characteristic for the upper Mwashya Subgroup (Wendorff and Key, 2009). The basal unit of the Nguba Group is a diamictite horizon, known as the Grand Conglomérat. It is composed out of unsorted pebbly mudstone (Binda and Van Eden, 1972) and the thickness ranges from 300 to 400 m (Cahen, 1954). These glaciogenic sediments were deposited in an asymmetrical rift during the Cryogenian (Wendorff and Key, 2009). It occurred after 765 Ma and before 735 Ma (Key et al., 2002) and has been linked to the global Sturtian glaciation (Bodiselitsch et al., 2005). The glacial origin is confirmed by several authors, who found features such as facetted and striated clasts, clast sizes ranging from boulders to granules, poor sorting, varves, … (e.g. Binda and Van Eden, 1972; François, 1973; Dumont and Cahen, 1977). The Grand Conglomérat is capped by a carbonate unit, the Kakontwe limestone formation (Fm). Such a carbonate complex is also found overlying the younger Petit Conglomérat, which is discussed later (Kennedy et al., 1998).

The rift phase (extensional tectonic) was followed by compressional tectonics and related subduction. This resulted in the collision of the Congo and Kalahari cratons (Figure 1) and eventually the assembly of the Gondwana supercontinent (Figure 2) (Wilson et al., 1987). A series of major orogenic events, called the Pan-African orogeny which includes the Lufilian orogeny (550-520 Ma) (Johnson et al., 2005), are set in motion. The latter led to closure of the Katangan basin and deformation into a predominantly north-verging, low- to medium grade fold- and thrust belt, called the Lufilian arc (Kampunzu et al., 2000). This inversion of the tectonic regime created a foreland basin, which was being filled with the Kundelungu Group. The basal unit of the Group is a glacial

diamictite, similar to the Grand Conglomérat in the Nguba Group, reflecting the Marinoan/Varanger glaciation (ca. 640 Ma) (Hoffman, 2005). It is made up of (sandy) shales, and dolomites (Wendorff and Key, 2009). As mentioned above, a cap carbonate unit (Gombela Subgroup) is overlying the Petit Conglomérat. This is followed by deposition of the Ngule and Biano Subgroup. The Ngule Subgroup consists of the Mongwe, Kiubo and Sampwe Fm. The first 2 formations are folded, whereas the Sampwe Fm and overlying Biano Subgroup are not folded, tabular rocks (Batumike et al., 2007).

Figure 2: Representation of Gondwana assemblage with involved cratons (modified from Gray et al., 2007).

Associated with the Lufilian orogeny, large megabreccias were found, with a matrix-supported texture and containing rocks belonging to the R.A.T., Mines and Dipeta subgroups (Jackson et al., 2003). Jackson et al. (2003) suggest these breccias were formed by a combination of salt extrusion and orogenic shortening. The Roan Group once contained evaporites, which might even have formed diapirs in some regions. These were deformed and extruded as a result of the Lufilian Orogeny, making them susceptible for dissolution. The mobile mixture of dismembered Roan Strata (by the Lufilian orogeny) and evaporites formed an evaporate-breccia mush. The latter leached, carried and redeposited a large amount of metals, forming the well-known ore deposits of the Roan Group. Only a thin layer of evaporates occurred in the Lubumbashi area, therefore the effect of salt tectonics was limited.

Kampunzu and Cailteux (1999) recognized three major deformation phases during the Lufilian orogeny: (1) Kolwezian folding and thrusting (D1), (2) left-lateral strike–slip faulting (D2) and (3) the Chilatembo deformation (D3). The D1 is associated with transport to the north and induced the detachment of the Roan rocks from the basement caused by major shearing. The D2 affected the deformed and thrusted terranes and D3 caused an east–west folding perpendicular to the orientation of the Katangan belt (Batumike et al., 2007 and references therein). The resulting Lufilian arc consists of three structural zones: the Inner Lufilian (Synclinorial Belt), Middle Lufilian (Domes Region), and Outer Lufilian (Fold-and-thrust Belt) (Figure 3). The study area is situated in the Outer Lufilian (Jackson et al., 2003).

Figure 3: Representation of the three structural zones of the Lufilian arc (modified from Jackson et al., 2003).

2.2 Weathering and formation of clay minerals

Variations in clay mineralogy in a soil profile are related to weathering intensity. Both physical and chemical weathering processes are important for the formation of clay minerals. These processes are quite complex, controlled and impacted by many factors. Especially climatic conditions (temperature, rainfall), geological settings (morphology and lithology) and composition and hydrodynamics of the soil solutions are affecting these processes (Nahon, 1991). Hydrolysis and acidolysis were recognized as weathering processes for the formation of clay minerals (Pedro, 1964). Hydrolysis is the process where elements such as Mg, Fe, Al, even Si and others are rearranged and/or removed and operates when the soil solution pH ranges between 5 and 9.6. Total hydrolysis breaks down the mineral completely and leads to the precipitation of gibbsite and iron oxides and the formation of kaolinite. Limited hydrolysis leads to the formation of smectite (Pedro, 1982). Acidolysis occurs when the soil solution pH is lower than 5 and can be total or partial as well. The minerals are completely solubilized with total acidolysis and Al does not precipitate (due to the low pH). The fixation of Al in octahedral and interlayer positions of layer silicates is a result of partial acidolysis. Overall hydrolysis and acidolysis mainly occur in warm temperate, tropical zones and cold temperate climates, respectively (Wilson, 1999). Besides, Proust and Meunier (1989) noticed two weathering pathways in a soilscape, a vertical and lateral one. The weathering intensity of the latter one increases towards micro-passages. As a consequence, soil solutions will flow through the soil and weathered rocks, following the primary flow paths (paths of least resistance), affecting the weathering rates and products (Velbel, 1993).

2.2.1 Weathering of talc

Talc is one of the least complex 2:1 phyllosilicates belonging to the talc-pyrophyllite group, with the IV VI ideal formula [Si8] [Mg6] O20(OH)4 (Figure 4), indicating the absence of octahedral vacancies. It has a basal spacing of 0.94 nm. The trioctahedral sheet is composed out of one planar Mg octahedral sheet (O) sandwiched between two planar sheets of Si tetrahedral (T). The octahedral sheet is linked to the tetrahedral sheets by sharing the apical oxygen of the latter ones. The T-O-T layers are only held together by very weak Van der Waals interactions between the tetrahedral layers due to the absence of any charge on the sheets. This weak force is the reason for the perfect basal (001) cleavage and causes a hardness of 1 and a quality called slipperiness (Moore and Reynolds, 1997).

As mentioned, ideally, the mineral is electrostatically neutral because there is not any tetrahedral or octahedral substitution, layer charge or interlayer material. Because of these properties, [1] talc does not absorb water (hydromorphic characteristics) and [2] has a distinct slipperiness. However, observations have shown that the mineral differs slightly from the ideal structure. Some tetrahedral substitution does takes place, resulting in a small excess of negative charge which is mostly neutralized by octahedral substitutions. This very small or lack of layer charge explains the low chemical reactivity and corresponding low cation exchange capacity (CEC) (~2 cmol(+)/kg), which is for a large part developed by variable charges at the edges of talc crystals (Zelazny et al., 2002).

O2- Si4+

OH- 2+ 2+ Mg or Fe

Figure 4: Schematic projection of talc along the c-axis (modified from Van Ranst, 2013).

In contrast to the previous explanation, heterovalent substitutions (Fe3+  Fe2+ and Si4+  Al3+, respectively) cannot occur in the octahedral or tetrahedral sheet, but homovalent substitutions (Al3+  Fe3+ or Mg2+  Fe2+) are possible albeit rare (Meunier, 2005). Although Paquet et al. (1982) studied a weathered profile, developed on an olivine pyroxenite, in western Ivory-Coast, analyses of talc particles revealed that Al3+ substituted Si4+ in the tetrahedral sheet, the amount of Si4+ was situated between 3.8 and 4. Also Fe-substitution in the octahedra took place, with Fe values comprising between 0.12 and 0.67. These alterations probably add to the swelling tendency of interlayers.

The high resistance of talc against acid attack is attributed to the minimal tetrahedral substitution (Temuujin et al., 2003), their unstrained structure and the occurrence of only oxygen atoms between the layer surfaces (Yang et al., 2006). The latter induces hydrophobic characteristics, thought to be the cause for its slow leaching behaviour (Temuujin et al., 2002). Octahedral cations are more sensitive to acid attacks than tetrahedral ones, so Mg2+ ions should be removed from the octahedral layer in an early phase of leaching (Yang et al., 2006). This removal leads to the genesis of relative unstable micropores and eventually mesopores (Okada et al., 2002; MacKenzie et al., 2004). These pores will collapse, which in turn causes a decrease of the specific surface area.

Talc is frequently found in Mg-rich metamorphic and hydrothermal environments (Whitney and Eberl 1982; Evans and Guggenheim, 1988). It is considered as a secondary mineral by geologists but as primary by pedologists. The occurrence of talc in soils is relatively uncommon, attributed to their instability to weathering (Pérez-Rodrigues et al., 1996) and to the fact that formation by pedogenesis is unlikely (Zelazny et al., 2002). However there are some exceptions. Harris et al. (1984) identified the mineral in a highly leached (weathered) acidic soil, a Virginia Piedmont Ultisol. Here, iron oxides released from Fe-rich minerals through chemical and physical weathering, coated the talc crystals, shielding them from the soil solution and preserving them. Talc was also detected by León (1964) in several soils of Columbia and by Lynn and Whittig (1966) in reduced tideland sediments along the San Pablo Bay shore, California.

Two soil profiles (Alfic Dystric Eutrochrept and Xerochreptic Haploxeralf) derived from metamorphosed dolomitic parent material, from south-west Spain were investigated by Pérez- Rodríguez et al. (1996), both containing talc. Soil I is more intense weathered than soil II, which is confirmed by a depletion of primary minerals and bases, low pH and CEC. In the middle of this soil, they observed a relative increase of illite attributable to weathering of chlorite and talc. But talc persisted after intense pedogenetic processes, probably due to iron oxides coatings (Harris et al., 1984). The less weathered soil II, showed a homogeneous mineralogical composition, with talc mainly concentrated in the silt fraction, and thus less mineral alteration.

Different authors suggested that talc weathers to the dioctahedral smectite, nontronite (Besnus et al., 1976) or iron oxides (Colin et al., 1981; Nahon and Colin 1982; Paquet et al. 1982). Later on this was confirmed by Zelazny et al. (2002). But talc is Mg-rich, therefore we would expect it to weather to the trioctahedral smectite, saponite, which has predominantly Mg in the octahedral sheet. Alietti (1958) recognized for the first time interstratified saponite-talc, chlorite-saponite and chlorite- vermiculite, as weathering products of talc. Later on, talc/saponite was also found by Veniale and Van der Marel (1968) and Guenot (1970). Garvie and Metcalfe (1997) investigated the clay mineralogy of quartz-calcite veins, cutting altered andesite at Builth Wells, UK. In the veins, talc, saponite and corrensite occur together. As mentioned above, talc is typically associated with Mg-rich hydrothermal and metamorphic environments (Whitney and Eberl 1982; Evans and Guggenheim, 1988). Unlike saponite, which is often a product of hydrothermal altered basic rocks (Cowking et al., 1983). It could be that the availability of aluminium in basic rocks allows for the formation of saponite, while its relative absence in dolomites promotes talc formation. However, they also have been reported coexisting in a core, taken near a brine discharge vent in the Red Sea (Singer and Stoffers, 1987; Shau and Peacor, 1992). Garvie and Metcalfe (1997) gave the following explanation: the simultaneous occurrence of hydrothermal activity and increasing regional metamorphic temperatures resulted in the formation of quartz-calcite veins and of (secondary) chlorite and albite in the andesites. This was followed by the ascending migration of Mg-fluids, derived from dewatering of the mudrocks, through fractures and veins in the overlying lavas. The latter acted as an impermeable cap, which eventually fractured, attributed to fluid overpressures. This Mg-rich fluid eventually leads to the formation of talc and saponite. Several authors have found saponite in veins cutting altered basaltic rocks (Curtis, 1976; Cowking et al., 1983) but did not find the unique relationship between talc and saponite as described by Garvie and Metcalfe (1997). Anyway, the transformation of talc towards saponite is observed but the mechanisms still remain unclear.

2.2.2 Weathering of chlorite

Chlorite is a 2:1:1 phyllosilicate. Its 2:1 structure, unlike talc, has isomorphic substitutions that occur in both tetrahedral and octahedral sheets. The resulting excess negative charge is balanced by substitutions in the octahedral interlayer sheet, which can be either tri- (brucite-like) or dioctahedral (gibbsite-like). Most naturally occurring chlorites are trioctahedral in both octahedral layers (Figure IV 2+ 3+ VI 2+ 3+ 2+ 5). Their ideal formula is: [(Si,Al)8] [(R ,R )6] O20(OH)4 [R ,R )6(OH)12], in which R is a divalent cation (Mg2+ or Fe2+) and R3+ is a trivalent cation (Fe3+ or Al3+) (Velde, 1995). The basal spacing is

about 1.4 nm and CEC ranges between 10 and 40 cmol(+)/kg. They are common in igneous, metamorphic and sedimentary rocks.

2- O + Si4

- OH

2+ 2+ Mg or Fe

Figure 5: Schematic projection of trioctahedral chlorite along the c-axis (modified from Van Ranst, 2013).

Chlorite can be altered to vermiculite by weathering processes, which is called chlorite vermiculitization. It occurs by a solid-state mechanism (Banfield and Murakami, 1998). An important reaction in this process, is the oxidation and release of Fe2+ in the chlorite hydroxide sheet, decreasing the negative charge on the lattice. Although Proust et al. (1986) showed that chlorite vermiculitization could also be driven solely by the release of ferrous iron (Fe2+) from the 2:1 layer of chlorite rather than the oxidation of Fe2+. The release of iron due to progressive weathering is confirmed by the occurrence of free iron oxides, especially goethite (Herbillon and Makumbi, 1975; Ross and Kodama, 1976; Ross et al., 1982). The loss of Fe2+ is accompanied by subsequent leaching of Mg2+ from the interlayer brucite-like sheets. The pH values of soils are lower than the pKa of Mg2+ (11.6), therefore the brucite interlayers dissolve easily. The overall excepted theory is that Al (pKa of 5) remains largely immobile during the transformation of chlorite to vermiculite (Proust et al., 1986). Nevertheless, Aspandiar and Eggleton (2002) observed a loss of Al during the alteration of chlorite to vermiculite as well.

Several authors observed the transformation of chlorite to chlorite-vermiculite (C/V) or chlorite- smectite (C/S) regular mixed-layers (e.g. Ross et al., 1982; Meunier, 2005). Aspandiar and Eggleton (2002) even observed the transformation of chlorite to corrensite (a trioctahedral variety of regular, 50:50 mixed-layer chlorite-smectite and chlorite-vermiculite (Beaufort et al., 1997). However, these mixed-layers are not generally observed as a reaction product. The immediately transformation of chlorite to vermiculite was observed by Bain (1977), Adams and Kassim (1983) and Argast (1991). Two possible explanations were given: the examined chlorites have a high Fe2+ content, which is oxidized rapidly. This was confirmed by Ross (1975), who carried out experiments using brominated 11

water and demonstrated that chlorites with a high Fe2+ content (e.g. diabantite) weather immediately to vermiculites within a few weeks. On the other hand, Anand and Gilkes (1984) found a Fe-rich chlorite altered via mixed-layer minerals to vermiculite. On the contrary, chlorites with intermediate Fe2+ content transform to regular interstratified chlorite-vermiculites and low Fe2+ chlorites are more resistant to weathering (Ross, 1975; Ross and Kodama, 1976; Argast, 1991). A second explanation is the occurrence of Al-hydroxy interlayered vermiculite (Argast, 1991), which are associated with more acid environments. Therefore rapid removal of the interlayer of chlorite occurs. Consequently, the absence of interstratified minerals indicate rapid reaction rates.

2.2.3 Formation of kaolinite

Kaolinite is a 1:1 layer silicate belonging to the kaolin-serpentine group, with the ideal formula IV VI [Si4] [Al4] O10(OH)8 (Figure 6). It has a basal spacing of 0.72 nm. The dioctahedral mineral is composed of an octahedral sheet (occupied by Al) and tetrahedral sheet, linked by sharing the apical oxygens of the latter ones. Normally no isomorphic substitution occurs in the tetrahedral or octahedral sheet of kaolinites present in geological deposits, explaining its constant composition. In some kaolinites of tropical soils, vermiculitic, micaceous or smectitic layers were found, leading to a negative charge (Talibudeen and Goulding, 1983a), contradicting the absence of isomorphous substitution.

- O2 4+ Si

- OH

+ Al3

Figure 6: Schematic projection of kaolinite along the c-axis (modified from Van Ranst, 2013).

Kaolinite can be formed directly from the soil solution by a dissolution-crystallization mechanism (neoformation) or through a solid-state reaction of a primary phyllosilicate. When kaolinite is formed by neoformation, primary minerals are dissolved and their composing elements are released. From these elements, kaolinite can then precipitate (Środoń, 1980; De Coninck et al., 1986). The remaining elements will be leached out or precipitate as (hydr)oxides.

Solid-state transformation involves the formation of kaolinite from e.g. smectite via kaolinite- smectite (K/S) intermediates. By improvements in XRD and HR-TEM techniques, mixed-layers of kaolinite and smectites in soils have been recognized. They were first reported by Sudo and Hayashi (1956) and Altschuler et al. (1963) and later on by several other authors (Delvaux et al., 1990; Vingiani et al., 2004; Fisher and Ryan, 2006; Ryan and Huertas, 2009). These K/S minerals were

found in Mediterranean, dry tropical, subtropical and temperate soils (Delvaux and Herbillon, 1995) and even in moist tropical soils (Fisher and Ryan, 2006; Ryan and Huertas, 2009). However, in some case K/S remains undetected (e.g. Cuadros et al., 1994; Środoń, 1999) or is confused with a stacking of single end-members of smectite and halloysite or disordered kaolinite.

Tetrahedral sheet K/S: kaolinite-smectite K: kaolin-layer Octahedral sheet S: smectite-layer Interlayer cation IL: interlayer

Figure 7: The two reaction mechanisms of the solid-state reaction, smectite kaolinization. The interlayer Al-atoms indicate Al-hydroxy-interlayer complexes (modified from Ryan and Huertas, 2009).

The solid-state reaction, called smectite kaolinization, can occur in two ways (Figure 7), probably working in cooperation and eventually leading to the transition of smectite-layers (S) into kaolin- layers (K). The first mechanism includes stripping of one of the tetrahedral sheets of a smectite layer. Protons are added to the oxygens of the remaining octahedral sheet, forming OH groups. This induces the development of a halloysite-like layer, which expands when glycolated (Hughes et al., 1993; Dudek et al., 2006). Further, Mg and Fe are being replaced by Al in the octahedral sheet and Si replaces Al in the tetrahedral sheet (Amouric and Olives, 1998; Dudek et al., 2006; Ryan and Huertas, 2009) reducing layer charges and transforming the halloysite like layer into a kaolinite like layer. The second mechanism, involves two processes, which occur simultaneously. First, the inversion of a 13

tetrahedral sheet of a smectite layer and lateral transition of a smectite interlayer into a kaolin sheet (Schultz et al., 1971; Brindley et al., 1983). This reaction can only occur in the presence of Al- hydroxy-interlayers in smectites (HIS) (Wada and Kakuto, 1983). HIS mainly occur in acidic soil environments, where Al is mobile and low amounts of other base cations are present. The term kaolin is used because both processes form a mixture of kaolinite and halloysite like layers. Eventually the process will lead to the lateral transition of one smectite layer into two kaolin layers (Amouric and Olives, 1998). Besides these two mechanisms, dissolution-precipitation reactions of smectite-layers into kaolin-layers are taking place within the smectite or kaolin-smectite crystal as well. These mechanisms are solid-state reactions, and hence occur within the solid, in a cell- preserved manner. So the majority of the pre-existing solid remains preserved, indicating infiltration of fluids via edges into the interlayers. These reactions are thus driven by compositional changes of the soil solution towards high Al and low Si concentrations. Eventually smectite-layers are progressively replaced by kaolin-layers (Ryan and Huertas, 2009).

3. MATERIAL AND METHODS

3.1 Environmental setting

The study area is located near Bumaki, about 12 km to the north-northeast of Lubumbashi (Figure 8), Upper Katanga, D.R. Congo. The coordinates of the study location are 11°34'5''S latitude and 27°29'38''E longitude.

Figure 8: Location of the D.R. Congo within Africa and of Lubumbashi (red) within D.R. Congo (modified from “marysrosaries” and “webokapi”).

The study area belongs to the miombo ecosystem, which was defined by Frost et al. (1986) as: “Those tropical and some near tropical ecosystems characterised by continuous herbaceous cover consisting mostly of heliophilous C4 grasses and sedges that show clear seasonality related to water stress. Woody species (shrubs, trees, palms) occur but seldom form a continuous cover paralleling that of the grassy layer”. The evergreen forest is destroyed by dry season fires, with fires returning every 1.6-3 years (Frost, 1996), leading to a secondary grass vegetation.

The WRB classicfication classified the soil as a Ferralsol (Ngongo et al., 2009). These deeply weathered soils are typical for the humid tropics and characterized by their low chemical fertility. The latter results from intensive leaching of base cations and nutrients, leading to a field pH lower than 7. The clay fraction consists of low-activity clays (kaolinite) and sesquioxides.

Termite mounds are another typical feature of the miombo scenery, covering about 4.3 to 7.8 % of the land surface (Sys, 1957; Malaisse, 1974). These mounds are present with a density of 3 to 5 per hectare, reaching heights of up to 8 m and widths of 15 m (Sys, 1957; Malaisse, 1974). The mounds are built by Macrotermes falciger (e.g. Aloni, 1975; Goffinet, 1976). In general, termite activity affects several soil properties. They will lead to an enrichment of fine particles (Arshad, 1981; Jouquet et al., 2007), SOM (Ackerman et al., 2007) and exchangeable base cations (Mujinya et al.,

2010, 2013) compared with the adjacent soil. Termite mounds are resistant to weathering as well, therefore an enrichment in 2:1 clay minerals arises (Mujinya et al., 2013). Thus, termites (e.g. Macrotermes falciger) act as ecosystem engineers and modify their direct surroundings through their construction and feeding activities.

The climate of Lubumbashi is classified as Cws6 (warm temperate rainy climate with a dry winter and summer, the rainy and dry season each have a duration of 5 months) according to the Köppen classification system (Köppen, 1936). The rainy season, with a duration of 118 days on average, lasts from November until March. The dry season occurs from May till September, with July and August ever dry. The remaining months (October and April) are transitional. The average annual rainfall is 1270 mm with 717 mm and 1770 mm being the minimum and maximum. The mean annual temperature is approximately 20 °C with a minimum of 16 °C during July and a maximum of 23 °C in October (Malaisse, 1974; Mbenza, 1990).

3.2 Sampling strategy

Half of the termite mound at Bumaki was removed, making sampling possible. The above ground samples were taken by Dr. Hans Erens (Laboratory of Soil Science, Ghent University) and Prof. Basile Bazirake Mujinya (Laboratory of Soil Science, Ghent University and Laboratory of Soil Science, University of Lubumbashi). For the underground samples an appeal was made on the ‘Groupe Forrest International’, who kindly provided the equipment and staff for a ca. 40 m deep cored drilling. These drilled cores were transferred to wooden boxes with depth indication. A small subset of samples was transported to Belgium immediately, the remainder was shipped to Belgium in October 2012 with the help of the Belgian Ministry of Defense. In Belgium, the samples were rearranged into cardboard boxes of roughly 10 cm and labelled with their laboratory number and depth interval. Based on previous work (Oostermeyer, 2014), 2 intervals were selected. These intervals are located at depths between 8.90-10.60 m (zones C and D) and 33.98-36.23 m (zones G and H), respectively. Appendix 1 gives an overview of the selected samples and the executed analyses.

3.3 Physico-chemical analyses

All the physico-chemical analyses are carried out according to the ‘Students Lab Manual’ and ‘Procedures for Soil Analysis’ (Van Reeuwijk, 1993). A brief description is given below.

3.3.1 Texture, clay separation and saturation

The samples are tested for the presence of carbonates by using HCl. The test indicates the absence of calcite and aragonite, consequently there is no need to destroy them. However, presence of dolomite, which reacts only limited with HCl, was later on detected by XRD. The sand fraction is separated from the clay and silt fraction by wet sieving (63 μm sieve). Successive sedimentation is used to separate the silt and clay fractions, using Na2CO3 as a dispersion agent. After stirring, the silt and clay is allowed to settle for 8 hours. The supernatant fluid (containing clay sized particles) is then siphonated into another beaker. Flocculation of the clay fraction is achieved using NaCl and

gravitational settling, after which the clear supernatant is decanted. This procedure (dispersion, settling, siphonating) is repeated until no clay remains in suspension (clear supernatant after 8 hours). Afterwards, the clay fraction is washed with distilled water, ethanol and acetone until chloride free. The weight percentage is obtained by weighing the three size fractions.

Ca2+-saturated silt (2-63 µm) and clay (0-2 µm) fractions were obtained by repeated washing with 1N solutions of CaCl2 and Ca(OAc)2, respectively. The excess of the saturating solution was washed with acetone and alcohol until free of Cl-.

3.3.2 Soil acidity

The pH H2O is measured by mixing 6 g of soil with 15 ml of distilled water. The mixture is left for 2 hours to equilibrate and stirred occasionally. The pH-measurement takes place with a glass-calomel combination electrode. Readings are recorded when the value has stabilized (DIN ISO 90.93 10390, 2005).

3.3.3 Cation exchange capacity and exchangeable base cations

Cation exchange capacity (CEC) is measured by leaching samples with a 1N NH4OAc at pH 7 using a mechanical vacuum extractor. A 1g ball of filter pulp is pressed against the bottom of a syringe tube after which 2.5 g of fine earth material is added. Four blanks are also added. After saturating the samples with 25 ml, another 45 ml of NH4OAc is added and extracted overnight (16h period), after which extracts are transferred to a 100 ml volumetric flask. This extract is used to determine exchangeable base cations by ICP-OES.

Excess ammonium acetate in the sample is removed by washing the samples with ethanol repeatedly. Samples and pulp are then transferred to a distillation tube together with 6g of NaCl.

The distillation unit automatically adds 10 ml of H2O and 20 ml of NaOH to the tube. The distilled

NH4 is collected in 50 ml of a boric acid solution. This distillate is then back-titrated using 0.05 N HCl solution. The CEC and BS are calculated using the following formulas.

푠푎푚푝푙푒 푏푙푎푛푘 ( 푉 − 푉 ) × 퐶퐻퐶푙 × 100% 퐶퐸퐶 = 퐻퐶푙 퐻퐶푙 푊푠푎푚푝푙푒

퐶푎 + 푀푔 + 퐾 + 푁푎 퐵푆 = × 100 퐶퐸퐶

3.3.4 Selective chemical extraction

The method of Mehra and Jackson (1960) is used to determine dithionite-citrate-bicarbonate (DCB) extractable Si, Fe, Al and Mn. Approximately 250 mg of fine earth is used to which a solution of sodium-citrate and sodiumbicarbonate (at pH 7.3) is added. This solution is stirred and brought up to a temperature of 75°C. The free oxides of Fe and Mn are reduced using sodium-dithionite, added in

powder to the samples. After centrifuging, the supernatant solution is decanted and the concentrations of Si, Fe, Al and Mn are determined using ICP-OES.

3.3.5 Total elemental analysis

Total elemental chemical analysis is made using an alkaline fusion method (DIN ISO 90.60 14869-2, 2002). An amount of 400 mg of finely ground sample is transferred into a platinum crucible. The sample is then pre-ignited at 850°C for 30 min to avoid damaging the platinum crucible by reduction of organic matter during fusion. The sample is re-weighed after heating to obtain the loss of weight on ignition (LOI). The pre-ignited sample is then mixed thoroughly with 2 g of lithium-meta/tetra- borate powder in the platinum crucible and fused for 15 minutes at 950°C in a preheated muffle furnace. The formed flux is allowed to cool for one night, transferred to a 100 ml Teflon centrifuge tube and dissolved in 4% HNO3 (nitric acid). When the sample is completely dissolved, it is transferred to a 100 ml volumetric plastic flask and made to volume. Blanks and reference samples are also added to the sample series and analysed using ICP-OES.

The elemental composition of the silicate fraction of soil selected soil samples at different depths is recalculated after the subtraction of the elemental contents (EC) with DCB and LOI and the multiplication with a conversion factor. Calculations are considered correct, when the calculated total of the silicate fraction equals the total elemental composition. The following general formula was used.

푠푖푙푖푐푎푡푒 푓푟푎푐푡푖표푛 = (표푥푖푑푒 (퐸퐶) − 표푥푖푑푒 (퐷퐶퐵) − 퐿푂퐼) ∗ 푓푎푐푡표푟

푡표푡푎푙 (퐸퐶) 푓푎푐푡표푟 = ( ) 푡표푡푎푙 (퐸퐶) − 푡표푡푎푙 (퐷퐶퐵) − 퐿푂퐼

3.4 Quantitative X-ray diffraction (QXRD)

The aim of quantitative analysis is the determination of the relative weight fractions for each mineral phase present in the sample. This study combines XRD, which is the most commonly used method for this purpose, with total elemental analysis. The qualitative analyses of clay-minerals is relatively unambiguous, however attempts of quantitative analyses for complex mixed-layer assemblages usually occurring in soils, still remain semiquantitative despite the large amount of research (e.g., Środoń, 2006; Velde and Meunier, 2008; Ufer et al., 2012a; b). This is commonly ascribed to the variable structural and chemical characteristics (Środoń, 2006; Velde and Meunier, 2008; Hubert et al., 2012) as well to the tendency of preferred orientation of clay minerals (Dohrman et al., 2009). All of these factors influence peak shape, position and intensity, which may lead to errors in the final interpretation.

A lot of disturbing factors can be excluded or minimized by the used software and the most suitable adapted powder sample preparation. Several authors (e.g. Bish and Reynolds, 1989; Moore and Reynolds, 1997) have described and compared different most suitable adapted powder preparation

methods. They came to the conclusion that no single universal procedure can be utilized because the properties of the sample determine the success of the preparation technique. The software does not directly correct for particle statistics and preferred orientation. Normally, the crystals would be truly randomly oriented, if they were all spherical. Nevertheless platy and needle shaped crystals have the tendency to exhibit preferential orientation, enhancing the 001 reflections and weakening the hkl reflections. It is possible that minerals with different degrees of preferred orientation are present in the sample. Of course, corrections will be needed to reduce this effect (Dohrmann et al., 2009). Achieving this random orientation has been and still is a long-lasting problem. Obviously complete randomness will never be achieved, as perfect orientation is unattainable. But various methods are developed to minimalize the effect of preferential orientation. For instance, mathematical software corrections can be applied by the user. But different sample-mounting techniques can be used as well. Overall, two categories can be distinguished: (1) no pre-processing of the dry sample powder and (2) pre-treatment of the dry sample powder. The advantage of the former is that the sample remains uncontaminated and comprises front, back and side loading (Bish and Reynolds, 1989). For each of these loading methods, a different direction of filling the XRD mounts is applied. With front loading the powder is pressed in the cavity against the base of the sample holder with a glass slide. With a shear motion the excess powder is removed and the surface is made smooth. With back loading, the pressed opposite side is analysed (e.g. Klug and Alexander, 1974). For side loading, the powder is inserted from an open side into the holder (e.g. Moore and Reynold, 1997). Another often used technique is the Razor Tamped Surface (RTS) method, which is a form of front loading where the powder is tamped by a razor blade (Figure 9, section 3.4.1). Zhang et al. (2003) showed that the RTS method is most effective in reducing the preferred orientation and minimizing the packing density in the powder mount compared to the other loading methods. The RTS method is then followed by side loading, back loading and at least front loading. The second technique (sample powder pre-treatment) can be subdivided further into different categories. Namely (a) formation of spherical aggregates holding random orientated particles, e.g. freeze drying (Moore and Reynolds, 1997) and spray-drying (Hillier, 1999) (section 3.4.1), (b) addition of adhesives e.g. gelatin or smearing grease (Brown and Brindley, 1980), (c) addition of filling material, e.g. powdered cork (Wilson, 1987), glass beads (Jenkins et al., 1986) or thermoplastic organic cement (Hinckley, 1963) and (d) forming clay suspensions in acetone (Paterson et al., 1986).

Particle statistics or also crystallite statistics have been investigated by several authors (e.g. Smith, 1992; Elton and Salt, 1996). These are the statistical variations in the amount of particles joining the diffraction. The measured diffraction intensities are considered reproducible, when an accuracy of ca. 1-2 % relative is obtained. Typically only a small part of the total amount of particles contributes to the computed diffraction intensities. Elton and Salt (1996) have tried to establish the amount of these diffracting particles by theoretical and experimental procedures. This estimation is then used to calculate the particle statistics error. The error decreases with an increasing number of diffracting particles and with decreasing particle size (Smith, 1992). Thus, particle size plays an important role in achieving the demanded accuracy. There are several ways to reduce the size of the particles and thereby increase the number of diffracting crystallites, e.g. adjustments of the diffractometer, sample spinning (Smith, 1992), mechanical breakdown of the particle size by instruments (McCrone micronizing mill).

Once an XRD pattern is collected, it can be processed by software in three ways: (1) single peak/natural standard, (2) whole pattern/natural standard and (3) whole pattern/computed standard. The first method (e.g., Hillier, 2000; Środoń et al., 2001), measures the mineral concentration by using a selected single peak. This is measured upon a natural standard added as a spike to the sample. The second technique (e.g., Smith et al., 1987; Batchelder and Cressey, 1998) matches the measured XRD pattern with preregistered standard patterns for obtaining the intensities of the mixture components and thereby the mineral concentration. The last method (Taylor, 1991; Bergmann et al., 1998) fits the measured XRD pattern with a calculated one. This calculated pattern is based on data from a crystallographic database. A refinement procedure is then commonly used to adapt the calculated pattern (e.g. peak width, unit cell dimensions, preferred orientation and others).

The past two decades a lot of research was done to develop software programs for the characterization of clay minerals and their quantification in mixtures of several phases. The above mentioned uncertainty factors make the study of phyllosilicates in a quantitative manner a real challenge. These factors have to be addressed in the most adequate way as possible, while a limited set of parameters has to remain physically meaningful. The NEWMOD-family (e.g. Pevear and Schuette, 1993; Yuan and Bish, 2010), MLM2C/3C and derivatives (Plançon and Drits, 2000; Plançon, 2002; Plançon and Roux, 2010), Sybilla (Aplin et al., 2006), WILDFIRE (Reynolds, 1980; 1993), SiroQuant (Taylor, 1991; Taylor and Hinczak, 2006; Alves and Omotoso, 2009) and BGMN (Ufer et al., 2012a; b) are all examples of computer programs which were developed for the computation of XRD patterns for clay minerals.

Most models are developed for the calculation of XRD patterns of oriented mounts. They only focus on the 001 (basal) reflections and the stacking mode of different layers in the c-direction. NEWMOD (e.g. Pevear and Schuette, 1993; Yuan and Bish, 2010), MLM2C/3C (Plançon and Drits, 2000), MODXRSD (Plançon, 2002; Plançon and Roux, 2010), Sybilla (Aplin et al., 2006), … are such examples of such frequently used software programs. These models are also referred to as one-dimensional models. The main disadvantages of these models are: (1) the effects of transitional and rotational disorder in the a and b directions are disregarded and (2) quantification of non-clay minerals (quartz, goethite, hematite, feldspars,…) is not possible. The identification of the latter requires three- dimensional models. The original NEWMOD (Reynolds, 1985) is rather outdated but provided the basis/ standard for calculating XRD patterns for single and mixed-layer clay minerals. Many extension (e.g. NEWMOD3C, NEWMOD+, FITMOD and others) were made and some of them are still used (e.g. Ehlmann et al., 2012). FITMOD is a recent derivative and has an automatic parameter refinement. The NEWMOD-family only model two-component mixed-layer minerals, apart from NEWMOD3C. This restriction leads to the fact that minerals with three or more components cannot be modelled, e.g. interlayered smectites (occurring in several hydration states (Moore and Reynolds, 1997), with other layer types (e.g. illite). The MLM2C and MLM3C programs (Plançon and Drits, 2000) can calculate one-dimensional XRD patterns of mixed-layer minerals consisting respectively of two or three components. Later on, Plançon developed MODXRSD (Plançon, 2002; Plançon and Roux, 2010), which is over whole similar to MLM2C/3C but considers the clay minerals as particles

instead of crystals. These particles can contain distortions: cracks, joints and wedges. These lead to a broadening of the XRD-peaks but they are taken into account by the program. Besides, the program also has an automatic parameter refinement. Sybilla (Aplin et al., 2006) models multiple components mixed-layer minerals. It is the most complete and user-friendly software available at this time and has an automatic parameter refinement. However, the software was developed for the Chevron ETC and thus largely unmodifiable.

The calculation of XRD patterns for random oriented powder diffraction patterns can be done by WILDFIRE, Sybilla 3D, SiroQuant, DIFFaX, BGMN and others. WILDFIRE (Reynolds, 1980; 1993) is similar to NEWMOD (e.g. Pevear and Schuette, 1993; Yuan and Bish, 2010), except instead of the one-dimensional it uses a three-dimensional approach. Sybilla 3D is an extension of the one- dimensional Sybilla (Aplin et al., 2006). It was developed by Chevron Energy Technology Company (ETC). BGMN (Ufer et al., 2012a; b) is a Rietveld model developed by Bergmann J.. BGMN implements the recursive algorithm originally implemented in DIFFaX (Bergmann and Kleeberg, 1998, Treacy et al., 1991). However, these two models remain somewhat user unfriendly. SiroQuant (Taylor, 1991; Taylor and Hinczak, 2006; Alves and Omotoso, 2009) is a Rietveld-software program originating from cement industry applications which does not calculate patterns for mixed-layers but can re-use observed patterns from pure phases. The main disadvantage is that for each phase present in the sample, a pure reference material needs to be available. This method also does not give detailed information on the structure and disorder of the phase. Nevertheless it is a user friendly and fast approach that yields fairly reliable results. Because of this, SiroQuant was chosen and the workflow and parameters in the model will be discussed in further detail in section 3.4.3.

In this study, a combination of SiroQuant (Taylor, 1991; Taylor and Hinczak, 2006; Alves and Omotoso, 2009) and PyXRD (Dumon, 2014; Dumon et al., 2014; Dumon and Van Ranst, unpublished) was used for modelling the random bulk powder XRD patterns and (001) oriented clay or silt XRD patterns, respectively. Both software programs apply the whole pattern/computed standard method and are discussed in sections 3.4.3 and 3.4.4, respectively.

3.4.1 Random powder samples

In a perfect powder sample, all crystals are arbitrarily oriented and of sufficiently small size (< 10µm). When this is the case, each atomic plane has an equal chance of yielding a diffraction maximum compared to every other atomic plane in that powder. In reality it is often difficult to assure the complete random orientation of all particles in a powder when packed in a sample holder. This is by large the result of the non-spherical shape of crystals, resulting in preferred orientation when packed in a sample holder. This is especially true for clay minerals because of their perfect cleavage and sheet-like crystal habit. However, several techniques have been developed to overcome this problem, of which spray-drying has proven to be superior, as it prevents preferred orientation and results in highly reproducible data (Hillier, 1999; Jonas and Kuykendall, 1966; Hughes and Bohor, 1970). Spray-drying a slurry or suspension of a finely ground material will result in the creation of very fine droplets in the oven, which then dry forming spherical aggregates. Packing these spheres in a sample holder therefore eliminate any possible preferred orientation.

As a first step the sample’s crystal size must be reduced. Samples are first crushed in a mortar, so it passes a 50 m sieve after which the sample was transferred to a McCrone micronizing mill. Ethanol was added as grinding fluid according to an approximately 1:4 soil to liquid ratio. After milling for 3 minutes, the sample is poured into an airbrush jar and spray-dried at 10-15 psi in an oven operated at 60°C (Hillier, 1999). The spray dried powder is recovered on a ream of glossy paper, which is then poured in an XRD sample holder. The sample is then evenly distributed across the cavity, after which any excess is removed using a razor blade (Figure 9). During sample loading no pressure is exerted on the powder, except for gently tapping to ensure sufficient packing.

Moving razor from Moving loci of the the center to the two razor blade sides

Sample holder Scotch tape Glass slide

Figure 9: Illustration of the RTS method (modified from Zhang et al., 2003).

To each spray dried sample 5 % of zincite (ZnO) was added as internal standard before milling.

Instead of zincite, corundum (Al2O3) (Chung, 1974) could be used as an internal standard, however zincite provides stronger and more apparent reflections so less standard is needed which might otherwise mask phases with low concentrations.

The powder XRD-patterns were collected on a Philips X'PERT SYSTEM with a PW 3710 based diffractometer (Laboratory of Soil Science, Prof. Dr. E. Van Ranst), equipped with a Cu tube anode, a secondary graphite beam monochromotor, a proportional xenon filled detector, and a 35 position multiple sample changer. The incident beam was automatically collimated. The irradiated length was 12 mm. The secondary beam side comprised a 0.1 mm receiving slit, a soller slit, and a 1° anti-scatter slit. The tube was operated at 40 kV and 30 mA, and the patterns were collected in a θ-2-θ geometry from 3.00° 2θ onwards, at a step size of 0.020° 2θ, and a count time of 2 seconds per step.

3.4.2 Oriented samples

Smear slides imply the correct mineral proportions in a better way than sedimentation slides (glass slide method). The latter is sensitive to particle size segregation (e.g. Moore and Reynolds, 1997), leading to a misrepresentation of the quantitative results towards the finest minerals, attributable to the preferential exposure of these finer particles to the X-rays. The glass slide method, where the 22

suspension is pipetted on the glass slide, may also results in severe crusting and curling of the sample upon drying. With this in mind the use of smear slides was preferred over sedimentation slides. Smear slides are prepared for either clay or silt fractions depending on the amount of available material. A glass slide is inserted in a custom made holder (Figure 10) and about 400 mg of the soil sample and a few droplets of H2O are placed on the slide. The sample is then stirred until a sticky, homogenous paste is obtained. The paste is then smeared evenly over the glass slide with a reciprocating motion using another glass slide (working slide). Adding extra (half) droplets of H2O on the working slide or waiting for the paste to dry slightly, was often necessary to get optimal cohesion between the sample slide and the paste. The smearing continued until reaching a (visually) perfect smooth surface. In the resulting sample mount, layer silicates are preferentially oriented with their c-axis perpendicular to the glass slide as a result of the pressure applied during smearing. Glycol solvation of Ca2+ saturated samples was carried out in vacuum with glycol vapor during 24 hours.

Figure 10: Preparation of a smear slide.

The XRD patterns for smear slides were recorded on a Brüker D8 ADVANCE ECO system (Laboratory of Soil Science, Prof. Dr. E. Van Ranst) with a position-sensitive LYNXEYE XE detector and a Cu anode tube. The primary beam path is automatically collimated to 15 mm. To prevent the detector from being saturated at low angles a motorized anti-scatter beam knife is present on the sample stage as well. The tube was operated at 40 kV and 25 mA and patterns were collected in a θ–θ geometry from 0.50° 2θ onwards with a step size of 4° 2θ and a count time per step of 280 sec.

3.4.3 SiroQuant

Siroquant (Taylor, 1991) is a Rietveld full-profile fitting model, which can quantify phases from bulk XRD patterns. The experimental XRD pattern of the powder sample, is fitted with a calculated multiphase pattern. The software provides a mineral database with crystal structure data used to calculate the pattern. The discrepancy between the calculated and observed patterns is minimized using a least-squares refinement of different parameters. These include the phase scales, the instrumental zero point, unit cell dimensions, halfwidth parameters, preferred orientation, lineshape, phase structural parameters, line asymmetry and peak shape functions. Not all of these parameters can or need to be refined, e.g. phases present in low concentrations do not always allow refinement of the unit cell dimensions. Obviously, the scale parameter always needs to be refined.

The measured and calculated patterns are put on the same scale by refining the scale factor which also determines the final phase abundances. The 2 values from the measured XRD pattern deviate from those of the calculated pattern according to 2휃(푚푒푎푠) = 2휃(푐푎푙푐) + 푍퐸푅푂. This deviation is due to a diffractometer aberration, ZERO can be refined during the Rietveld analysis. The addition of a standard with well-known peak positions is of great help to distinguish variability in unit cell dimensions from this zero point deviation. The halfwidth parameters describe the shape of the peaks, usually only W is refined. If this is not sufficient to describe the peak shape, additionally V and U are included in the refinement as well. Most crystals are not spherical, hence a correction for their preferred orientation can be made if necessary. Other parameters are not discussed further in detail since they were not used in the refinement (Taylor and Hinczak, 2006).

The refinement procedure consist of several stages in which in general an increasing number of parameters is selected for refinement. Each stage consists of several cycles and has a damping factor to reduce overshoots or undershoots due to the least-squares refinement procedure. The background is removed manually in advance or automatically during the refinement. For most samples a manual background subtraction was performed as the automatic refinement often yielded unrealistic background curves, e.g. due to the presence of phyllosilicate reflections at low angles.

Amorphous or non-diffracting phases can still produce broad background humps. The amount of X- ray amorphous material can be quantified using an internal standard. This method can work, if the Rietveld method scales the sum of all phases to 100%. If an amorphous phase is present this will result in an overestimation of all phases, including the internal standard. It requires a good quantification of all other phases, and often removal of the background humps produced by the amorphous material.

The strength of the Rietveld method is the utilisation of a whole pattern-fitting algorithm, where all reflections are considered, even the overlapping ones. In this way, the uncertainty in the derived weight fractions and the effects of preferred orientation, primary extinction and nonlinear detection are minimized. Problems still arise when dealing with disordered clays, mixed-layer minerals and smectites. Their quantification still remains difficult and requires the combination of information from multiple sources. Therefore SiroQuant is used in combination with PyXRD, which is discussed in the next section.

3.4.4 Python X-ray Diffraction (PyXRD)

PyXRD version 0.6.4 is like SiroQuant a multi-specimen, full-profile fitting technique, written in Python 2.7 (Dumon and Van Ranst, 2015). It allows the simultaneous calculation of multiple saturation states. By optimizing a structural model for each clay species, the complete XRD pattern is calculated. The computer calculations are based on the algorithm developed initially by Drits and Sakharov (1976) and subsequent developments by Drits and Tchoubar (1990), Drits et al. (1997), Plançon (2002) and references therein.

A single project object can be generated in PyXRD. The created object bundles the AtomType, Phase and Component, Specimen and Mixture object(s). The latter ones are used to calculate X-ray diffraction patterns.

The AtomType object contains the physical constants, such as atomic weight, charge, atomic scattering factors (Waasmaier and Kirfel, 1995) for single ions, small and larger molecules. The single ions and small molecules are considered as spherical electron clouds, the larger molecules are regarded as a combination of several spheres. Consequently mineral components are obtained and used in the calculation of the XRD pattern.

With the Phase objects a one-dimensional XRD pattern of a (mixed-layer) mineral is attained by using the matrix formalism (Drits and Tchoubar, 1990). Markovian statistics involving the Reichweite key-concept (Drits and Tchoubar, 1990) is used for describing the order or disorder of the mixed- layered minerals e.g. R0, R1, R2... R0 is the chance to find a certain layer independent from the previous layers, R1 means it only depends on the first preliminary layer, R2 means it depends on the two closest preliminary layers and so on. The average coherent scattering domain (CSDS) distribution is a measurement for the crystallinity. It is described by a generic log-normal or a log- normal distribution, values are published in Drits et al. (1997). The σ* factor stands for the standard deviation of the Gaussian distribution of incomplete preferred orientation of the clay mineral (Dohrmann et al., 2009). The factor is expressed in degrees, normally varies between 6-12° and is considered the same for each phase present in the smear slide. Which contrasts to Zeelmaekers (2011) where σ* can differ significantly from phase to phase. In some cases σ* was set lower than 6° i.e. 3°. Another important factor is ∆c, describing anomalies in basal spacing along clay layers, caused by e.g. isomorphous substitution, wedges or cracks, weathering.

The Specimen objects are linked with the phases by Mixture objects and define instrumental and experimental parameters such as beam divergences, goniometer radius and sample length. The whole pattern can be shifted to a slightly higher or lower 2θ, using a reference peak (which is present in the analysed sample). Here the goethite (0.4183 nm) or the quartz (0.4257 nm) peak was used. In this way small variations in the vertical position of the sample-top were corrected. After setting the previous parameters a qualitative interpretation is done. First the clay minerals with a rational series of peaks are identified and added to the model. Due to a high scattering noise, the reflections of non-clay minerals (no 001-reflections) are excluded from the model. Subsequently the sensitive parameters (e.g. CSDS, octahedral-Fe content, basal spacing) are optimized to obtain a better fit.

The Mixture objects connect the phases and specimens. A single XRD pattern can be modelled in various ways as a result of the inherent variability of clay minerals. Hence a multi-specimen approach was suggested by Sakharov et al. (1999). This approach makes a comparison between air- dry (AD) and glycolated (EG) patterns possible, allowing the identification of swelling layers and add the correct (mixed-layer) phase for them. This is accomplished in the interface, where a table with columns and rows respectively represent the amount of specimens and phases. The specimens and the related phases are selected by the user, enabling the user to keep the unaffected phases

unaltered, while adding varying states of affected phases for an AD and an EG specimen. The sensitive parameters (CSDS, octahedral Fe-content, basal spacing) can be adjusted again.

After the Mixture is constructed, a refinement procedure can start. PyXRD incorporates different refinement algorithms. For a more detailed description of the functionality of PyXRD reference is made to Dumon and Van Ranst, unpublished; Dumon, 2014 and references therein.

3.4.5 Calculation of the mineral composition

To further constrain the results, a theoretical compositional range was calculated based on the obtained weight percentages for the different identified minerals. Since some of these minerals can have substitutions, minimum and maximum values were entered for the different major oxides for each phase. The theoretical composition obtained in this fashion is then compared with the measured composition. If a large discrepancy was observed this indicates a flaw in the XRD model, and adjustments were made accordingly. These adjustments include adding and/or removing phases and attempting to improve the fit if possible. This procedure results in a more reliable quantification.

3.5 High resolution transmission electron microscope (HR-TEM)

A small amount of air-dried sample (ca. 20 mg) is suspended in a 2 ml eppendorf tube using acetone for 48 hours. After centrifuging (5 minutes, 4000 RPM), the supernatant is decanted. After this initial dehydration step, the sample is further dehydrated during 8 hours using a 1:1 mixture of aceton and propylene-oxide and for another 8 hours using pure propylene-oxide.

Impregnation starts using Spurr-propyleenoxide mixtures: first a mixture of 1:3 (24 hours), then a 1:1 mixture followed at last by a 3:1 mixture. Each of these impregnation steps lasts 24 hours. The final impregnation is done using pure Spurr for 62 hours in an embedding mould. Hardening of the resin is achieved by heating the samples to 70°C for 8 hours.

The obtained block of resin is then trimmed using a Leica EM-trimmer. After trimming semi-thin slices (approx. 500 nm thick) are cut using a diamond knife ultra-microtome (Leica UC7). These slices are then transferred to a glass slide to check (using a binocular microscope) if sufficient sample was cut. If so ultra-thin slices (approximately 70-80 nm thick) are cut using the same ultra-microtome and transferred to a carbon coated copper grid. These grids are then coated a second time with carbon to improve stability. Low-resolution TEM images were made using a JEOL JEM 1010 at 60 keV, high resolution TEM images were made using a Cs-corrected transmission electron microscope (JEOL FE2200) operated at 200 keV.

3.6 Micro X-ray fluorescence (XRF)

Micro X-ray fluorescence (µXRF) is a non-destructive analysing technique with a wide range of applications in various disciplines including clay mineralogy and evaluation of multi-layered coatings. Here µXRF was used as an elemental analysing and mapping technique. The X-ray excitation beam is focussed on a small spot of interest in the sample, inducing characteristic X-ray fluorescence

emission of that spot leading to an elemental characterization of that spot. Tack et al. (2014) describes the executed procedures in more detail. An energy dispersive (ED) pnCCD detector and SLcam are used for full-field fluorescence mode X-ray absorption near-edge structure spectroscopic (XANES) measurements. This results in spatially resolved information on the chemical state for millimeter sized sample areas. For each measurement (20 µm x 20 µm), a XANES profile is received for each of nearly 1300 points.

3.7 Micromorphological analysis

The samples were dried in the oven at 50-60°C and impregnated under vacuum with an unsaturated polyester resin. Subsequently they were cut into slices, polished to a thickness of 25 to 30 µm and fixed on a glass slide. Some of the uncovered thin sections were analysed by the µXRF. Afterwards they were studied with a polarizing microscope and described according to the terminology of Stoops (2003).

4. RESULTS

4.1 Morphological properties

This section is based on the visual core descriptions made by Oostermeyer (2014). Below a summary of these observations is given for the studied intervals. A detailed overview can also be found in Appendices 2 and 3.

The first selected interval (8.90-10.60 m) comprises the two zones C and D. Zone C is a gravel layer and extends from 7.43 m to 9.30 m depth. The colour of this zone is reddish yellow (dry: 7.5 YR 6/8, moist: 7.5 YR 5/8) to red (dry: 2.5 YR 5/8, moist: 2.5 YR 4/8). From the lower boundary of this zone (8.90-9.30 m), 3 samples were taken for further analysis. Zone D (9.30-22.80 m) is marked by the first intercalations of talc at different depths. The colour of the matrix is mainly brownish yellow (dry: 10 YR 6/8, moist: 10 YR 6/8) and pink (dry: 5 YR 7/3, moist: 10 R 6/4). Samples were taken from the uppermost part of the zone (9.30-10.60 m).

The second selected interval (33.98-36.23 m) contains zones G and H. Zone G (32.68-34.80 m) is characterized by very pale brown (dry: 10 YR 7/4, moist: 10 YR 6/6) and brownish yellow (dry: 10 YR 6/8, moist: 10 YR 6/8) colours. A dark yellowish brown (dry: 10 YR 3/4, moist: 10 YR 3/3) colour can be observed in the uppermost part of the interval (33.98-34.10 m). Between 34.36-34.46 m and 34.65-34.80 m, white (dry: 10 YR 8/0, moist: 10 YR 8/1) material, which feels like talc can be observed. Samples were taken from 33.98 m to 34.80 m depth for further analysis. Zone H ranges from 34.80 m to 37.32 m depth. The colours of this zone are pale olive (dry: 5 Y 6/3, moist: 5 Y 5/3) and light olive gray (dry: 5 Y 6/2, moist: 5 Y 6/2) and greenish gray (dry: 10 GY 5/1, moist: 10 GY 4/1). Samples were taken within the first two meters of the zone for more detailed study.

4.2 Physico-chemical properties

The physico-chemical properties of interval 1 and 2 are presented in Tables 2 and 3, respectively. The weight percentages (wt%) of sand in interval 1 are the highest for zone C. The clay fraction increases towards zone D and stays relatively constant throughout the interval, except from 10.55 m to 10.60 m (13/165) which displays the lowest clay content. The silt fraction increases until a depth of 9.64-9.80 m, followed by a decrease until ca. 10.30 m. From 10.55 m to 10.60 m (13/165) the highest silt fraction (80 %) occurs, contradicting the previous decreasing trend.

Interval 2 is almost devoid of clay, except for a few percentages in the upper part. From 33.98 to 34.50 m depth, the silt fraction shows a strong increase to almost 80%. Subsequently, the silt content diminishes again until a minimum of 19 % is reached at ca. 35 m. The next interval at 35.25- 35.40 m depth is again very silty with about 47 % of silt. In the remainder of the interval the silt content stays more or less constant around 30 %. Logically, the amount of sand is inversely proportional to the amount of silt because the clay fraction is very low or even absent.

Table 2: Physico-chemical properties of selected samples of interval 1.

Sample Depth interval Texture (wt%) pH-H O Exchange complex (cmol(+)/kg) BS Zone 2 number (m) Sand Silt Clay (1:2.5) Ca2+ Mg2+ K+ Na+ CEC (%) 13/152* 8.90-9.00 48 43 9 5.32 0.31 0.49 0.05 0.12 13.81 7 C 13/153 9.00-9.10 n.d. n.d. n.d. 5.15 0.52 0.82 0.13 0.09 6.72 23 13/154* 9.10-9.30 30 58 12 6.08 0.75 1.06 0.26 0.11 6.97 31 13/155* 9.30-9.40 22 59 19 4.59 0.06 0.70 0.03 < 0.01 19.17 4 13/156 9.40-9.50 n.d. n.d. n.d. 4.63 0.13 0.77 0.06 0.09 5.89 18 13/157* 9.50-9.64 8 66 26 4.53 0.10 2.22 0.06 0.11 10.07 25 13/158* 9.64-9.80 7 71 22 4.48 0.10 2.42 0.07 0.10 10.26 26 13/159 9.80-9.93 n.d. n.d. n.d. 4.68 0.11 3.21 0.07 0.10 12.19 29 D 13/160* 9.93-10.10 20 64 16 4.96 0.06 1.51 0.06 0.08 6.09 28

13/161 10.10-10.20 n.d. n.d. n.d. 4.66 0.15 0.93 0.06 0.17 4.12 32 13/162* 10.20-10.30 29 56 15 5.05 0.29 1.20 0.13 0.22 14.97 12 13/163* 10.30-10.40 23 59 18 4.73 0.13 2.32 0.11 0.14 7.20 38 13/164 10.40-10.55 n.d. n.d. n.d. 5.20 1.81 3.96 0.14 0.32 9.58 65 13/165* 10.55-10.60 15 80 5 4.86 0.09 3.68 0.07 0.03 8.28 47 n.d. : not determined * : successive sedimentation for clay segregation and texture analysis

BS : base saturation

The wt% of sand, silt and clay were plotted on the USDA soil textural triangle for both intervals (Figure 11). The samples of the first interval all plotted in the silt loam texture class, with two exceptions. Sample 152 belongs to the loam class and sample 165, falls on the boundary between silt and silt loam. In interval 2, the texture classes silt loam, sandy loam and loamy sand are recognized.

Figure 11: USDA soil textural triangle, bleu represents the samples and numbers from interval 1 and red the samples and numbers from interval 2.

Most of the CEC values of interval 1 stay below or around 10 cmol(+)/kg. But four exceptions could be recognized, where the CEC values are lying between 12.19 cmol(+)/kg and 19.17 cmol(+)/kg. The exchangeable Ca2+ of the first interval is not very variable (average of 0.33 cmol(+)/kg), with the exception of a peak value of 1.81 cmol(+)/kg from 10.40 m to 10.55 m depth. The values of exchangeable Mg2+ show a general increase until a depth of 9.80-9.93 m, where a value of 3.21 cmol(+)/kg is attained. Then, a decrease over a small distance is observed, which is followed once more by an increase. The exchangeable Na+ (average of 0.13 cmol(+)/kg) and K+ (average of 0.09 cmol(+)/kg) are very low compared to Mg2+ (average of 1.81 cmol(+)/kg) and relatively uniform. The base saturation (BS) of interval 1 is relatively stable, apart from the interruption of some higher (65 and 47 %) and lower values (7 and 4 %).

Table 3: Physico-chemical properties of the selected samples of interval 2.

Sample Depth interval Texture (wt%) pH-H O Exchange complex (cmol(+)/kg) BS Zone 2 number (m) Sand Silt Clay (1:2.5) Ca2+ Mg2+ K+ Na+ CEC (%) 13/355 33.98-34.10 n.d. n.d. n.d. 6.45 0.74 1.68 0.07 0.16 3.29 80 13/356* 34.10-34.23 53 46 1 6.39 2.50 5.92 0.15 0.14 9.28 94 13/357* 34.23-34.36 21 73 6 6.54 12.96 37.58 0.53 0.13 53.00 97 G 13/358* 34.36-34.46 16 77 7 6.78 14.84 41.91 0.68 0.13 54.30 >100 13/359* 34.46-34.50 17 79 4 6.83 14.03 42.34 0.68 0.12 58.09 98 13/360* 34.50-34.65 45 54 1 6.80 3.73 14.38 0.18 0.12 20.83 88 13/361 34.65-34.80 n.d. n.d. n.d. 7.73 4.56 29.48 0.12 0.05 39.25 87 13/362 34.80-34.90 n.d. n.d. n.d. 8.21 6.45 16.66 0.08 0.15 14.31 >100 13/363* 34.90-35.09 81 19 - 8.14 7.48 14.75 0.09 0.25 11.65 >100 13/364 35.09-35.25 n.d. n.d. n.d. 8.05 5.92 14.92 0.07 0.20 14.45 >100

13/365* 35.25-35.40 53 47 - 7.78 6.31 54.27 0.23 0.18 57.63 >100 13/366 35-40-35.56 n.d. n.d. n.d. 8.17 5.10 22.89 0.10 0.15 23.50 >100 H 13/367* 35.56-35.70 71 29 - 8.17 8.03 21.71 0.10 0.17 22.01 >100 13/368 35.70-35.85 n.d. n.d. n.d. 8.21 7.76 30.37 0.20 0.15 31.95 >100 13/369* 35.85-35.97 70 30 - 8.83 6.04 14.70 0.03 0.03 13.29 >100 13/370 35.97-36.13 n.d. n.d. n.d. 8.55 5.33 11.58 0.08 0.14 10.11 >100 13/371* 36.13-36.23 76 24 - 8.58 7.88 12.25 0.18 0.15 10.27 >100 - : absent n.d. : not determined * : successive sedimentation for clay segregation and texture analysis BS : base saturation

The pH-H2O values in interval 1 vary around an average of 4.92, with a minimum and maximum value of 4.48 and 6.08, respectively. Throughout interval 2, the pH-H2O increases from 6.45 towards 8.58, a difference of more than two pH-units.

The CEC values of interval 2 vary enormously. The lowest value (3.29 cmol(+)/kg) occurs in the uppermost sample of the interval. CEC values exceeding 50 cmol(+)/kg are established in two units (34.23-34.50 m and 35.25-35.40 m). Aside from these, two intervals exceeding a value of 30 cmol(+)/kg could be recognized as well, from 34.65 to 34.80 m depth and from 35.70 to 35.85 m depth. Below this depth CEC varies around a mean of 15.60 cmol(+)/kg. The exchangeable Ca2+ of interval 2 varies between 0.74 cmol(+)/kg and 14.84 cmol(+)/kg. First an increasing trend is recognized until a value of 7.48 cmol(+)/kg is reached. However, this trend is interrupted by a unit with values ranging between 12.96 cmol(+)/kg and 14.84 cmol(+)/kg (34.23-34.50 m), which are the highest values of Ca2+ throughout the interval. The trend is followed by a decrease with values going down to 5.10 cmol(+)/kg from 35.40 to 35.56 m depth. Below this depth exchangeable Ca2+ increases again and varies around an average of 7.01 cmol(+)/kg. When considering the exchangeable Mg2+ of the second interval, some units with exceptional high values are recognized. These units occur from 34.23 to 34.50 m depth, from 34.65 to 34.80 m depth, from 35.25 to 35.40 m depth and from 35.70 to 35.85 m depth, with values ranging between 29.48 cmol(+)/kg and 54.27 cmol(+)/kg. The other Mg2+ values vary around a mean of 15.98 cmol(+)/kg, apart from the uppermost two values (33.98- 34.10 m and 34.10-34.23 m). The exchangeable Na+ (average of 0.14 cmol(+)/kg) and K+ (average of 0.21 cmol(+)/kg) are very low compared to Mg2+ (average of 22.79 cmol(+)/kg) and relatively uniform. The BS of interval 2 shows a general increasing trend with depth, eventually a BS of more than 100 % is reached.

4.3 Total elemental composition

Table 4 represents the total elemental composition of the selected samples at different depths. The

SiO2 wt% of interval 1 mostly increases with depth. When considering the SiO2 wt% of interval 2, a decreasing trend with depth is identified, with the exception of one value (60.19 %). A remarkably high value of 80.06 % is recognized in the uppermost sample of the second interval. Both the highest and lowest reported values for SiO2 (80.06 % and 26.91 %, respectively) occur in the second interval.

The Al2O3 wt% of interval 1 increases until a depth of 9.64 m, corresponding to a value of 24.03 %, followed by a decrease. Within interval 2, a unit with a mean value of 12.74 Al2O3 % (34.23-34.65 m) is distinguished. The variation of the other values is relatively small, between 1.33 and 5.76 %.

A decreasing trend of the wt% of Fe2O3 in interval 1 is observed, until a value of 8.12 % in the lowest sample of the interval is reached. The Fe2O3 wt% variability of interval 2 is quite small, the values fluctuate between 2.05 and 4.89 %. When considering the wt% values of MnO, no trends are noticed in both intervals. The values remain more or less constant. However, two peak values of 1.30 and 0.97 % occur in the uppermost part of interval 1. The MgO values of interval 2 are higher than those of interval 1. In interval 2, a clear increasing trend could be distinguished with a maximum value of 18.98 %. Such a trend could also be recognized in interval 1, but is less evident.

Table 4: Total elemental composition (in wt%) of selected samples at different depths.

Sample Depth interval Zone SiO Al O Fe O MnO MgO CaO Na O K O TiO P O LOI Total number (m) 2 2 3 2 3 2 2 2 2 5

13/152 8.90-9.00 50.49 10.17 28.84 1.30 0.49 0.03 0.06 0.43 0.46 0.09 8.19 100.54 C 13/154 9.10-9.30 52.12 13.70 21.26 0.97 0.98 0.03 0.05 0.71 0.67 0.06 8.41 98.96

13/155 9.30-9.40 58.07 15.22 10.53 0.17 4.78 < 0.01 0.06 0.80 1.08 0.02 7.39 98.11

13/157 9.50-9.64 46.12 24.03 11.84 0.02 2.76 < 0.01 0.13 0.43 2.18 0.02 10.55 98.09 13/158 9.64-9.80 46.67 23.08 12.46 0.02 4.71 < 0.01 0.15 0.42 1.91 0.02 10.20 99.64

Interval 1 Interval D 13/160 9.93-10.10 59.70 18.06 8.59 0.20 2.70 < 0.01 0.09 2.07 0.97 0.01 7.15 99.55 13/162 10.20-10.30 63.61 15.84 7.98 0.15 3.71 0.03 0.08 2.89 0.84 0.02 4.88 100.01 13/163 10.30-10.40 56.43 18.08 9.29 0.10 4.32 < 0.01 0.09 2.55 1.25 0.02 7.09 99.23 13/165 10.55-10.60 62.91 16.10 8.12 0.09 3.53 0.02 0.12 1.24 1.11 0.01 6.28 99.55 13/356 34.10-34.23 80.06 3.08 2.40 0.13 9.64 0.11 0.06 0.29 0.19 0.02 3.37 99.34 33 13/357 34.23-34.36 52.69 13.19 4.89 0.17 11.31 0.41 0.25 3.41 0.83 0.03 13.34 100.51

G 13/358 34.36-34.46 52.15 12.27 3.81 0.14 13.43 0.71 0.22 2.04 0.96 0.21 14.44 100.39

13/359 34.46-34.50 50.05 13.59 3.00 0.07 12.65 0.88 0.08 2.61 1.16 0.35 15.18 99.62 13/360 34.50-34.65 60.19 11.91 4.45 0.21 11.62 0.37 0.10 3.43 0.83 0.21 7.04 100.37 13/363 34.90-35.09 49.92 1.33 2.05 0.07 13.28 13.00 0.06 0.10 0.09 0.07 19.56 99.53 interval 2 interval 13/365 35.25-35.40 44.54 5.76 2.73 0.04 18.98 6.94 0.04 0.27 0.26 0.04 18.23 97.82 H 13/367 35.56-35.70 45.57 2.04 3.12 0.09 14.25 13.15 0.05 0.18 0.13 0.05 21.04 99.67 13/369 35.85-35.97 26.91 3.45 3.56 0.13 18.45 16.88 0.05 0.32 0.23 0.06 28.26 98.32 13/371 36.13-36.23 26.96 5.58 2.73 0.13 18.03 16.40 0.09 1.61 0.37 0.12 27.68 99.70 LOI : loss on ignition (weight loss after heating at 1000°C)

The wt% values of CaO of interval 1 are very low, with even 5 out of the 9 measurements under the detection limit. Within the second interval, two units could be separated. The first unit with an average wt% of CaO of 0.50 %. This unit corresponds to zone G, at a depth of 34.65 m. The second unit corresponds consequently with zone H. The values of unit 2 are much larger than those of unit 1. The second unit has an average of 13.27 % and contains a minimum of 6.94 %, which differs significantly from the other values within that unit.

The Na2O contents of both intervals are more or less constant. The mean of interval 1 and 2 amounts 0.09 and 0.10 %, respectively. Two peak values (0.25 and 0.22 %) can be differentiated in interval 2.

Interval 1 can be divided into 2 units based on the K2O content. Unit 1 has an average of 0.56 % and occurs between a depth of 8.90 m and 9.80 m. The second unit appears from 9.93 m until 10.60 m

(average of 2.19). Interval 2 shows within zone G an increasing trend in K2O, followed by an alternation of decreases and increases.

The TiO2 content of interval 1 increases up to 2.18 %, followed by a decrease until 0.84 % and again followed by an increase. Within interval 2, the wt% of TiO2 first increases to 1.16 %, followed by a decrease and an increase again.

The P2O5 values are low throughout both intervals, but still small trends could be identified. Interval 1 shows a decreasing trend. Interval 2 demonstrates an opposite trend, apart from the unit at a depth between 34.36 and 34.65 m (average of 0.26 %). Loss on ignition (LOI) represents the weight loss after heating at 1000°C. The LOI of interval 1 does not display a real trend, decreases and increases alternate, with a maximum and minimum of 10.55 and 4.88 %, respectively. For interval 2, LOI generally increases with depth, except for one value (7.04 %) corresponding to the deepest sample in zone G. Another low value (3.37 %) occurs at the top of zone G, namely the minimum of the second interval.

4.4 Weathering indices and molar ratios

The weathering indices and molecular ratios of selected soil samples at different depths are given in Table 5. The weight percentages obtained by the total elemental analysis were rescaled to 100 %. These weights were used to calculate the molecular weights, from which the weathering indices and molecular ratios (Al2O3/MgO and TiO2/MgO) were computed. The weathering indices: Chemical

Index of Alteration (CIA) (Nesbitt and Young, 1982), the ratio of SiO2 and Fe2O3 (SF) (Jenny, 1941) and Ruxton ratio (R) (Ruxton, 1968) were calculated using the following formulae.

퐴푙2 푂3 퐶퐼퐴 = ∗ 100 (퐴푙2푂3 + 퐶푎푂 + 푁푎2푂 + 퐾2푂 )

푆푖푂2 푆퐹 = 퐹푒2푂3

푆푖푂2 푅 = 퐴푙2푂3

The indices and ratios are based on the mass transfer of labile (mobile) elements. As a result of weathering, the mobile elements will be removed inducing a relative enrichment in the less mobile elements. Thus with increasing weathering, the R and SF will decrease due to the immobility of Al2O3 and Fe2O3 and the mobility of SiO2. The CIA and the other two molecular ratios present an opposite trend, i.e. an increase with increasing weathering, related to the retention of Al2O3 and TiO2 and the removal of Ca, Mg, Na, K. The R of interval 1 varies between 3.00 and 8.50. Interval 2 contains three extreme values, 44.15, 63.63 and 37.87. The first two, represent the first sample in zone G and H, respectively. The third less extreme value occurs at a depth between 35.56 and 35.70 m. Within zone G, the R fluctuates around a mean of 7.20, apart from the extreme value of 44.15. In the lower part of interval 2 (zone H), the R ranges between 8.00 and 13.50, without taken into account the two extremes. The SF values of interval 1 lie below 21.50 and are all smaller than those of interval 2, without considering the SF (20.08) of interval 2 at a depth between 35.85 and 35.97 m. Within interval 2, two peak values could be identified and coincide with the first two extremes of the Ruxton ratio. The remaining values range between 20.00 and 45.50. The CIA values of interval 1 are higher than those of interval 2 and vary between 82.50 and 97.50. In interval 2, the values decrease with depth and a minimum of 5.29 occurs between 34.90 and 35.09 m. Interval 1 shows a decreasing trend of Al2O3/MgO with depth. Within interval 2, the values are relatively stable, with the exception of the last four samples of zone G. Their Al2O3/MgO values fluctuate around an average of 0.41. The values of the ratio TiO2/MgO could be neglected for interval 2, contrary to interval 1 which presents an increasing trend towards the top.

Table 5: Weathering indices and molar ratios of selected samples at different depths.

Depth Molecular ratios and weathering indices Sample Zone interval number (m) R SF CIA Al2O3/MgO TiO2/MgO 13/152 8.90-9.00 8.43 4.65 94.20 8.26 0.47 C 13/154 9.10-9.30 6.46 6.52 93.79 5.52 0.34

13/155 9.30-9.40 6.48 14.66 94.05 1.26 0.11

13/157 9.50-9.64 3.26 10.35 97.22 3.44 0.40 13/158 9.64-9.80 3.43 9.95 97.05 1.94 0.20

Interval1 D 13/160 9.93-10.10 5.61 18.48 88.28 2.64 0.18 13/162 10.20-10.30 6.82 21.20 82.69 1.69 0.11 13/163 10.30-10.40 5.30 16.15 86.17 1.65 0.15 13/165 10.55-10.60 6.63 20.59 91.02 1.80 0.16 13/356 34.10-34.23 44.15 88.77 83.77 0.13 0.01 13/357 34.23-34.36 6.78 28.65 73.12 0.46 0.04 G 13/358 34.36-34.46 7.22 36.36 76.04 0.36 0.04

13/359 34.46-34.50 6.25 44.36 74.87 0.42 0.05 13/360 34.50-34.65 8.58 35.96 72.39 0.41 0.04 13/363 34.90-35.09 63.63 64.75 5.29 0.04 0.00 Interval2 13/365 35.25-35.40 13.12 43.43 30.76 0.12 0.01 H 13/367 35.56-35.70 37.87 38.86 7.79 0.06 0.00 13/369 35.85-35.97 13.24 20.08 9.98 0.07 0.01 13/371 36.13-36.23 8.20 26.22 14.97 0.12 0.01 35

4.5 Total elemental composition of the silicate fraction

The elemental composition of the silicate fraction of selected soil samples at different depths is given in Table 6. The SiO2 content of interval 1 decreases downwards up to 56.08 %, followed by an general increase. The minimum and maximum of interval 1 are 82.21 % and 56.08 %, respectively. Beside the values 66.18 and 63.07 %, interval 2 presents a decreasing trend with depth. The first and last SiO2 values of interval 2 represent the maximum (84.32 %) and minimum (37.41 %), respectively.

In interval 1, two units with values above 20 % Al2O3 are recognized, occurring at depths between 9.50-10.10 m and 10.30-10.40 m. The remaining values display a mean of 16.61 %. Within interval 2, a similar unit is suggested from 34.23 until 34.65 m, but the values range between 13 and 16 %. This unit comprises almost the entire zone G, except for the top value. Zone H of interval 2 is characterized by lower Al2O3 values. They vary between 1 and 8 %. The Fe2O3 content of interval 1 and 2 is quite constant. Interval 1 displays, a zero value at a depth between 9.30 and 9.40 m. In interval 2, the lowest (0.99 %) and highest value (4.30 %) of Fe2O3 are consecutively. When considering the MnO values of interval 1, a decreasing trend is established. In contrast to interval 1, interval 2 increases towards the bottom. Based on the MgO content, interval 1 could be separated into 2 sections. These sections coincide with zone C and D. For interval 2, the MgO content generally increase with depth. The CaO content of interval 1 are (much) lower than those of interval 2. Interval 1 even has five out of the nine measurements that are below the detection limit (< 0.01 %).

Interval 2 displays a general progressive trend of the CaO values with depth. Almost no variation occurs in the Na2O content of both intervals. Their Na2O values lie between 0.05 and 0.18 %, apart from the peak values (0.29 and 0.27 %) of interval 2. A subdivision of interval 1 could be made based on the K2O values. The first section appears from the top until 9.80 m and has K2O values below 1 %.

The second section (9.80-10.60 m) is characterized by K2O values higher than 1 %. Within interval 2, two units (34.23-34.65 m; 36.13-36.23 m) with a K2O content higher than 2 % are observed. Both intervals display no patterns regarding to the TiO2 values. However, in both intervals a section with higher TiO2 values could be identified. These sections have a mean of 2.50 and 1.27 % for interval 1

(9.50-9.80 m) and 2 (34.36-34.50 m), respectively. The P2O5 content increases towards the top, with a maximum of 0.15 %, in the top sample. The second interval contains a unit (34.36-34.65 m) with higher values compared to the others. The unit has an average of 0.30 %.

Table 6: The elemental composition of the silicate fraction (in wt%) of selected samples at different depths.

Silicate fraction Sample Depth interval Zone SiO Al O Fe O MnO MgO CaO Na O K O TiO P O Total number (m) 2 2 3 2 3 2 2 2 2 5 13/152 8.90-9.00 82.21 13.03 2.45 0.29 0.80 0.06 0.10 0.71 0.75 0.15 100.54 C 13/154 9.10-9.30 72.01 17.07 5.99 0.40 1.37 0.05 0.07 0.99 0.93 0.08 98.96

13/155 9.30-9.40 71.86 17.82 < 0.01 0.04 5.95 < 0.01 0.07 1.00 1.35 0.02 98.11

13/157 9.50-9.64 56.08 28.78 6.47 0.01 3.37 < 0.01 0.16 0.53 2.67 0.02 98.09 13/158 9.64-9.80 56.53 27.54 6.79 0.01 5.73 < 0.01 0.18 0.51 2.32 0.03 99.64

Interval 1 Interval D 13/160 9.93-10.10 68.81 20.42 3.53 0.01 3.13 < 0.01 0.11 2.40 1.12 0.01 99.55 13/162 10.20-10.30 70.35 17.23 4.02 0.02 4.12 0.03 0.09 3.21 0.93 0.02 100.01 13/163 10.30-10.40 65.64 20.60 3.33 0.02 5.06 < 0.01 0.10 2.99 1.46 0.02 99.23 37 13/165 10.55-10.60 71.22 17.89 3.57 0.01 4.01 0.02 0.14 1.41 1.26 0.02 99.55

13/356 34.10-34.23 84.32 3.13 0.99 0.02 10.19 0.11 0.06 0.30 0.20 0.02 99.34 13/357 34.23-34.36 61.53 15.39 4.30 0.04 13.40 0.49 0.29 4.04 0.99 0.04 100.51 G 13/358 34.36-34.46 61.66 14.54 3.08 0.03 16.11 0.85 0.27 2.45 1.15 0.26 100.39

13/359 34.46-34.50 59.36 16.19 2.77 0.02 15.18 1.06 0.09 3.13 1.39 0.41 99.62 13/360 34.50-34.65 66.18 13.13 2.59 0.03 12.94 0.41 0.11 3.82 0.93 0.24 100.37 13/363 34.90-35.09 63.07 1.66 1.05 0.05 16.84 16.48 0.07 0.12 0.11 0.09 99.53

Interval 2 Interval 13/365 35.25-35.40 55.37 7.15 2.21 0.04 23.65 8.64 0.05 0.33 0.33 0.05 97.82 H 13/367 35.56-35.70 58.74 2.61 2.25 0.10 18.43 17.00 0.07 0.23 0.17 0.07 99.67 13/369 35.85-35.97 38.38 4.91 3.34 0.18 26.40 24.15 0.08 0.46 0.33 0.09 98.32 13/371 36.13-36.23 37.41 7.73 3.47 0.17 25.07 22.80 0.12 2.25 0.52 0.16 99.70

4.6 Trace elements

The trace element composition of selected soil samples at different depths is presented in Table 7. The Ba content in zone C (on average 1407 ppm) is much higher than in zone D (on average 258 ppm), zone G (on average 183 ppm) and zone H (on average 51 ppm). Thus the Ba content is decreasing through the intervals. In zone D and G, some low Ba concentrations are occurring, which lower the average of the zone. Like the Ba content, the Co concentrations of zone C are much higher compared to the concentrations of the other zones. The Co content of zone D is larger than that of zone G and zone H. So the Co content shows a similar variation as the Ba content. This trend could also be recognized in the Cr and Cu content. The Ni content is also the highest in zone C (on average 151 ppm) and lower towards zone D (on average 46 ppm). The concentrations of zone G (on average 65 ppm) are higher than those of zone D which is in contrast with the previous trend. From zone G to zone H the content lowers again. Through interval 1, the Sr content is slightly decreasing. The Sr concentrations of zone G are similar to those of interval 1, with one exception namely the top concentration. Zone H presents the highest content of Sr, with a mean of 189 ppm. The Zn content is generally decreasing through interval 1. No real trend could be recognized in the Zn content of interval 2. The Be, Bi, Cd and Pb content were under the detection limit and therefore not represented in the table.

Table 7: Concentration of trace elements (in ppm) in the selected samples at different depths.

Sample Depth interval Zone Ba Co Cr Cu Ni Sr Zn number (m) 13/152 8.90-9.00 1465 871 387 1268 191 22 75 C 13/154 9.10-9.30 1350 499 225 868 111 19 51

13/155 9.30-9.40 317 82 131 366 55 13 47

13/157 9.50-9.64 75 28 91 218 47 11 26 13/158 9.64-9.80 65 31 108 187 60 10 28

Interval 1 Interval D 13/160 9.93-10.10 443 88 103 183 37 15 26 13/162 10.20-10.30 373 79 70 156 42 13 22 13/163 10.30-10.40 325 47 79 229 38 11 37 13/165 10.55-10.60 208 57 86 176 47 12 27 13/356 34.10-34.23 69 24 14 22 14 7 49 13/357 34.23-34.36 301 69 60 39 55 15 102 G 13/358 34.36-34.46 186 45 67 28 92 17 94

13/359 34.46-34.50 128 59 73 24 92 21 75 13/360 34.50-34.65 228 42 64 39 70 14 64 13/363 34.90-35.09 44 14 11 47 8 147 18 interval 2 interval 13/365 35.25-35.40 66 29 26 13 19 69 39 H 13/367 35.56-35.70 48 22 11 <10 7 163 23 13/369 35.85-35.97 39 23 18 <10 14 260 27 13/371 36.13-36.23 56 26 27 14 18 307 30

4.7 Quantitative X-ray diffraction (QXRD)

The mineral compositions (in wt%) of selected bulk samples (powder samples) of the first interval and second interval are given in Tables 8 and 9, respectively. The different minerals determined in oriented XRD patterns of the clay or silt fraction of some selected samples in interval 1 and 2, are presented in Appendix 4. The bulk XRD patterns (powder) are presented in Appendix 6. The oriented XRD patterns of the clay or silt fractions of interval 1 and 2 are given in Appendices 7 and 8, respectively.

Generally, kaolinite (peaks around 0.720, 0.357 and 0.234 nm), and quartz (peaks at 0.334, 0.425 and 0.182 nm) are the main components of interval 1. From 9.50 to 9.93 m depth, a unit with more kaolinite (average of 49 %) is recognized. Everywhere else, kaolinite ranges between 21 and 35 %. This abundant occurrence of kaolinite is confirmed by the oriented XRD patterns of the clay or silt fraction. The quartz content shows a decreasing trend at the transition between zones C and D until a value of 11 % (9.80-9.93 m) is reached. Below this depth a sharp increase of quartz occurs. Two units could be recognized based on the mica (peaks at 1.00, 0.500 and 0.333 nm in the oriented XRD patterns of the clay or silt fraction) and/or mica/smectite (M/S) content (peak at approximately 1.052 nm in the oriented XRD patterns of the clay or silt fraction). The first unit occurs from 8.90 until 9.93 m depth, with values that never exceed 10 %. The second interval has values ranging between 14 and 27 % (9.93-10.60 m). When considering talc (peaks at 0.94, 0.470 and 0.313 nm), a sharp transition from zone C (2 %) to zone D (19 %) is observed. Within zone C, talc sharply increases upwards. Throughout zone D the talc content varies around a mean of 14 %, with a peak value of 19 % (9.80-9.93 m) and a minimum of 10 % (10.10-10.20 m). This observation was once more confirmed by the oriented XRD patterns of the clay or silt fraction. Zone C (average of 18 %) contains higher amounts of goethite (0.498 nm) than zone D (average of 5 %). Hematite (0.270 nm) ranges between 3 % and 8 %, apart from three values which exceed or equal 10 % (9.50 m - 9.93 m).

In interval 2, orthoclase (e.g. peaks at 0.378, 0.354, 0.324, 0.299, 0.248 nm) and plagioclase (e.g. peaks at 0.378, 0.254, 0.241, 0.239 nm) only occur in zone G from 34.23 to 34.65 m depth. Both non clay minerals are decreasing with depth. The peak value of quartz (51 %) (peaks at 0.334, 0.425 and 0.182 nm) is found from 34.10 to 34.23 m depth and the minimum numbers 8 % (34.36-34.46 m). These extremes both occur in the uppermost part of zone G. Between 34.50 and 35.70 m, the quartz content varies around a mean of 33 %. The remaining values are relatively close to the minimum and range between 8 and 14 %. A clear delineation can be seen between zone G and H based on the dolomite (peaks at 0.289 and 0.288 nm) content. Zone G has an average dolomite content of 3 %. The dolomite content in zone H ranges between 42 and 61 %, apart from the lower value of 23 % (35.25-35.40 m). The delineation of these two zones is also recognized in the smectite (peak at around 1.50 nm) and talc (peak at 0.94, 0.470 and 0.313 nm) content.

Table 8: Mineral composition (in wt%) of selected bulk samples of the first interval.

Clay minerals Non clay minerals

Sample Depth interval Zone Kaolinite Mica and M/S Talc Goethite Hematite Quartz Others number (m)

13/152* 8.90-9.00 22 5 1 25 5 41 1 C 13/153 9.00-9.10 28 5 7 14 6 39 1 13/154* 9.10-9.30 27 7 2 15 6 41 2 13/155* 9.30-9.40 35 9 19 7 5 24 1 13/156 9.40-9.50 n.d. n.d. n.d. n.d. n.d. n.d. n.d. 13/157* 9.50-9.64 52 8 10 5 10 13 2 13/158* 9.64-9.80 54 9 11 - 12 11 3

40 13/159 9.80-9.93 42 10 19 5 11 12 1

D 13/160* 9.93-10.10 30 18 11 - 8 32 1 13/161 10.10-10.20 29 20 10 3 3 35 0 13/162* 10.20-10.30 21 27 15 - 5 32 - 13/163* 10.30-10.40 31 19 13 6 4 26 1 13/164 10.40-10.55 n.d. n.d. n.d. n.d. n.d. n.d. n.d. 13/165* 10.55-10.60 21 14 15 4 7 38 1 - : absent

* : successive sedimentation for clay segregation and texture analysis

n.d. : not determined

Table 9: Mineral composition (in wt%) of selected bulk samples of the second interval.

Clay minerals Non clay minerals Sample Depth interval Zone Chlorite and C/S Mica and M/S Smectite Talc Dolomite Orthoclase Plagioclase Quartz number (m) 13/355 33.98-34.10 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 13/356* 34.10-34.23 - 7 7 34 - - - 51 13/357* 34.23-34.36 23 3 6 26 3 13 18 8 G 13/358* 34.36-34.46 19 3 9 31 4 16 10 8 13/359* 34.46-34.50 31 3 10 24 3 11 7 11 13/360* 34.50-34.65 31 13 5 3 3 6 6 34 13/361 34.65-34.80 18 - 3 28 3 - - 48 13/362 34.80-34.90 12 - - - 61 - - 27 13/363* 34.90-35.09 9 1 - 5 45 - - 39

13/364 35.09-35.25 7 - - - 62 - - 31 13/365* 35.25-35.40 49 3 - 2 23 - - 24 13/366 35-40-35.56 16 - - - 48 - - 36 H 13/367* 35.56-35.70 27 - - 3 42 - - 28 13/368 35.70-35.85 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 13/369* 35.85-35.97 26 - - 3 57 - - 14 13/370 35.97-36.13 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 13/371* 36.13-36.23 24 - - 2 48 8 4 13 - : absent

* : successive sedimentation for clay segregation and texture analysis

n.d. : not determined

In zone G, smectite varies around a mean of 6 %. In Zone H, smectite is completely absent. This observation was mostly confirmed by the oriented XRD patterns of the silt fraction. However, these patterns indicate the occurrence of smectite traces in zone H as well. The talc content of zone G is higher than 23 %, apart from the exceptional low value of 3 % from 34.50 to 34.65 m depth. In zone H, the content ranges between 0 and 5 %. This agrees with the oriented XRD patterns of the silt fraction. Mica and/or M/S mainly occur in zone G. In zone H, small amounts are only present from 34.90 to 35.09 m depth (13/363) and from 35.25 to 35.40 m depth (13/365). Based on the oriented XRD patterns of the silt fraction, mica (peak at 1.00 nm) and/or M/S (peak at approximately 1.04 nm) are present from 35.85 to 35.97 m depth (13/369) and from 36.13 to 36.23 m (13/371) depth as well. The chlorite (peak at 1.42 nm) and/or chlorite/smectite (C/S) (peak at 1.43 nm) content of zone G varies around a mean of 24 %. Within zone H, the values fluctuate more, ranging from 7 towards 49 %. Once again, the oriented XRD patterns of the silt fraction confirm this observation, with the exception of sample 13/356 from 34.10 to 34.23 m depth. In this sample, the oriented XRD patterns of the silt fraction and the bulk XRD pattern contradict each other. Chlorite and C/S could even be recognized separately based on the oriented XRD patterns of the silt fraction.

Figures 12 and 13 are two examples of the observed (black line) and modelled (red line) XRD patterns of two bulk samples from interval 1 (13/154 and 13/160). Figures 14 and 15 represent the observed (black line) and modelled (red line) XRD patterns of two bulk sample of interval 2 (13/360 and 13/369), respectively. In both modelled XRD patterns of interval 1 (13/154 and 13/160), the minerals: kaolinite, mica and/or M/S, talc, hematite, quartz and rutile were added to explain the observed XRD pattern. But in sample 13/154, goethite and pyrolusite were added as well to obtain a better result. A good similarity between the modelled and the observed XRD pattern of sample 13/369 was attained by the use of four or five minerals: talc, chlorite and/or C/S, dolomite and quartz. In addition to the last mentioned minerals, smectite, mica and/or M/S, orthoclase and plagioclase were needed as well to obtain a good model of the XRD pattern of sample 13/360.

Figure 12: Observed (black line) and modelled (red line) XRD patterns of the bulk sample 13/154 from interval 1.

Figure 13: Observed (black line) and modelled (red line) XRD patterns of the bulk sample 13/160 from interval 1.

Figure 14: Observed (black line) and modelled (red line) XRD patterns of the bulk sample 13/360 from interval 2.

Figure 15: Observed (black line) and modelled (red line) XRD patterns of the bulk sample 13/369 from interval 2.

4.8 High resolution transmission electron microscope (HR-TEM)

Figure 16 represents the HR-TEM images of intervals 1 and 2. TEM is used to obtain lattice-fringe images of layer minerals. The spacing between fringes gives us information about the mineral. Figure 16a represents talc, with a spacing of 0.95 nm. Figure 16b shows kaolin layers, with spacings of 0.7 and 0.8 nm. Figure 16c and d shows again talc, with a spacing of 0.94 or 0.93 nm.

a b

c d

Figure 16: HR-TEM images of (a) talc from 9.30 m to 9.40 m depth (sample 155), (b) kaolin from 9.30 m to 9.40 m depth (sample 155) and (c, d) talc from 34.90 m to 35.09 m depth (sample 363) (modified from Dumon et al., 2015).

4.9 Micro X-ray fluorescence (µXRF)

Figures 17, 18 and 19 represent the X-ray fluorescence intensity maps of the composing elements of selected areas on thin sections of sample 356, 358 and 366, respectively. These samples were choosen with the intention to map weathered chlorite crystals. The Fe content in the selected areas demonstrates distinct spots with high intensities. These spots coincide with the brown weathered zones seen in the photomicrographs of sample 356 and 358. For sample 366 these spots correspond to Fe nodules. In all the samples, several higher intensity spots were recognized in the K content as well. The Mn content of sample 358 confirms the presence of a Mn nodule. The Mn XRF map of sample 366 presents some high intensity zones; they match again with the weathered areas observed in the microphotograph. The Ti distributions of the selected area of sample 366 represent approximately the weathered zones in the photograph. The Ti content of the other samples show spots, which could not be directly linked to distinct zones in the photograph. The Ca content of sample 366 confirms the presence of dolomite, it corresponds to the weathered spots as well. For the other samples, the Ca content shows some regions with a higher content but no clear relation with the photograph is found. The Ni, Cu and As content of sample 358, all display a higher intensities in the upper left corner. In sample 366, high intensity areas were recognized in the Ni, Cu, Zn and Ga XRF maps as well. The XRF maps of other elements not discussed here, show no significant spatial pattern.

Figures 20, 21 and 22 represent the cumulative X-ray fluorescence spectrum from the selected areas of sample 356, 358 and 366, respectively. The peak areas give information about the quantity of the element. The spectra in Figures 20 and 21 are quite similar in shape. They differ from each other mainly in intensities e.g. the Mn and Cr intensity of sample 358 is higher than that of sample 356. The cumulative XRF spectrum of sample 366 shows a clear peak right of Si and one right of S, these are not well recognized in the other samples. The Ca and the subsequent peak display higher intensities compared to the other cumulative X-ray fluorescence spectra. On the contrary, the Ti and V intensities are lower than those of sample 356 and 358.

The Rayleigh scattering (Rayl) is induced by an elastic collision between the emitted X-rays and the electron cloud of an atom within the sample. The Compton scattering (Compt) is similar to the Rayleigh scattering, but an inelastic collision takes place instead of an elastic one. These scatterings are measured by the detector as well. The ratio of the Rayleigh to the Compton scattering gives an idea about the average atomic mass. Atoms with a higher atomic number (Z) produce more Rayleigh scatter than Compton scatter. Atoms with a lower Z exhibit an opposite trend. Combination of the two can give us information about the presence or absence of sample material.

Figure 17: X-ray fluorescence intensity maps of the composing elements of the studied area of sample 356. A photomicrograph of the study area is shown in the uppermost left corner.

Figure 18: X-ray fluorescence intensity maps of the composing elements of the studied area of sample 358. A photomicrograph of the study area is shown in the uppermost left corner.

Figure 19: X-ray fluorescence intensity maps of the composing elements of the studied area of sample 366. A photomicrograph of the study area is shown in the uppermost left corner. The D represents dolomite.

Figure 20: Cumulative X-ray fluorescence spectrum from the studied area of sample 356.

Figure 21: Cumulative X-ray fluorescence spectrum from the studied area of sample 358.

Figure 22: Cumulative X-ray fluorescence spectrum from the studied area of sample 366.

4.10 Micromorphological properties

Generally, interval 1 is characterized by a micromass with a reddish brown to yellow colour, which is typical for (highly) weathered tropical soils. Mottling and clay illuviation are observed as well. The dominant colour through interval 2 is colourless, apart from sample 13/355. The interval contains saprolite (2:1 phyllosilicates, weathered chlorite, quartz) and bedrock (dolomite) components. The relative amount of soil material in the interval is small. Talc (Figure 23) is recognized in both intervals, in varying amounts. It is present in the soil material component, and also as part of the saprolite or parent rock. Alteration of talc (Figure 23a, b and c) is recognized at several depths. Chlorite (Figure 24) is only observed in interval 2, showing an increased weathering towards the top of the interval.

Below a detailed micromorphological description and photographs of several thin sections are given.

Interval 1:

Sample # Depth (m) Micromorphological description 13/152 8.90-9.00 angular blocky microstructure; weak developed pedality; close porphyric; undifferentiated b-fabric; cracks filled up with fine mass of oxides; colour: brownish to reddish groundmass with black parts. 13/154 9.10-9.30 subangular blocky microstructure; moderately developed pedality; porphyric; crystallitic b-fabric; large oxide nodules with small quartz and 2:1 phyllosilicate inclusions (Figure 25a); colour: brown to reddish groundmass with green to yellow parts.

13/155 9.30-9.40 cracked microstructure; weak developed pedality; porphyric; crystallitic b-fabric; oxide nodules; colour: brown reddish groundmass with yellow bands, indicating reductomorphic features (Figure 25b). 13/156 9.40-9.50 cracked microstructure; moderately developed pedality; porphyric; speckled to crystallitic b-fabric; oxide nodules and hematite crystals: colour: black, reddish and yellow groundmass. 13/157 9.50-9.64 cracked microstructure; weak developed pedality to apedal; crystallic b- fabric; oxide nodules, cavities filled with oxides (Figure 25c); colour: brown reddish groundmass with yellow and black parts; altered 2:1 phyllosilicates. 13/159 9.80-9.93 weak developed subangular blocky to massive microstructure; weak developed pedality; undifferentiated to crystallitic b-fabric; oxide nodules, oxide filled up cavities; colour: yellow to red-brown groundmass; altered 2:1 phyllosilicates. 13/160 9.93-10.10 subangular blocky to massive microstructure; moderately developed pedality; crystallitic b-fabric; oxide nodules, fragments composed out of fine mass of quartz, 2:1 phyllosilicates and oxides; colour: colourless fragments, yellow to red brown groundmass; indications of clay illuviation (Figure 25d). 13/162 10.20-10.30 massive to cracked microstructure; weak developed pedality to apedal; crystallitic b-fabric; oxide nodules, fragments composed out of fine mass of quartz, 2:1 phyllosilicates and oxides,; colour: colourless fragments, yellow to red brown groundmass. 13/164 10.40-10.55 massive to cracked microstructure; weak developed pedality to apedal; crystallitic b-fabric; oxide nodules; colour: colourless, yellow to red brown groundmass. 13/165 10.55-10.60 1. cracked to massive microstructure; apedal; speckled to crystallitic b-fabric; oxide nodules; colour: red-brown groundmass; quartz, kaolinite and 2:1 phyllosilicate are encompassed in the groundmass. 2. massive microstructure; apedal; undifferentiated b-fabric; colour: yellow to red-brown groundmass; clay and oxides are segregated in bands with semi-parallel orientation.

Interval 2:

Sample # Depth (m) Micromorphological description 13/355 33.98-34.10 vughy microstructure; large pieces of altered talc, hematite crystals and oxide nodules; colour; colourless, brown to yellow and green. 13/356 34.10-34.23 massive to cracked microstructure; very weakly developed pedality; occurrence of few small single hematite crystals; colour: colourless with random distributed yellow spots. 13/358 34.36-34.46 massive to cracked microstructure; apedal; small oxide spots; colour: colourless to yellow to green, with some red to brown parts. 13/360 34.50-34.65 massive microstructure; apedal; monic c/f related distribution; colour: colourless to grey, with randomly distributed brown spots of oxides. 13/361 34.65-34.80 channel to chamber microstructure; apedal; colour: reddish brown aggregates and colourless aggregates; altered 2:1 phyllosilicates and complex altered chlorite (Figure 24c, d, e and f). 13/362 34.80-34.90 vughy microstructure; apedal; close to single spaced porphyric c/f related distribution; cristallitic b-fabric; hematite crystals (Figure 25e and f) and oxide nodules; colour: colourless to brown with greenish parts. 13/363 34.90-35.09 massive microstructure; apedal; colour: colourless with small brown/ black spots; highly weathered chlorite (Figure 24c, d, e and f). 13/365 35.25-35.40 subangular blocky microstructure; moderately developed pedality; crystallitic b-fabric; good accommodated peds; planar voids; massive dolomitic fragments, highly weathered chlorite (Figure 24c, d, e and f), hematite crystals; colour: colourless with brown parts. 13/366 35-40-35.56 1. cracked microstructure; good accommodation; crystallitic b- fabric; colour: green to brown; clay and oxides are segregated in bands with semi-parallel orientation. 2. cracked microstructure; good accommodation; colour: colourless to green. 13/367 35.56-35.70 subangular to angular blocky microstructure; moderately developed pedality; partially accommodated peds; planar voids and vughs; highly weathered chlorite (Figure 24c, d, e and f), hematite crystals; colour: colourless. 13/369 35.85-35.97 subangular blocky microstructure; moderately developed pedality; partially accommodated peds; planar voids and vughs; massive dolomitic fragments, alteration of dolomite to talc (Figure 23e and f) weathered chlorite (Figure 24c, d, e and f), hematite crystals; colour: colourless. 13/371 36.13-36.23 subangular blocky microstructure; moderately developed pedality; partially accommodated peds; planar voids and vughs; massive dolomitic fragments, alteration of dolomite towards talc (Figure 23e and f), weathered chlorite (Figure 24c, d, e and f), hematite crystals; colour: colourless with some green parts.

a b

c d

e f

Tc Tc D

Figure 23: Microscopic characteristics of talc: (a) alignment of weathered talc (9.93-10.10 m); (b, c) altered talc (9.50-9.93 m); (d) talc in dolomite dominated material (36.13-36.23 m); (e, f) talc associated with dolomite (35.85-36.23 m). The D and Tc represent dolomite and talc, respectively.

a b

371

c d

100 µm 100 µm

e f

100 µm

Figure 24: Micrographs of chlorite at depths between 34 and 36.5 m (interval 2). (a, b) radial shaped chlorite (PPL and XPL) and (c, d, e, f) weathered chlorite (PPL).

a b

c d

e f

H H

Figure 25: Micrographs of (a) (hydr)oxide nodule with quartz and 2:1 phyllosilicate inclusions (9.00-9.10 m); (b) reductomorphic feature (9.30-9.40 m); (c) cavity filled with (hydr)oxides (9.50-9.64 m); (d) illuvial clay (9.93-10.10 m); (e, f) hematite crystals (H) (34.80-34.90 m and 35.70-35.85 m).

5. DISCUSSION

5.1 Evolution of the physico-chemical and mineralogical properties of the two selected intervals

Figure 26 represents the visualisation of interval 1 and Figure 27 represents the visualisation of interval 2, both are based on their physico-chemical properties, DCB (Appendix 5) and mineral contents. These figures envisage the evolution of some distinctive parameters with depth. However, reference is made to the tables of results (section 4) for the exact values.

5.1.1 Interval 1

A general inverse relationship can be recognized between goethite (FeO(OH)) and hematite (Fe2O3). Goethite content increases towards the top. The mineral originates directly from Fe-containing minerals while hematite is then formed by dehydration of goethite (Pérez-Rodríguez et al., 1996).

Their presence is confirmed by comparing the Fe2O3 content of the total elemental composition with that of the silicate fraction. The Fe2O3 content of the silicate fraction is significantly lower, which indicates that iron is mainly present as Fe oxides rather than in silicate minerals. A small amount of ‘other’ non-clay minerals were used to explain some peaks in the bulk XRD patterns. Due to their low concentration it is hard to tell which mineral species these peaks belong to, but are most likely Ti and Mn oxides (e.g. rutile (TiO2) and pyrolusite (MnO2)). Their presence is also suggested by the significant amounts of Ti and Mn in the total elemental composition (Table 4), nevertheless these minerals cannot be determined with certainty.

The dominant clay minerals of interval 1 are kaolinite, mica and talc. The occurrence of these minerals side by side is striking. While the large amount of kaolinite and iron oxides are an indicator for highly weathered conditions, the presence of mica and talc are more typical for less weathered conditions.

Two processes can be responsible for the formation of kaolinite, namely neoformation (Środoń, 1980; De Coninck et al., 1986) and solid-state transformation (Ryan and Huertas, 2009). These formation processes are already discussed in more detail in section 2.2.3. Kaolinite is a typical mineral for strongly weathered soils and is considered as an advanced weathering product. It is a very stable mineral due to the almost absence of isomorphic substitution. In addition, some kaolinites with vermiculitic, micaceous or smectitic layers were found in tropical soils, leading to a net negative charge (Herbillon et al., 1981; Talibudeen and Goulding, 1983a; Righi et al., 1999; Vingiani et al., 2004), making the theoretical absence of isomorphous substitution somewhat relative. Anyway, if a charge is formed on the layer, it still is relatively low compared to the ones originating in micas, vermiculites, smectites and chlorites.

Figure 26: Visualisation of interval 1 based on the physico-chemical properties, DCB and quantitative analyse. T: texture, QPA: quantitative phase analysis, DCB: dithionite-citrate-bicarbonate and CEC: cation exchange capacity (modified from Dumon et al., 2015).

Mica eventually weathers to other minerals with time, typically forming intergrades, i.e. mica → mica-vermiculite → vermiculite or mica-smectite → smectite (Jackson et al., 1952, Wilson, 2004, Velde and Meunier, 2008). Dioctahedral micas occur predominantly in clays of soils and have a basal spacing around 1.00 nm. Trioctahedral micas exhibit a larger basal spacing than the dioctahedral ones. This is related to the bigger octahedral cations. The XRD patterns of the clay and silt fraction (Appendix 7) show peaks with a d-value somewhat larger than 1.00 nm, thus suggesting the presence of trioctahedral micas. However, trioctahedral micas are rare in weathered soils as they are more prone to weathering compared to dioctahedral species. Another possible explanation for these large d-values is the occurrence of mica-smectite. This is confirmed by looking at the second order peaks, revealing slightly irrational series. However, these peaks in the oriented clay and silt XRD patterns (Appendix 7) do not shift clearly when glycolated. So only faint indications for weathering of mica are present. The total elemental composition of sample 13/163 indicates a dioctahedral mica. After subtracting the goethite, hematite and talc components, no more iron and magnesium remains.

Talc is more stable than mica because talc is electrostatically neutral, leading to hydrophobic characteristics. However, observations have shown that the mineral differs slightly from the ideal structure. No mixed-layer minerals of talc were found. The weathering of talc is further discussed in detail in section 5.2. Talc was observed for the first time in zone C (Oostermeyer, 2014). This zone contains lower amounts of talc than zone D. This is confirmed by an increase in Mg content of zone D compared to zone C.

Generally, the persistence of talc and mica besides kaolinite in this highly weathered interval is striking and in contrast with the high amount of kaolinite. A plausible hypothesis to explain this is outlined in section 5.5.

The evolution of the free Fe oxides is similar. They increase upwards starting from 9.30 m depth. The

Fe2O3 increase is explained by the increase in goethite. The quantity of DCB extractable Fe is much larger than the DCB extractable Al and Mn content. This Al and Mn appears as substitutions in Fe

(hydr)oxides. The Fe2O3 content in the silicate fraction shows that only minor amounts of iron occur in the silicate minerals. Thus most iron must be present in the Fe (hydr)oxides.

High amounts of pH dependent variable charge particles, such as kaolinite, talc and sesquioxides are present in the interval. The pH under field conditions is comparable to the pH-H2O, which is around a mean of 4.92. The CEC measurements are conducted with a buffer at pH 7. This may lead to an overestimation of the CEC and thereby an underestimation of the BS. The occupied exchange complex is dominated by Mg2+. Three peak values in the CEC could be recognized from 8.90 to 9.00 m depth, from 9.30 to 9.40 m depth and from 10.20 to 10.30 m depth. These are located at positions with only slightly higher talc and mica contents, respectively. Perhaps these primary minerals are slightly more weathered at those positions, causing the higher CEC values.

5.1.2 Interval 2

The textural analysis of the interval indicates that the clay content is very low for zone G (an average of 3.80 %) and practically absent for zone H. As a result, all the oriented XRD patterns were taken from the silt fraction. The interval can be separated into two units based on the dolomite content. These units coincide approximately with zone G and H, respectively from 34.80 to 34.10 m depth and from 34.80 to 36.23 m depth.

In zone H, dolomite and quartz are the dominant non-clay minerals. The high amount of dolomite is corroborated by high Mg and Ca contents. The mineral is also recognized abundantly in thin sections. In addition, sample 13/366 used for µXRF mapping (Figure 19) contains a fragment of dolomite, which is clearly delineated by the Ca Kα intensity map. Besides the high dolomite content, the following parameters in zone H show high values as well: sand content, LOI and Sr content. The dolomite content from zone H to zone G sharply decreases. This boundary could indicate that zone G is different lithological material.

Figure 27: Visualisation of interval 2 based on the physico-chemical properties, DCB and quantitative analyse. T: texture, QPA: quantitative phase analysis, DCB: dithionite-citrate-bicarbonate and CEC: cation exchange capacity (modified from Dumon et al., 2015).

Talc and mica are mainly present in zone G. This zone exhibits a high Mg content, which is in accordance with higher amounts of talc. The higher amounts of talc are indicated by a high Mg content. Logically, the Mg content of zone H is higher than that of zone G. This is attributed to the presence of high amounts of dolomite (up to 62 %) in zone H against lower amounts of talc (up to 34 %) in zone G. The Ca content confirms this as well, it is higher in zone H than in zone G, again due to the presence of dolomite and talc, respectively.

Smectite was only noticed in zone G based on the bulk XRD pattern. The oriented silt XRD patterns show the occurrence of smectite mainly in zone G, but some traces of the mineral could be recognized in zone H as well.

Chlorite and the transformation product (C/S) were recognized throughout the whole interval. Most of the primary chlorites are trioctahedral. Throughout the interval, a high Mg content is present. It can be concluded that magnesium is present within the structures of talc, chlorite and dolomite. In addition, dioctahedral chlorites contain more Al (32 to 50 wt% Al2O3) and Si (35 to 43 wt% SiO2) than trioctahedral ones (Van Ranst, 2013). The Al content of the silicate fraction does not even reach 20 % in zone G, of which the larger part has to be present in the feldspars, and is lower than 8 % throughout zone H. Therefore, it can be safely concluded that the chlorites are trioctahedral.

Feldspars are present in the upper and lower part of the interval, respectively from 34.23 to 34.65 m depth and from 36.13 to 36.23 m depth. They are inherited from the parent material, probably from arkoses. The sharp increase from 34.23 to 34.65 m depth, could be an indication of a lithological boundary, as proposed earlier. The presence of orthoclase is also confirmed by an increase in the K content. The Ca and Na contents slightly increase at these depths as well, which indicates plagioclase. However, these increases are not as clear. As mentioned, the higher Al contents are also related to these feldspars. They are somewhat resistant to weathering, to a lesser extent than quartz and to a higher extent than trioctahedral chlorite.

The amounts of free Fe and Mn oxides are significantly lower than those of interval 1. Based on the XRD patterns no Fe or Mn oxides were recognized, apart from a small peak at 0.418 nm in the oriented silt XRD pattern of sample 13/363, which could indicate the presence of goethite. However, oxides were observed in thin sections, mainly present as nodules. The Mn Kα intensity map of sample 13/358 (Figure 18) and the Fe Kα intensity map of sample 13/366 (Figure 19) show the presence of a Mn nodule and several Fe nodules, respectively.

The CEC varies between a minimum of 3.29 cmol(+)/kg and a maximum of 58.09 cmol(+)/kg. Literature data suggests chlorite has a CEC between 10 and 40 cmol(+)/kg, the higher values associated with mixed-layered chlorite-smectite. This explains the higher CEC values of zone H. C/S is occurring within this zone and could explain the maximum value of 57.63 cmol(+)/kg from 35.25 to 35.40 m depth. The interval from 34.23 to 34.50 m depth in zone G is characterized by an average CEC value of 55.13 cmol(+)/kg, attributed to the occurrence of chlorite, C/S and smectite. In general the CEC seems to be determined largely by the presence of chlorite and its mixed-layers.

5.1.3 Weathering degree

The weathering indices and molar ratios (Table 5) clearly indicated that interval 1 is more weathered than interval 2. The higher weathering degree of interval 1 is confirmed by morphological (section 4.10) and quantitative phase analysis (sections 5.1.1 and 5.1.2) as well. Morphological analysis show that the first interval is characterized by free iron oxides, which colour the soil red to brown. While

the second interval is mainly colourless. Interval 1 has a silt loam texture. Interval 2 generally has a much more sandy texture, although the upper part (zone G) still has more silt.

Within interval 2, a subdivision could be made, with zone G (occurrence of smectite) being more weathered than zone H. The pH-H2O generally decreases with increasing degree of weathering, due 3+ to liberation of Al .6H2O upon weathering of silicate minerals, leaching of base cations (expressed by a lowering of the BS) and formation of colloidal constituents like kaolin minerals and/or sesquioxides, which have amphoteric groups at their surfaces. Interval 1 has an average pH-H2O of

4.92 (Table 2). In interval 2, a distinction between zone G and H could be made based on the pH-H2O

(Table 3). Zone G and H have an average pH-H2O of 6.79 and 8.29, respectively. Thus, the first interval is more weathered than the second. Only based on pH-H2O a division within he second interval could be made as well, namely that zone G has a higher weathering degree than zone H. However, the mineralogy should also be taken into account. Dolomite, which is a carbonate and leads to an increase in the pH-H2O, is much more pronounced in zone H (average of 48.3 %) than in zone G (average of 2.6 %). Thus a difference in weathering within the second interval cannot be ambiguously established.

Quartz is generally abundant in the two intervals. The mineral is absent in the clay fraction and present in the silt and sand fractions of interval 2. For interval 1, quartz is also recognized in the clay fraction. This evolution is attributed to the low weathering potential of quartz (Goldich, 1938). Thus quartz will be stable in the coarse size classes far longer compared to other minerals.

5.2 Weathering of talc

Interval 1 shows a decreasing trend of talc with depth. Interval 2, contains 20 to 35 % of talc in zone G, the amounts in zone H are much lower or even zero. Talc is recognized in the clay and silt fractions of interval 1 and 2, respectively. The mineral is present in the silt fraction of interval 1 as well (Oostermeyer, 2014). The dominant presence of talc in the silt fraction could indicate that talc is inherited from the parent material.

Some authors state that the occurrence of talc in a soil is rare due to their instability to weathering (Pérez-Rodrigues et al., 1996) and pedogenesis formation is unlikely (Zelazny et al., 2002). But the mineral was recognized several times in different soils (e.g. Lynn and Whittig, 1966; Harris et al., 1984; Pérez-Rodríguez et al., 1996). Harris et al. (1984) concluded that talc was coated by iron oxides, leading to the protection against the soil solution and preventing dissolution. This was not the case in this study, because iron oxides and talc are clearly separate phases, and not intimately mixed with each other, or present as coatings.

Others suggest that talc weathers to smectites, nontronite (Besnus et al., 1976; Zelazny et al., 2002) or more likely saponite (e.g. Veniale and Van der Marel, 1968; Guenot, 1970) through interstratified T/S or to iron oxides (e.g. Paquet et al., 1982; Zelazny et al., 2002). The peaks of the oriented XRD patterns (Appendices 7 and 8), which indicate talc are sharp and narrow and do not shift when glycolated. This indicates the presence of pure talc and not of T/S. TEM images of both intervals

(Figure 16a, c and d) also show the presence of talc instead of T/S. Consequently, it can be concluded that talc in this material does not weather through solid-state transformation.

The mineral and alteration products were recognized microscopically. The only explanation left for the weathering of talc is progressive dissolution. Cations within the mineral will be removed gradually and this will eventually lead to the collapse of the structure. Octahedral cations are more sensitive to leaching than tetrahedral ones, so Mg2+ ions will probably be removed more quickly from the octahedral layer (Yang et al., 2006).

However, talc can in theory be quite resistant to weathering due to its minimal tetrahedral substitution leading to hydrophobic characteristics (Temuujin et al., 2003), unstrained structure and the absence of interlayer atoms (Yang et al., 2006). The occurrence of talc together with other primary minerals such as feldspars and micas in the second interval fits with the general idea that it is a less weathered zone. However, interval 1 is characterized by a higher weathering degree (section 5.1.3) as suggested by the large amounts of kaolinite and iron oxides. In such an environment, talc should already be weathered and/or completely dissolved but is still present (section 5.1.1). Besides, micromorphological analysis indicate the presence of talc in a bimodal size distribution (e.g. Figure 23a, b), mainly in the first interval. The smaller fragments of talc are made off the disintegrated larger ones. Chemical and physical weathering lead to this disintegration. The presence of the larger pieces, is unexpected considering the high weathering degree of the first interval. These observations seem to suggest talc is an allochthonous material in the first interval, but probably in- situ in the second interval. A plausible hypothesis how this came to be is presented in section 5.5.

5.3 Weathering of chlorite

The structure of chlorites is discussed in section 2.2.2. Chlorite and C/S are only recognized in interval 2. The bulk (Appendix 6) and oriented (Appendix 8) silt XRD patterns of the interval confirm the presence of these minerals. The C/S peak in the oriented XRD patterns shift when glycolated. Overall, it suggests a solid-state transformation of chlorite to smectite. The thin sections confirm the presence of weathered chlorite (complex altered).

The presence of trioctahedral chlorite was confirmed earlier (section 5.1.2). According to Velde (1995), homo-ionic substitutions (Mg2+  Fe2+) often occur in these trioctahedral chlorites. The µXRF mapping of sample 13/358 (Figure 18) and 13/366 (Figure 19) shows weathered chlorite and indicates the presence of iron within these structures. The presence of Fe in the chlorite structure is also confirmed by total elemental composition and by total composition of the silicate fraction. The total elemental composition of the chlorite and C/S phases of sample 13/365, is chosen for its high chlorite and C/S content and given in Table 10. The weight percentages were calculated by subtracting the LOI, DCB, dolomite, talc, mica and quartz phases from its total elemental IV VI composition (Table 4). It was assumed that dolomite (CaMg(CO3)2), talc ([Si8] [Mg6] O20(OH)4), mica IV VI ([Si6Al2] [Al4] O20(OH)4K2) and quartz (SiO2) were present in their ideal form. A second assumption is made namely that all the carbon and hydrogen is lost as CO2 and H2O on ignition. The structural

IV VI formula is obtained as being [Si7.79Al0.21] [Mg8.17Al2.10Fe(III)0.55] O20(OH)16, considering Fe in its most oxidated state. Table 10: Total elemental composition of the chlorite and C/S phases (in wt%) of sample 13/365.

SiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O TiO2 P2O5 LOI Total

40.55 10.05 3.81 0.06 28.52 0.49 0.08 0.00 0.55 0.08 15.86 100.00

Several authors (Ross, 1975; Ross and Kodama, 1976; Argast, 1991) concluded that chlorites with higher Fe content are less resistant to weathering. The 3.81 Fe2O3 wt% can be considered as intermediately high. The solid-state formation of chlorite to smectite through mixed-layer minerals starts with the oxidation of Fe2+ (e.g. Ross et al., 1982; Meunier, 2005), followed by the progressive hydration of the hydroxide sheet (Ross et al., 1982). The release of iron leads to the formation of free iron oxides, which were recognized in the thin sections as nodules. Anyway, the transformation process is discussed more detailed in the literature review (section 2.2.2).

Oostermeyer (2014) recognized unweathered chlorite at 39.58 m depth. In this research, it was recognized in thin sections from 35.97 to 36.23 m depth. Weathered chlorite has been recognized (Figure 24c, d, e, f) throughout interval 2. Smectite (intermediate weathering product of chlorite) occurs only in zone G from interval 2, indicating that chlorite in zone G is in a more progressed weathering stage than zone H. The chlorite weathering sequence is eventually completed with the formation of kaolinite out of smectite (sections 2.2.3 and 5.4).

5.4 Formation of kaolinite

Kaolinite was only observed abundantly in interval 1. Its presence was mainly confirmed by XRD patterns (Appendices 6 and 7) and a TEM image (Figure 16b). This observation was also confirmed by Oostermeyer (2014), who recognized the mineral abundantly from zone A until zone D. Zone F contains traces of kaolinite and from zone G to zone I, kaolinite seems to be absent.

Sample 13/163 was chosen as example for the calculation of the Fe content in kaolinite. The content is obtained by subtracting the LOI, goethite, hematite, rutile, mica, talc and quartz phases from the total elemental composition. No significant amount of iron is found in the kaolinite phase of this sample, which agrees with the general composition of kaolinite. If kaolinite would contain significant amounts of iron, it could be an indication that kaolinite is formed out of iron rich smectite (solid- state transformation). However, this was not the case, suggesting kaolinite is formed as a result of neoformation. This was confirmed by the oriented silt and clay XRD patterns of interval 1 (Appendix 7), they do not show indications of K/S. Neoformation or the dissolution-crystallization mechanism is more likely. The remaining elements will leach out or precipitate as (hydro)oxides, e.g. goethite and hematite (Środoń, 1980; De Coninck et al., 1986). These last two minerals were observed in interval 1 as well.

5.5 Relief inversion hypothesis

Like stated before, the occurrence of talc and slightly weathered mica (2:1 phyllosilicates) together with kaolinite and Fe (hydr)oxides in interval 1 is curious. In a normal soil evolution these 2:1 phyllosilicates should already be largely broken down. Thin sections of interval 1 have revealed a heterogeneous parent material, a mishmash of rough and coarse fragments embedded in a groundmass. The 2:1 phyllosilicates (talc and mica) are recognized in different sizes (e.g. talc in Figure 23a, b). The groundmass is composed of kaolinite, 2:1 phyllosilicates and Fe (hydr)oxides. The smaller 2:1 phyllosilicates indicate weathering (namely dissolution) but the larger fragments show no disintegration yet. This indicates the copresence of weathered and less weathered material. Thus this contradiction needs an extra explanation.

Imagine that the original surface was a peneplain (Figure 28a), consisting of parent material at the bottom, followed by saprolite and completely weathered soil material at the top.

This peneplain was incised by a river (Figure 28b), leading to the formation of a valley and the removal of saprolite and weathered material. The remaining bare material at the footslope was less weathered saprolite, consisting of zones G and H. The clear distinction between zone G and H is attributed to different lithological materials. The parent material of zone G contains arkoses, explaining the presence of feldspars. At the bottom of interval 2 (zone H), some feldspars are recognized as well, also indicating arkoses. The zones G and H both contain chlorite and C/S, zone G includes smectite as well (Figure 27). These are signs that the zones are slowly weathering, with zone G more weathered than zone H (section 5.1.3).

It is possible that zone F is also part of this saprolite remnant. Because zone F (Table 11) contains inter alia smectite, hematite, goethite and traces of kaolinite. The presence of these minerals indicates that this zone is even more weathered than the zones G and H. A plausible explanation could be that chlorite is completely transformed to smectite through interstratified minerals. Smectite will eventually dissolve or transform through solid-state to kaolinite. However, the mineral is still abundantly present, which could be attributed to the poorly drained footslope. No information has been gathered on zone E. As a result, the position of zones E and F within the original landscape still remain unclear. This should be solved by investigating both zones in more detail. For this reason, the zones will not be further incorporated in the explanation.

Eventually an escarpment was formed (Figure 28b) and the river started to retract. The resulting escarpment became unstable and collapsed under gravitational force (Figure 28c). This led to an entrainment of weathered and unweathered materials (colluvium), which was deposited at the footslope above zone G. This deposited material represents zones C and D. After deposition the material started to weather again, which is indicated by an increase in Fe oxides towards the top (Figure 26). It was also confirmed by the occurrence of a gravel layer from 7.43 to 9.30 m depth, the larger part of zone C (Oostermeyer, 2014). The gravel layer can be explained as the result of continued erosion, leading to depletion and enrichment of finer and coarser materials, respectively.

Bedrock Laterish material Valley infillings

Saprolite Colluvium River

Figure 28: Illustration of the relief inversion hypothesis: (a) peneplain, (b) river incision, (c) gravitational collapse and (d) fill up of the valley. Red: indications of the zones.

Ultimately, the river pulled further away and a dried up valley was left. Ultimately, the valley was filled up with material, coming from higher ground (Figure 28d). As a result, weathered material was deposited in the valley above zone C. This material weathered further, forming zones A and B. At the bottom of zone A (2.70-3.03 m), a second gravel layer is recognized, the so-called stone layer. This layer might be the result of termite activity. The termites will bring up finer particles, used for building their mound (Oostermeyer, 2014), leading to an accumulation of coarser material below the mound. The top of zone A is weathered the most and a laterite layer was developed. This hardened layer is more resistant than the adjacent material. The latter started to weather more easily and was removed slowly. So the former valley bottom was filled up and evolved to a ridge.

Table 11: Mineralogical composition of the whole core (modified from Oostermeyer, 2014).

Clay minerals Non clay minerals depth interval Mica Zone Chlorite Ortho- Plagio- (m) Kaolinite and Smectite Talc Dolomite Gibbsite Goethite Hematite Quartz Others and C/S clase clase M/S A 0.00-3.03 - +++ + - + - ++ ++ ++ - - +++ + B 3.03-7.43 - +++ + - + - - ++ ++ - - +++ + C 7.43-9.30 - +++ ++ - ++ - - ++ + - - +++ + D 9.30-22.80 - +++ +++ - ++ - - + + - - ++ + E 22.80-28.89 n.i. n.i. n.i. n.i. n.i. n.i. n.i. n.i. n.i. n.i. n.i. n.i. n.i. F 28.89-32.68 - * + +++ +++ - - + + - - +++ + G 32.68-34.80 +++ - ++ ++ +++ + - - - ++ ++ +++ -

70 H 34.80-37.32 +++ - + * ++ +++ - - - + + +++ -

I 37.32-40.00 - - - - +++ + - - - - - ++ - + : few ++ : common +++ : many * : traces - : absent n.i. : not identified

6. CONCLUSION

This research is a continuation of previous work on a 40 m soil-saprolite sequence under a Macrotermes mound near Lubumbashi in the Katanga province of the D.R. Congo (Oostermeyer, 2014). Two intervals (8.90-10.60 m and 33.98-36.23 m) were selected, based on this previous work. Different aspects such as kaolinite formation pathways and weathering behaviour of talc and chlorite are examined. This section presents a brief summary of all work, with the obtained conclusions and findings and finalized with suggestions for possible future research.

The quantitative mineralogy of both intervals was compared. Interval 1 consists mainly of quartz, Fe oxides, kaolinite, talc and mica. Interval 2 is composed out of quartz, feldspars, dolomite, talc, mica, smectite C/S and chlorite. These compositions, supported by weathering indices and molar ratios, indicate that interval 1 is more weathered than interval 2.

Interval 1 contains an inconsistency: talc and poorly weathered mica still occur together with kaolinite and iron oxides, both indicative of an advanced weathering stage. Little information is available on weathering of talc, mainly attributed to its rare presence in soils. In this research, no evidence is found that talc is protected by iron oxides. Yet, the XRD patterns and TEM images of interval 1 suggest the presence of pure talc instead of T/S. Thus these observations suggest that talc weathers by progressive dissolution.

Another anomaly is that 2:1 phyllosilicates are observed as rough and coarse fragments embedded in a groundmass of phyllosilicates and Fe (hydr)oxides. This indicates the copresence of weathered and less weathered material. This admixture could be attributed to a relief inversion. The former land surface was incised by a river, forming a valley. At the footslope of the valley wall saprolite still occurred, which is now included in interval 2. The compositional difference between the upper- and lower part of interval 2 is attributed to a geological difference of the parent material. Eventually a gravitational collapse of the valley wall occurred, leading to deposition of material at the footslope. This deposited material is the former mixture of weathered and less weathered material, which is now included in interval 1. Finally, the valley is filled up with material from higher ground, which evolves into zones A and B after a while.

Chlorite and C/S are only recognized in interval 2. The presence of C/S proves that chlorite weathers according to a solid-state transformation. Chlorite is trioctahedral with mainly magnesium in the octahedral sheet. But some homo-ionic substitutions occur where magnesium is replaced for iron. This iron oxidizes and preludes the weathering process of chlorite. No significant indications were found that kaolinite forms out of smectite (solid-state transformation), suggesting kaolinite is formed mostly by neoformation.

6.1 FUTURE RESEARCH

In order to continue this work, following suggestions can be made. In this research, the focus was concentrated on two intervals (8.90-10.60 m and 33.98-36.23 m). To get a more complete overview of the core and to confirm the relief inversion hypothesis, analytical analyses, bulk and oriented XRD patterns and XRD modelling should be executed on samples of the other zones. Besides, neighboring cores could be taken to invigorate the relief inversion hypothesis even more.

The oriented XRD patterns of both intervals (Appendices 7 and 8) should be modelled in more detail in PyXRD. This could give more insight in the clay mineral composition and a more precise idea about the type and presence of mixed-layer clay minerals. Successive sedimentation for clay segregation and texture analysis can be redone as well, with removal of the carbonates, to obtain a larger clay fraction.

Further study of the TEM images could confirm the presence of intermediate weathering stages of several clay minerals, more specific the presence of T/S, C/S and K/S. Enhancement of µXRF mapping could reveal the weathering sequence of chlorite and give more information on the process of chlorite weathering.

Experimental dissolution of talc in vitro can reveal the weathering sequence of the mineral and could lead to a better understanding of the weathering process. The mineral should be subjected to several solutions with different acidity levels and cation compositions (mainly variation of the Mg content). The main difficulty in this approach is the separation of pure talc from the rest of the soil material.

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8. APPENDIX 1

The tables below give an overview of the selected samples and the executed analyses.

Sample Depth interval XRD XRD Texture Total elemental pH-H O Zone DCB CEC 2 Thin sections HR-TEM µXRF number (m) (R) (O) analyse composition (1:2.5)

13/152* 8.90-9.00 x x x x x x x n.d. n.d. n.d. C 13/153 9.00-9.10 x n.d. n.d. n.d. n.d. x x x n.d. n.d. 13/154* 9.10-9.30 x x x x x x x x n.d. n.d. 13/155* 9.30-9.40 x x x x x x x n.d. x n.d. 13/156 9.40-9.50 x n.d. n.d. n.d. n.d. x x x n.d. n.d.

13/157* 9.50-9.64 x x x x x x x x n.d. n.d. 87 13/158* 9.64-9.80 x x x x x x x x n.d. n.d.

13/159 9.80-9.93 x n.d. n.d. n.d. n.d. x x x n.d. n.d. Interval 1 Interval D 13/160* 9.93-10.10 x x x x x x x x n.d. n.d. 13/161 10.10-10.20 x n.d. n.d. n.d. n.d. x x x n.d. n.d. 13/162* 10.20-10.30 x x x x x x x x n.d. n.d. 13/163* 10.30-10.40 x x x x x x x x n.d. n.d. 13/164 10.40-10.55 x n.d. n.d. n.d. n.d. x x x n.d. n.d. 13/165* 10.55-10.60 x n.d. x x x x x x n.d. n.d. * : successive sedimentation for clay segregation and texture analysis. R : random oriented O : oriented n.d. : not determined

Sample Depth interval XRD XRD Texture Total elemental pH-H2O Thin Zone DCB CEC HR-TEM µXRF number (m) (R) (O) analyse composition (1:2.5) sections

13/355 33.98-34.10 x n.d. n.d. n.d. n.d. x x x n.d. n.d. 13/356* 34.10-34.23 x x x x x x x x n.d. x 13/357* 34.23-34.36 x x x x x x x x n.d. n.d. G 13/358* 34.36-34.46 x x x x x x x x n.d. x 13/359* 34.46-34.50 x x x x x x x x n.d. n.d. 13/360* 34.50-34.65 x x x x x x x x n.d. n.d. 13/361 34.65-34.80 x n.d. n.d. n.d. n.d. x x x n.d. n.d. 13/362 34.80-34.90 x n.d. n.d. n.d. n.d. x x x n.d. n.d.

13/363* 34.90-35.09 x x x x x x x n.d. x n.d. Interval 2 Interval 13/364 35.09-35.25 x n.d. n.d. n.d. n.d. x x x n.d. n.d.

88 13/365* 35.25-35.40 x x x x x x x x n.d. n.d.

13/366 35-40-35.56 x n.d. n.d. n.d. n.d. x x x n.d. x H 13/367* 35.56-35.70 x x x x x x x x n.d. n.d. 13/368 35.70-35.85 x n.d. n.d. n.d. n.d. x x x n.d. n.d. 13/369* 35.85-35.97 x x x x x x x n.d. n.d. n.d. 13/370 35.97-36.13 x n.d. n.d. n.d. n.d. x x x n.d. n.d. 13/371* 36.13-36.23 x x x x x x x x n.d. n.d. * : successive sedimentation for clay segregation and texture analysis. R : random oriented O : oriented n.d. : not determined

9. APPENDIX 2

Visual core description of interval 1 is given below (modified from Oostermeyer, 2014).

Zone Top (m) Bottom (m) Colour(s) of the matrix Mottling 8.90 9.00 dry: 7.5 R 4/8 (red) dry: 10 YR 3/1 (very dark gray) moist: 10 R 3/6 (dark red) moist: 10 YR 3/1 (very dark gray) (dominant) + dry: 7.5 YR 6/8 (reddish yellow) moist: 7.5 YR 5/8 (strong brown) + dry: 10 YR 8/2 (white) moist: 10 YR 7/1 (light gray) C

9.00 9.30 dry: 2.5 YR 5/8 (red) dry: 10 YR 3/1 (very dark gray) moist: 2.5 YR 4/8 (red) moist: 10 YR 3/1 (very dark gray) (dominant) + dry: 10 YR 7/8 (yellow) moist: 10 YR 6/8 (brownish yellow)

9.30 9.50 dry: 2.5 Y 8/0 (white) dry: 10 R 5/6 (red) moist: 2.5 Y 8/2 (white) moist: 10 R 4/8 (red) (feels like talc) (dominant) + dry: 10 YR 7/4 (very pale brown) moist: 10 YR 5/4 (yellowish brown) + dry: 10 YR 7/8 (yellow) moist: 10 YR 6/8 (brownish yellow)

D 9.50 9.93 dry: 10 R 5/6 (red) dry: 2.5 Y 8/0 (white) moist: 10 R 4/8 (red) moist: 2.5 Y 8/2 (white) (feels like talc, dominant) + dry: 10 YR 7/4 (very pale brown) moist: 10 YR 5/4 (yellowish brown) + dry: 10 YR 7/8 (yellow) moist: 10 YR 6/8 (brownish yellow)

Zone Top (m) Bottom (m) Colour(s) of the matrix Mottling 9.93 10.40 dry: 7.5 R 4/6 (red) dry: 10 R 6/2 (pale red) moist: 7.5 R 3/6 (dark red) moist: 10 R 5/3 (weak red) (dominant) + dry: 10 YR 7/4 (very pale brown) moist: 10 YR 5/4 (yellowish brown) + dry: 10 YR 2/1 (black) moist: 10 YR 2/1 (black) + dry: 10 YR 7/8 (yellow) moist: 10 YR 6/8 (brownish yellow) + dry: 2.5 Y 8/0 (white) moist: 2.5 Y 8/2 (white) (feels like talc) D

10.40 10.55 dry: 2.5 Y 8/0 (white) dry: 10 YR 6/8 (brownish yellow) moist: 2.5 Y 8/2 (white) moist: 7.5 YR 5/6 (strong brown) (feels like talc, dominant) + dry: 5 YR 8/3 (pink) moist: 10 R 5/4 (weak red)

10.55 10.60 dry: 7.5 R 4/6 (red) dry: 10 YR 6/8 (brownish yellow) moist: 7.5 R 3/6 (dark red) moist: 7.5 YR 5/6 (strong brown) + + dry: 5 YR 8/3 (pink) dry: 2.5 Y 8/0 (white) moist: 10 R 5/4 (weak red) moist: 2.5 Y 8/2 (white) (feels like talc)

10. APPENDIX 3

Visual core description of interval 2 is given below (modified from Oostermeyer, 2014).

Zone Top (m) Bottom (m) Colour(s) of the matrix Mottling 33.98 34.10 dry: 10 YR 3/4 (dark yellowish dry: 10 YR 6/8 (brownish brown) yellow) moist: 10 YR 3/3 (dark brown) moist: 10 YR 6/8 (brownish yellow) (dominant) + dry: 10 YR 8/0 (white) moist: 10 YR 8/1 (white) (feels like talc) 34.10 34.36 dry: 10 YR 8/3 (very pale brown) dry: 10 YR 6/8 (brownish moist: 10 YR 8/4 (very pale yellow) brown) moist: 10 YR 6/8 (brownish yellow) (dominant) + dry: 10 YR 3/1 (very dark gray) moist: 10 YR 3/1 (very dark gray) 34.36 34.46 dry: 10 YR 8/0 (white) dry: 10 YR 6/8 (brownish moist: 10 YR 8/1 (white) yellow) (feels like talc) moist: 10 YR 6/8 (brownish G yellow) + dry: 10 YR 3/1 (very dark gray) moist: 10 YR 3/1 (very dark gray) 34.46 34.65 dry: 10 YR 6/8 (brownish dry: 10 YR 3/1 (very dark gray) yellow) moist: 10 YR 3/1 (very dark moist: 10 YR 6/8 (brownish gray yellow) + dry: 10 YR 8/3 (very pale brown) moist: 10 YR 8/4 (very pale brown) 34.65 34.80 dry: 10 YR 8/0 (white) dry: 10 YR 3/1 (very dark gray) moist: 10 YR 8/1 (white) moist: 10 YR 3/1 (very dark (feels like talc) gray) + + dry: 10 YR 6/8 (brownish dry: 5 Y 6/3 (pale olive) yellow) moist: 5 Y 5/3 (olive) moist: 10 YR 6/8 (brownish yellow)

Zone Top (m) Bottom (m) Colour(s) of the matrix Mottling 34.80 34.90 dry: 5 Y 6/3 (pale olive) dry: 10 YR 8/3 (very pale moist: 5 Y 5/3 (olive brown) moist: 10 YR 8/4 (very pale brown) (dominant) + dry: 10 YR 6/8 (brownish yellow) moist: 10 YR 6/8 (brownish yellow) 34.90 35.25 dry: 5 Y 6/2 (light olive gray) dry: 10 YR 6/8 (brownish moist: 5 Y 6/2 (light olive gray) yellow) moist: 10 YR 6/8 (brownish H yellow 35.25 36.23 dry: 5 Y 6/3 (pale olive) dry: 2.5 Y 8/6 (yellow) moist: 5 Y 5/3 (olive) moist: 2.5 Y 7/4 (pale yellow) + dry: 10 YR 8/6 (yellow) moist: 10 YR 5/6 (yellowish brown) + dry: 10 R 5/4 (weak red) moist: 10 R 4/8 (red) + dry: 10 YR 8/0 (white) moist: 10 YR 8/1 (white) (feels like talc)

11. APPENDIX 4

Below the identified minerals in the clay or silt fraction determined by oriented XRD patterns of some selected samples in interval 1 and 2.

Depth Sample Zone interval C/S Chlorite Kaolinite Mica* Smectite Talc number (m) 13/152 8.90-9.00 - - +++ + - + C 13/154 9.10-9.30 - * +++ + - +

13/155 9.30-9.40 - - + + - +++ 13/157 9.50-9.64 - - +++ + - + 13/158 9.64-9.80 - - +++ + - -

Interval 1 Interval D 13/160 9.93-10.10 - - +++ ++ - + 13/162 10.20-10.30 - - +++ ++ - ++ 13/163 10.30-10.40 - - +++ +++ - ++ 13/356 34.10-34.23 ++ - - + + +++ 13/357 34.23-34.36 +++ - - + + +++ G 13/358 34.36-34.46 ++ - - + ++ +++

13/359 34.46-34.50 +++ + - + + ++ l 2 l 13/360 34.50-34.65 ++ ++ - ++ + + 13/363 34.90-35.09 +++ + - + * ++ Interva 13/365 35.25-35.40 +++ - - + * + H 13/367 35.56-35.70 +++ - - - * ++ 13/369 35.85-35.97 +++ ++ - + * ++ 13/371 36.13-36.23 +++ ++ - + * ++ - : absent * : traces + : few ++ : common +++ : many Mica* : mica and traces of mica-smectite

12. APPENDIX 5

The Al, Fe, Mn and Si contents extracted with DCB (in wt%) of selected samples at different depths.

Depth DCB Sample Zone interval number (m) Al Al2O3 Fe Fe2O3 Mn MnO Si SiO2 13/152 8.90-9.00 1.21 2.29 19.13 27.36 0.87 1.12 0.37 0.80 C 13/154 9.10-9.30 0.78 1.47 11.86 16.97 0.53 0.69 0.26 0.55

13/155 9.30-9.40 0.48 0.92 7.47 10.68 0.10 0.13 0.19 0.40

13/157 9.50-9.64 0.25 0.48 4.58 6.55 0.01 0.01 0.11 0.23 13/158 9.64-9.80 0.24 0.45 4.81 6.88 0.01 0.01 0.10 0.21

Interval 1 Interval D 13/160 9.93-10.10 0.23 0.43 3.88 5.54 0.15 0.19 0.13 0.28 13/162 10.20-10.30 0.17 0.31 3.04 4.35 0.11 0.14 0.10 0.21 13/163 10.30-10.40 0.26 0.49 4.50 6.44 0.07 0.09 0.17 0.36 13/165 10.55-10.60 0.19 0.35 3.48 4.97 0.06 0.08 0.10 0.22 13/356 34.10-34.23 0.06 0.12 1.02 1.46 0.09 0.12 0.15 0.32 13/357 34.23-34.36 0.10 0.20 0.88 1.26 0.10 0.13 0.35 0.75 G 13/358 34.36-34.46 0.08 0.15 0.87 1.24 0.09 0.12 0.36 0.76

13/359 34.46-34.50 0.05 0.10 0.48 0.69 0.05 0.06 0.28 0.59 13/360 34.50-34.65 0.06 0.12 1.49 2.13 0.14 0.19 0.35 0.75

erval 2 erval 13/363 34.90-35.09 0.01 0.02 0.86 1.22 0.03 0.03 0.08 0.17 Int 13/365 35.25-35.40 0.01 0.02 0.66 0.95 0.01 0.01 0.05 0.10 H 13/367 35.56-35.70 0.01 0.02 0.96 1.37 0.01 0.01 0.07 0.14 13/369 35.85-35.97 0.01 0.02 0.86 1.23 0.01 0.01 0.04 0.09 13/371 36.13-36.23 0.01 0.02 0.17 0.24 0.01 0.01 0.03 0.06

13. APPENDIX 6

XRD powder patterns of the bulk samples of the two intervals are given below.

100

101

102

14. APPENDIX 7

XRD patterns of some selected oriented clay and silt samples of interval 1 after different treatments are given below.

103

104

105

106

107

108

109

110

15. APPENDIX 8

XRD patterns of some selected oriented silt samples of interval 2 after different treatments are given below.

111

112

113

114

115

116

117

118

119

120