Illitization decrypted by B and Li isotope geochemistry

of nanometer-sized illite crystals from bentonite

beds, East Slovak Basin

Norbert Clauer¹, Lynda B. Williams², Damien Lemarchand¹,
Pierre Florian³ and Miroslav Honty⁴

¹ Ecole et Observatoire des Sciences de la Terre, Université de Strasbourg (UdS-CNRS),
F-67084 Strasbourg, France

² School of Earth & Space Exploration, Arizona State University, Tempe, AZ 85287-1404,
USA

³ CNRS, CEMHTI UPR3079, Université d'Orléans, F-45071 Orléans, France

⁴ SCK-CEN, Health, Environment and Safety Institute, B-2400 Mol, Belgium

**Corresponding author:** Dr. Norbert Clauer, Laboratoire d'Hydrologie et de Géochimie de Strasbourg (CNRS/UdS), 1 Rue Blessig, 67084 Strasbourg, France; office phone: 033 390 24 04 33; fax: 033 390 24 04 02; e-mail: nclauer@illite.u-strasbg.fr

**Keywords:** B and Li concentrations and isotope compositions; B solid-state nuclear magnetic resonance; illite-type nanometer-sized crystals; crystal nucleation and growth; bentonite beds; East Slovak Basin

**ABSTRACT** – Boron and lithium contents and isotopic compositions were determined in nucleating and growing nanometer-sized illite crystals (<0.02 to 0.2 µm) separated from mixed-layered illite-smectite particles of bentonite beds from East Slovak Basin. Substituted into tetrahedral and octahedral sites of illite, respectively, they were used to deduce fluid sources and chemical changes during nucleation and crystal growth. The positively correlated contents in most size separates suggest that the substitution of both B and Li is probably
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# INTRODUCTION

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Studying mixed-layered illite-smectite minerals has been very useful for decades in the reconstruction of thermal histories of and associated fluid flow events in sedimentary basins (e.g., Dunoyer de Segonzac, 1970; Hower et al., 1976; Velde, 1985; Jennings and Thompson, 1986; Śucha et al., 1993; Blaise et al., 2016). Significant progress has been made since separation of illite-type crystals by Nadeau et al. (1984) that they called “fundamental particles”, as they nucleate from and grow in contact with saturated aqueous solutions in volcano-clastic sequences of sedimentary basins (e.g., Clauer et al., 1997; Środoń et al., 2009; Clauer et al., 2013; Williams et al., 2013). Formation of these nanometer-sized particles requires nucleation and, either growth of illite-type layers next to dissolving smectite precursor layers or surface-controlled growth. The transition from smectite- to illite-type particles occurs within temperature and pressure ranges that match hydrocarbon (oil and gas) generation and/or maturation of sedimentary organic matter (kerogen), which motivated an abundant literature on the topic (e.g., Williams et al., 2013 and references therein). While illite growth is determined by the aqueous chemistry (e.g., saturation state controlling the cation supply) and while the crystal surface area controls the growth rate (Eberl et al., 2011), varied geochemical and isotopic signatures of illite particles show that their growth can last variably over millions of years to depths of several thousands of meters (Clauer et al., 2014a).

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Common in sedimentary environments, natural fractionation of light stable isotopes, like B or Li, has been investigated during the last two decades or so in varied types of materials and fluids (rivers, brines, seawater), as well as in organic matter from hydrocarbon-bearing rocks, evaporites and remnants of continental weathering processes (e.g., Pistiner and Henderson, 2003; Teng et al., 2008; Cividini et al. 2010; Lemarchand et al., 2012). Isotopic signatures of sedimentary pore fluids being recorded in illite crystals that precipitated from those fluids during burial, the potential exists for reconstructing the compositions of the paleo-fluids that interacted with maturing hydrocarbons (e.g., Clauer et al., 2009; Williams et al., 2015). Complicating factors include the fluid composition, while pH and temperature impact also the B-isotopic composition of authigenic illite. For instance, the impact of the pH is due to the aqueous B-speciation depending on it (Palmer and Swihart, 1996). Tetrahedral coordinated B(OH)₄⁻ is dominant at pH values >~9 (pKa ambient). Therefore, there is only little isotopic fractionation of B between fluids and tetrahedral mineral sites (e.g., illite) at high pH, since there is no coordination change (Hervig et al., 2002). However, the pH of pore fluids is generally buffered at pH <7 in hydrocarbon-bearing clastic sediments undergoing
burial, where trigonal B(OH)₃ dominates and where B fractionation between illite and water is significant and well established (Williams et al., 2001; Williams and Hervig, 2003). This leaves the primary challenge of determining whether temperature or variations in the fluid δ¹¹B is inducing the measurable isotopic differences in authigenic illite. When studying nanometer-sized illite-type crystals that nucleate and grow in an aqueous environment subjected to progressively increasing temperature due to burial, one needs to consider potential changes in the fluid chemical composition. In turn, interpretation of the measured isotopic changes is far from being trivial, because both variables may last millions of years.

Therefore, the contents and fractionation factors of B and Li are expected to delineate changes in crystallization processes of nucleating and growing illite-type crystals precipitating from pore fluids (e.g., Moldovanyi et al., 1993; Williams et al., 2001; Bottomley and Clark, 2003). To understand best these light-element fractionation processes in illite crystals during nucleation and growth, their concentrations and isotopic compositions were evaluated here in especially small nanometer-sized crystals extracted from illite-enriched illite-smectite layers (labeled I-S hereafter) that were recovered from bentonite beds of the East Slovak Basin (Clauer et al., 2014a). Studying preferentially bentonites is because they represent essentially pure clay material that resulted from alteration of the initial volcanic ash layers during their diagenetic change into bentonites. Ideally this type of volcanic ash eliminates any contribution of detrital minerals that may have been supplied from varied geologic environments, as is the case for shales.

We have also included a few preliminary results of Nuclear Magnetic Resonance (NMR) data into this review article devoted primarily to B and Li isotope geochemistry of nucleating and growing illite crystals, in order to provide further clues on how B(OH)₃ and B(OH)₄ compete for their crystal sites in the initial illite structure. Indeed, NMR shows good sensitivity to ¹¹B (abundance = 80.4%) making it detectable under Magic Angle Spinning conditions and, therefore, potentially useful in mineralogy as it allows straight distinction and quantification of four- and three-fold coordinated B environments (Stebbins and Xue, 2014). Hence, it has an interesting ability to discriminate between the two identified processes of illite formation; keeping an inherited parent structure during smectite transformation into illite or minimizing the bond energy by direct chemical precipitation of illite from a supersaturated fluid. A preliminary attempt to locate B in illite-type crystals of varied size by NMR appeared therefore appropriate and timely to complement the B and Li isotope study of identical mineral fractions. This comparison is even more justified as laboratory experiments (Williams et al., 2002; 2007) showed changes in the B-isotope composition during the crystallization
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trend from smectite to illite and, importantly, showed that B absorbed in clay interlayers is not
necessarily in equilibrium with that in the tetrahedral sites, if interlayer water changed during
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137 crystal growth. To justify the NMR experiment is also the fact that the sole NMR study yet
138 available on clay-type materials is, to the best of our knowledge, on B adsorbed on illite
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crystal surfaces, not on B structurally incorporated into mineral structures (Kim and
140 Kirkpatrick, 2006).

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SOME BASICS OF TRACING ILLITIZATION BY B AND Li
ISOTOPE GEOCHEMISTRY

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Illite-rich nanometer-sized fractions (<0.02 µm = <20 nm labeled “nano-sized” hereafter) consist of the smallest clay-type particles that can technically be separated presently. Theoretically, they represent crystals that nucleated from a supersaturated solution at diagenetic temperatures (Eberl et al., 1998), which may have derived their composition from unstable smectite layers in the mixed layered I-S (Nadeau et al., 1984; Inoue et al., 1987; Eberl and Środoń, 1988). From a technical point of view, the smectite-type layers are broken apart by infinite osmotic dispersion that leaves only the illite layers (e.g., Clauer et al., 1997). Assuming that particle separation was successful, varied geochemical and isotopic information can be extracted from these nano-sized crystals of varied thickness representing progressive growth, such as during burial diagenesis.

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At this point, the available K-Ar studies of nano-sized illite extracts that precipitated in bentonites show that the different fractions yield crystallization ages that apparently relate to the diagenetic history, after verifying the absence of any detrital contamination by electron-microscope observation or by the Bertaut-Warren-Averbach technique (e.g., Eberl et al., 1998a; Clauer et al., 2014b). In fact, the thinnest, smallest particles without pyroclastic and/or detrital contribution yield K-Ar ages that are either similar to or older than those of the corresponding slightly thicker, larger particles, suggesting that illitization proceeded by surface-controlled crystal growth. The smallest crystals that failed for any reason to grow after nucleation yield the oldest ages and, therefore, K-Ar ages of nano-sized particles that underwent a continued, long-lasting crystal-growth process, represent an integral of variable amounts of K progressively incorporated into the I-S mineral structure. The amounts of radiogenic ⁴⁰Ar generated by decay of the incorporated ⁴⁰K are consequently variable in the separates, depending on the supply and growth rate, therefore on the size of the particles and the local evolution of each host rock (Clauer et al., 1997). It is also generally agreed that the

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δ¹⁸O values of illite can help to evaluate its crystallization temperature if a constant fluid chemistry interacted during illitization. However, this is theoretically only valid when illitization occurs during continuous increasing temperature trend, and not when temperature decreases, which has been questioned recently (Clauer et al., 2003). In addition to temperature, the water/rock ratio and the fluid chemistry are also entitled to change during long-lasting crystal growth and, therefore, need to be included in interpreting the isotopic compositions of mineral constitutive elements.

## THE GEOLOGICAL SETTING

The Western Carpathians are part of a young orogenic setting that connects the Eastern Alps to the Western Carpathians (Fig. 1A). Due to their geodynamic evolution, the latter are divided into the Externides representing the neo-Alpine accretion prism of the flysch nappes that thrust over the fore-deep, and the Internides that correspond to the paleo-Alpine and meso-Alpine consolidated parts of the orogen mutually separated by the tectonic suture of the Pieniny Klippen belt (Kováč et al., 1997). One of the driving mechanisms of the neo-Alpine geodynamic evolution of the Western Carpathians was a subduction of the basement (fore-deep) of the Externides Flysch Belt under the overriding orogenic plate (Royden, 1993; Tomek and Hall, 1993).

Most central Western Carpathians are covered by thick Tertiary sediments and volcanics that belong to the Pannonian back-arc basin. Part of this cover is also constitutive of the young Neogene East Slovak sedimentary basin that has been extensively studied for its sedimentary and tectonic characteristics, as well as for its mineral and organic compositions. In turn, this East Slovak Basin is part of the Transcarpathian depression, together with the Mukathcevo and Solotvino basins. It contains many bentonite beds that are relicts of ash layers consisting originally of volcanic glass. During progressive burial, the volcanic glass altered first into smectite because of an initial K lack, and then gradually into I-S depending on the further K supply and the duration of the illitization process. The stratigraphic age of the bentonite beds that are basically empty of any detrital material was set from 17.0 to 11.5 Ma (Table 1; Soták et al., 1993). Some of the basin features are of importance for this study, such as: (1) the occurrence of surface (Vihorlat and Slanské Mountains) and subsurface volcanic bodies, as well as of silicic and calc-alkaline pyroclastic infillings in the host sediments, (2) the high thermal gradient ($Q = 82-113 \text{ mW.m}^{-2}$) that induced temperatures of ~210°C at 4-km depth (Král et al., 1985), and (3) the well-established thermal history of the basin that
allowed correlations of the illitization degree with other paleo-temperature indicators (e.g., Clauer et al, 2014a and references therein).

The studied samples were collected in bentonite beds interstratified in the shales of seven drill cores into the East Slovak Basin (Fig. 1B). Evolution of the volcano-clastics was monitored by the collision of the Carpathian orogen with the North-European platform and later by the crustal extension and thermal subsidence in the Pannonian domain. The consequences were a change from a compressive into an extensional tectonic regime, together with varied paleo-geographic changes during the basin evolution. The basin is considered as a complex back-arc feature (Kováč et al., 1995) with an asymmetric Neogene infilling in both the longitudinal and latitudinal directions, together with a steep fault scarp on the NE and SE margins (Rudinec, 1990). This sedimentary infilling reached a maximum thickness of 8-9 km, comprising several shallow regressive sequences (Soták et al., 1993). The studied bentonite beds are parts of the Lower Badenian to Lower Sarmatian (13.5 to 16.5 Ma) volcano-sedimentary sequences (Table 1).

## SAMPLE DESCRIPTION AND ANALYTICAL METHODS

The selected bentonite samples preserved features typical of rhyolitic and andesitic protoliths, such as corroded quartz phenocrysts, biotite, zoned feldspar and volcanic glass. They were collected in bore-holes drilled into the southeastern part of the East Slovak Basin (Fig. 1B) including that of Cicarovce-1 (samples Cic1/11, Cic1/15 and Cic1/18) to the south of the basin, and those of Ptruksa-47 (sample Ptr47) and Ptruksa-13 (sample Ptr13) at respectively 15 and 5 km to the SE of the Cicarovce-1 drilling. The Senné-13 (sample Sen13) and Senné-21 (sample Sen21) holes were drilled about 25 km to the N of the Cicarovce-1 hole. The Trhoviste-1 drilling (sample Trh1) is located about 30 km to the WNW of the previous, and that of Bidovce-1 (sample Bid1) about 50 km to the W of the Trhoviste drilling. The Bid1, Ptr47, Ptr13, Sen13 and Sen21 samples were collected near salt domes that modified the crystallization conditions of the nano-sized particles (e.g., Honty et al., 2004). The collected samples were recovered between approximately 1,555 and 3,015 m depth and their I-S contain a widely variable content of smectite layers from 10 to 90% (Table 1; Honty et al., 2004). A sequence of four <2 µm fractions of shales collected next to some of the studied bentonite samples were also treated for removal and analysis of the organics that could have been a source of B and Li transferred by grain-boundary diffusion into the bentonite beds, and therefore could have contributed to the B and Li contents of the nano-
sized illite crystals. The B and Li contents and isotopic compositions of these organic compounds were deduced from extracted and treated clay size-fractions.

The overall sample preparation was described elsewhere (Clauer et al., 2014b), together with the analytical aspects. The clay-fraction (<2 µm) separation of the bentonite and shale samples were subjected to Jackson's (1975) procedure, including carbonate removal, mild organic matter oxidation and Fe (oxy-) hydroxide removal followed by centrifugation and sedimentation. Finer fractions (<0.02, 0.02-0.05 and 0.05-0.2 µm) were separated by continuous flow high-speed centrifugation (with a Beckman J2-MC centrifuge) after initial separation of the <2 µm fraction by sedimentation and “infinite” dispersion at a ratio of 1g of powder in 40L of deionized water (Środoń et al., 1992). Oriented diffraction patterns of these fractions were recorded on an X-ray diffractometer (Philips PW1830) in air-dried and ethylene-glycolated states for some of the studied bentonites (Ptr47, Ptr13 and Sen21). The qualitative and quantitative compositions of the separated clay fractions were determined by modeling diffraction patterns with the Sybilla software*.

## Principles and basics of B and Li isotope geochemistry

The B and Li isotope compositions are usually expressed, as for many other elements, as the deviation (in ‰) from a standard, which is boric acid (NIST SRM 951) with an atomic abundance ratio ¹¹B/¹⁰B of 4.0437 ± 0.0001 for δ¹¹B (Catanzaro et al., 1970). The NIST lithium carbonate standard L-SVEC is used for δ⁻Li with a certified ⁷Li/⁶Li ratio of 12.17285 ± 0.00023 (Coplen, 2011). In the present case, the internal reference material IMt-1 illite was calibrated against the NIST standard for evaluation of instrumental mass fractionation (Williams et al., 2012). The delta notation for the two elements is then:

$$ \delta^{11}\text{B} (\text{‰}) = (^{11}\text{B}/^{10}\text{B}_{\text{sample}} / ^{11}\text{B}/^{10}\text{B}_{\text{NIST 951}} - 1) \times 10^3, \text{ and} $$

$$ \delta^{7}\text{Li} (\text{‰}) = (^{7}\text{Li}/^{6}\text{Li}_{\text{sample}} / ^{7}\text{Li}/^{6}\text{Li}_{\text{L-SVEC}} - 1) \times 10^3, $$

for which the B and Li isotopic ratios are determined in standards and samples that were both processed along the same chemical preparation procedure.

Seawater is enriched in ¹¹B relative to the oceanic and continental crust (Lemarchand et al., 2000; 2002). This is considered to result from a preferential transfer of the ¹⁰B(OH)₄⁻

<footnote>*Footnote: Sybilla Version 2.2.2, © Chevron Technology Company, a division of Chevron, U.S.A., Inc., San Ramon, CA. McCarty, D. [personal communication].</footnote>
ion from solutions onto clay particles (Spivack et al., 1987), and from exchange reactions between the two dominant aqueous species $ ^{11}B(OH)_3^0 $ and $ ^{10}B(OH)_4^- $ that control the B isotope fractionation in fluids (Schwarcz et al., 1969). This pH dependence has been used in attempts to derive paleo-seawater pH information from foraminifera and other carbonaceous fossils (e.g., Sanyal et al., 1996; Klochko et al., 2006; Hemming and Honisch, 2007). In the late 1960's, Schwarcz et al. (1969) suggested that enriched $ ^{11}B $ values in seawater relative to continental inputs are due to a fractionation during preferential adsorption of $ ^{10}B $ onto clay particles. The authors calculated an enriching $ ^{10}B $ fractionation by about 30-40‰ in clays from an experiment, and noticed that this factor was in good agreement with a value calculated when assuming a steady state of the ocean with respect to both the B concentration and isotope composition. In fact, this steady state is assumed as remaining stable by a combined B input from continental weathered rocks and $ ^{11}B $ desorption from sediments.

Subsequently, Williams et al. (2001a; 2002) performed laboratory experiments to measure the B-isotope fractionation between illite and water as a function of temperature during illitization of smectite. While previous studies showed that I-S preferentially incorporate $ ^{10}B $ into the illite-type structures during diagenesis (Spivack et al., 1987), those studies had not recognized that B was both adsorbed in the clay interlayer and fixed in the tetrahedral sites. Williams et al. (2001) showed changes in the B-isotope composition that track the mineralogical reorganization of smectite recrystallizing into illite. Importantly, Williams et al. (2002) showed also that the B absorbed in clay interlayers is not necessarily in equilibrium with that in the tetrahedral sites, if the interlayer water changed after crystallization. On the other hand, the B-content of the associated pore waters did not decrease with increasing burial depth in the Gulf of Mexico sedimentary basin (U.S.A.) as expected if B were incorporated into illite (Land and Macpherson, 1992; Moldovanyi and Walter, 1992). Instead, an increase in $ ^{10}B $-enriched pore water (>100 mg/L) was reported, suggesting another source of $ ^{10}B $ in deep organic-rich basin environments. In studying also the B isotopic composition of organic matter from hydrocarbon reservoirs of the Gulf Coast, Williams et al. (2001a, 2001b) found that there is a $ ^{10}B $-enrichment of the brines associated with hydrocarbons. The pore water also became relatively $ ^{11}B $-enriched during progressive migration through clay-rich sediments, as $ ^{10}B $ is preferentially adsorbed on smectite or incorporated into illite. Little is known yet about the B content of oil, but significant amounts are reported in kerogen with $ \delta^{11}B $ similar to authigenic, pore-filling clay minerals in sandstone reservoirs (Williams and Hervig, 2004; Williams et al., 2013). Boron isotopes are also
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fractionated during hydrothermally driven mineral crystallization (Oi et al., 1989; Spivack et al., 1990; Leeman et al., 1992).

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In summary, there is no mineral-specific fractionation of B isotopes, like there is for oxygen, but there is an ion-coordination dependence of the fractionation (Hervig et al., 2004). In diagenetic environments where pH is buffered to <7, B is dominantly in trigonal coordination in the fluids (Palmer and Swihart, 1996), but it substitutes for Si in tetrahedral sites of nucleating and growing illite crystals, inducing a significant fractionation as a function of temperature (Williams et al., 2001; 2007). It is probably appropriate to remember that a precise measurement of the equilibrium fractionation of B substituted in the silicate framework of clay minerals requires removal of all B adsorbed and exchangeable in the interlayer sites. As adsorbed B may have an isotopic composition representing later fluids that interacted with the clay particles either in nature or during sample preparation in the laboratory, it is essential to remove this adsorbed or "exchangeable" interlayer B by cation exchange.

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Li is commonly very enriched in oilfield brines and saline formation waters of sedimentary basins (Collins, 1975; Connolly et al., 1990; Fontes and Matray, 1993; Moldovanyi et al., 1993; Stueber et al., 1993; Wilson and Long, 1993). Also, because of its long residence time, ocean water is fairly homogeneous in its Li concentration (about 180 µg/L) and isotopic composition (δ⁷Li = +32‰; Chan and Edmond, 1988; You and Chan, 1996), but the ⁷Li/⁶Li secular variation in seawater remains basically unknown despite a time-dependent variation in seawater suggested recently on the basis of Li-isotope compositions of foraminifera. Misra and Froelich (2012) reported that the seawater δ⁷Li rose by 9‰ during the last 60 millions years due to changes in continental and seafloor weathering. However, the authors did not provide the necessary evidence for lack of diagenetic alteration of the analyzed foraminifera, which is critical as it occurs easily and might produce similar trends. The potential contribution of Li from organic sources seems also to have not been considered. While it is not known which organic compounds contain Li, or what their biological use may be, it has been shown that Li-C alkanes are stable under diagenetic conditions up to 80°C (Kamiensky et al., 1965). High Li contents (up to 350 µg/L) in oilfield brines (Collins, 1975) also suggest an organic source (Williams et al., 2015). This Li-concentration is orders of magnitude greater in abundance than in seawater (Edmond and Chan, 1988). Marine sediments also contain high Li contents with concentrations up 100 µg/g in pelagic clays, while carbonates and biogenic silica have concentrations close to 1 µg/g (Chan et al., 2002).
## The B and Li isotope analytical procedure

The B and Li contents and isotope compositions of three size fractions from nine selected bentonite samples were analyzed on a Cameca 6f secondary ion mass spectrometry (SIMS) following Hervig's procedure (1996). Each analysis averages 100 isotope ratio measurements and the tabulated value averages three or more analyses on different spots of the same sample (Table 2). The spot size was intentionally defocused to ~30 µm to average more oriented clay crystallites and to decrease the impact of ion bombardment through the thin clay sample. Therefore, each SIMS analysis represents a "mini-bulk" analysis of the size-fractionated population of illite crystals, consuming on the order of 10¹¹ atoms. The obtained B-isotope values were corrected for instrumental mass fractionation (IMF) by comparison with SIMS analyses of the reference Silver Hill Illite (Clay Minerals Society Source Clay, IMt-1) material. The <2.0 µm fraction of this IMt-1 sample was bulk-analyzed by thermal ionization mass spectrometry (TIMS) in two facilities, yielding δ¹¹B of -9 ± 0.4 ‰ (Williams et al. 2001a). The IMF correction is important because the B-isotope instrumental fractionation can reach 3 to 4%, while the IMF usually varies by less than 5‰ during a single analytical session. Measuring the IMF of a standard material before and after each sample analysis helps to efficiently constrain the instrumental drift.

Separation of nano-sized fractions that includes infinite osmotic dispersion for the collection of illite-rich particles removes the interlayer B from smectite crystals by Na-exchange. To insure that all surface- or edge-adsorbed B contaminants are entirely removed, the samples were sonified with a 1.8% mannitol solution (a B-complexing agent) and shaken for 24h (Hingston, 1964; TONARINI et al., 1997). The mannitol was rinsed from particles by three washes in distilled-deionized water that had been filtered through Amberlite resin to remove the exchangeable traces of B (Williams et al., 2012). Clay suspensions were mounted as 5 µL aliquots (2-10 mg clay), dried flat onto B-free glass slides and Au-coated for electrical conduction to compensate for sample charging. The B contents and isotope ratios of the nano-sized crystals represent then theoretically only the B substituted in tetrahedral sites of the illite in equilibrium with the fluid from which it precipitated. The analytical errors of these direct clay analyses are <1‰ (1σ) for the isotope ratios and <5% for the contents.

Measurement of Li-contents and isotope compositions by SIMS follows the same protocol as for B; it is described in detail elsewhere (Williams et al., 2012). Lithium can be held in clay interlayer, octahedral sites and possibly pseudo-hexagonal cavities of the I-S. However, cation exchange removes all but the Li substituted in the octahedral, and possibly in the pseudo-hexagonal cavities of nucleating illite crystals. The analyses were calibrated at
372 $\delta^7$Li = 8.8±1‰ (2σ) using SIMS measurements of <0.2 µm Mg-saturated IMt-1 illite
373 completed at the University of Maryland using multiple-collection, inductively coupled
374 plasma mass spectrometry (ICP-MS) (Roberta Rudnick; pers. comm.). Used to correct the
375 instrumental mass fractionation, the Li isotope composition of this internal reference is
376 determined between each sample analysis, the raw data being normalized accordingly.

377 The <2 µm fractions of the four shale samples were prepared the same way as the
378 nano-sized bentonite fractions. The “bulk” fractions were cation exchanged and mannitol
379 washed before analyzing the “total” structural B and Li. Then, the samples were bleached to
380 remove the organics (Anderson, 1963), rinsed again with mannitol before SIMS analysis to
381 determine the nominally “organic-free” B and Li. The SIMS analytical conditions were the
382 same as for the bentonite nano-sized fractions, the B and Li contents and isotope
383 compositions of the organics being calculated by mass balance.

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385 Principles and basics of ¹¹B solid-state Nuclear Magnetic Resonance

386 The ¹⁰B and ¹¹B isotopes are NMR-active quadrupolar nuclei, which means that their
387 NMR lines are strongly influenced by the electric field gradient generated by the surrounding
388 electrons and atoms. In this respect, ¹⁰B is most often out of reach to current NMR equipment
389 due to an extremely large quadrupolar broadening of its NMR line (Marakami et al., 2009).
390 On the other hand, ¹¹B has received much attention over the last 40 years due to its good
391 sensitivity, being often used to decrypt the structure of various types of borate systems,
392 including minerals (e.g., Stebbins and Xue, 2014). Recent technological improvements such
393 as very-high field NMR magnets (>20.0 T) allow now the analysis of samples with B content
394 even as low as 10 µg/g. (Rollion-Bard et al., 2011; Cusack et al., 2015; Mavromatis et al.,
395 2015). While reliable deduction of the total B amounts is still difficult to achieve from NMR
396 analyses, the determination of the relative occurrence of B in the 3 and 4 coordination states is
397 readily achieved.

398 No mineral-specific fractionation of B isotope compositions, like there is the case for
399 oxygen, has been reported but instead an ion-coordination dependence of the fractionation
400 (Hervig et al., 2002). Boron is also dominantly in trigonal coordination in fluids of diagenetic
401 environments with pH buffered at <7 and ¹⁰B substituted in tetrahedral sites of nucleating and
402 growing illite crystals, which induces a large temperature-dependent fractionation (Williams
403 et al., 2001; 2007). A confirmation of B scavenging tetrahedral positions of illite
404 crystallographic structures, instead of Si, is confirmed by a negative correlation between the B
405 and SiO₂ contents of the illite crystals that will be discussed in a further section.
## Some basics about the Nuclear Magnetic Resonance analytical procedure

A Magic Angle Spinning probe of Doty Scientific with an almost undetectable ¹¹B background signal and an optimized sensitivity at the ¹¹B frequency on a 20.0 T magnet (i.e. 272.7 MHz) was used here for determination of the NMR data at the CEMHTI of Orleans, France. The samples were spun at 20 kHz as a single-pulse (“Bloch decay”) sequence. A small pulse angle of less than π/12 was used to retrieve a quantitative signal (Lippmaa et al., 1986) with an acquisition of typically 24 to 48 hours. The extremely small amount of the ¹¹B background signal has been removed by subtracting quantitatively a spectrum recorded with an empty rotor. The ¹¹B spectra are referenced to BF₃O₂, and the [BO₃]/([BO₃]+[BO₄]) ratios were approximated by direct integration on experimental spectra, thanks to the natural separation between the two BO₃ and BO₄ species obtained at the high magnetic field used in this study.

## RESULTS

### Summary of the available data

The nano-sized separates from bentonite beds consist of I-S mixed layers (Honty et al., 2004) modified by variable degrees of illitization, as expressed by the routine expandability, that is to say, the percentage of smectite layers in the I-S (labeled SₓRD in Table 1). Expandability ranges widely from 90 to 10%, the I-S interlayer's order varying from R0 when >40 % of smectite layers (labeled %S) to R1/R3 when <40 %S. Correlated with the present-day collection depth of the different samples, the %S of the <2 µm size fractions plots along two sub-parallel lines (Honty et al., 2004), suggesting an illitization process mainly driven by a burial-induced temperature increase. These two sub-parallel lines suggest also that illitization recorded different thermal gradients, unless fluid compositions differed at the different sampling locations.

### The XRD data

I-S is the dominant mineral phase of all size fractions from all studied samples (Fig. 2). The R1 ordered I-S with an I/S ratio of about 2.5 is the main mineral of sample Ptr47. A reflection at 14.5 Å that shifts towards 15.2 Å upon ethylene-glycolation was also visible in the air-dried <0.02 and 0.02-0.05 µm fractions, its origin being not yet clear. Sample Ptr13 contains mainly R3 ordered I-S with about 88% of illite layers. Quartz was detected together
with kaolinite and I-S in the <0.02 and 0.02-0.05 µm fractions, whereas not in the coarser
0.05-0.2 µm fraction. Kaolinite was present in all three fractions in decreasing amounts from
finer to coarser fraction. R1-ordered I-S with about 84% illite layers dominated in all fractions
of sample Sen21. Similarly to sample Ptr13, quartz and kaolinite were detected in decreasing
amounts from finest to coarsest fraction. Again, an ill-defined peak occurred at ~14.5 Å at the
air-dried state, swelling to 15.2 Å after ethylene-glycolation. The abundance of the smectite
layers in the I-S of the <0.2 µm fractions agrees with data obtained on the <2 µm fractions by
Honty et al. (2004).

448

449 *The K-Ar data*

From K-Ar results published earlier (Clauer et al., 2014a), it can be stated that
illitization was episodic in the two Ptr and Sen samples, lasted variably from as long as about
4.5 Ma in samples Cic1/18 and Ptr47 to as short as less then1 Ma in samples Cic1/11 and
Ptr13. It started variably after deposition, rapidly within 2 million years after deposition in
sample Ptr47, or significantly delayed by more than 7 million years after deposition in sample
Cic1/15. In fact, illitization does obviously not operate like a unique identical process
spreading all over the basin: it was multi-episodic, more or less delayed relative to deposition
and probably the interacting fluids were variably temperatured with changing chemical
compositions: nucleation and growth of the illite crystals were obviously not strictly
contemporaneous for all collected samples. Also, the K-Ar ages either increased or decreased
with decreasing size of the nano-sized particles (Table 2).

Comparing the stratigraphic age of the different bentonite beds with their K-Ar
isotopic ages shows that illitization related somehow to the subsidence episode(s), but not
systematically. The comparison suggests also that temperature was not the only or dominant
driving parameter. Earliest illitization coincides with the Middle-Miocene (Lower Sarmatian)
subsidence from 13.5 to 11.5 Ma in the Trh1, Sen21 and Ptr47 beds, whereas it was delayed
in the Bid1 and Ptr13 samples. Only samples Cic1/15 and Cic1/18 yield younger K-Ar ages
for the onset of illitization, significantly after the Middle-Miocene subsidence. These variable
ages allowed calculation of subsidence rates that organized into three sets (Clauer et al.,
2014a): high rates (>500 m/Ma) for samples Trh1 and Ptr47 before the earliest detected
illitization, intermediate rates (from 500 to 300 m/Ma) for samples Cic1/18, Sen45, Sen13 and
Ptr13, and for samples Cic1/11 and Sen21 before the oldest illitization episode, and low rates
(<300 m/Ma) for samples Cic1/15 and Bid1 before initial illitization, and for samples
Cic1/11, Sen21 and Ptr47 before a further illitization episode. The thermal gradients could be

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grouped the same way into three categories: high (>60°C/km) for none of the samples studied here, intermediate (from 60 to 50°C/km) for samples Cic1/15, Sen21 and Sen13, and low (<50°C/km) for samples Cic1/11, Cic1/18, Trh1 and Bid1, and for samples Ptr13 and Ptr47 before the second illitization episode. In comparing the subsidence rates and the thermal gradients of the different samples, the high rates correlate with the low gradients, such as in samples Trh1 and Ptr47, whereas intermediate values characterize all other studied samples.

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## The B and Li data of illite from bentonites and of organics from host shales

The B contents of the bentonite nano-sized separates scatter widely from 30 to 491 µg/g, with a large δ¹¹B range from -26.9‰ to +6.8‰ (Table 2). Comparing the B contents and δ¹¹B values of the studied size fractions shows narrow plots, within analytical uncertainty, for the different Cic1/18, Bid1, Trh1 and Ptr47 size fractions. The two finest fractions of samples Sen13 and Ptr13 yield the same K-Ar age (Table 2; Clauer et al., 2014a), plotting narrowly while the coarsest and the older illite fractions are ¹¹B enriched, suggesting that they precipitated from a different fluid. The three data points of sample Cic1/15 are widely scattered, as well as those of sample Cic1/11 (Fig. 4).

The organic matter of the <2 µm fractions from host shales yields low B contents (50 µg/g) relative to those of the bentonite clays. Nonetheless, its δ¹¹B ranges widely from +16 to -12‰, with three negative values at -12 and -4‰ and one positive at +16‰ (Fig. 3). Combining the δ¹¹B and B contents of the illite crystals from bentonites and of the organics from hosting shales shows that almost all illite fractions, especially the finer, are within the range of the negative δ¹¹B values of the organics, the isotopic fractionation between B in organics and water having not been quantified. The δ¹¹B and B contents of the coarsest illite fractions trend either towards positive or negative δ¹¹B values when their B contents increase (Fig. 4). A more detailed examination of the combined B contents and δ¹¹B values outlines trends that depend on the crystal size. The coarsest Ptr13, Sen21 and Sen13 fractions trend towards higher B contents and δ¹¹B, like they could have interacted with fluids concentrated in B of organic origin similar in isotope composition to the organics from nearby host shales with a +16‰ δ¹¹B value. Also, the data points of the three Cic1/11 fractions trend towards high B contents, but with low δ¹¹B implying a different origin for the interacting fluids. The δ¹¹B data of all other illite crystals are characterized with narrowly ranging δ¹¹B from -2‰ to -14‰, whereas their B contents vary widely from about 20 to 400 µg/g. In the detail, the B contents might be split into successive amounts from 20 to 60 µg/g, from 100 to 200 µg/g and

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more widely from 250 to 400 µg/g (Fig. 4). At a first glance, it looks like the fluids with which nucleating and growing illite crystals interacted had different B concentrations.

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Basically, there is no reason that δ¹¹B and B contents of minerals correlate. However, the situation is specific here, as the younger coarser crystals could either have grown from finer that nucleated first, or could have plausibly resulted from an older illitization episode that was different than that of the finer. In any case, changing B contents and isotope compositions of nano-sized illite crystals reflect changing fluids from which they equilibrated. Therefore, the obtained B isotopic values can result from either: (1) a changing δ¹¹B of the nano-sized crystals without a significant change of the B content, which suggests varied crystallization temperatures, or (2) changing B contents due to varied fluid composition, with possibly but not necessarily a different B source, or (3) changing δ¹¹B values and B contents that reflect variations in both the temperature and water composition. It can then be deduced from analytical database that: (1) the chemistry of the interacting fluids (and probably the burial temperature) remained about constant during illite nucleation and growth in the samples Cic1/18 and Ptr47, (2) the B dilution of the fluids interacting with samples Bid1, Sen 21 and Trh1/37 was slightly different, (3) the chemistry of the fluids changed continuously while crystals were growing towards higher B contents and lower δ¹¹B in sample Cic1/11, and (4) the B contents and δ¹¹B of the fluids interacting with the illite crystals increased during crystal growth in samples Sen13, Ptr13 and Cic1/15, which suggests combined changing fluid chemistry and temperature. The overall increase in the B contents and often in the δ¹¹B of the coarser fractions apparently indicates crystallization at different temperature in connection with different fluids (Fig. 4).

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The Li contents of the same sample fractions are also widely scattered from 3 to 139 µg/g, however with most contents below 20 µg/g and only three above 100 µg/g (Table 2). The δ⁷Li values range from -21.6 to +7.0‰ with a correlation between Li content and δ⁷Li that can be claimed for the same reasons as those for B, most data points characterizing either an almost unchanged or a significantly increasing Li content with two size fractions of sample Cic1/11 off the two trends (Fig. 5). In fact, sample Bid1 is the only one with all three δ⁷Li values within analytical uncertainty. Two Cic1/18 fractions and two of sample Cic1/11 are also quite within analytical uncertainty, whereas those of Sen13 plot on an almost vertical trend. Those of samples Ptr47 and Cic1/15 yield lighter δ⁷Li for the coarser fractions than for the two others, suggesting an interaction with a different fluid.
The organic matter of the <2 µm fractions from hosting shales also outlines low Li contents of about 10 µg/g, yet an order of magnitude greater than seawater. Nonetheless, the δ⁷Li of the extracted organic matter calculated by mass balance ranges widely from +7 to -17‰. In turn, the clay material contains more Li than the associated organics, but with lower δ⁷Li (except for sample Cic1/18). As clays preferentially take up the light isotope from interacting fluid, it has been inferred that kerogen releases ⁶Li-enriched compounds to the water (Williams et al., 2013). Lithium of the clays will consequently be lighter than that of the water by about 10-15‰ at the burial temperatures assumed here. By combining δ⁷Li values and Li contents of the illite crystals from bentonites with those of the organic matter from hosting shales, the three data points of the organic matter plot along the steepest trend (Fig. 5). Again, the difference in the organic δ⁷Li with changing temperature suggests that ⁶Li is released into the fluid, leaving the high-temperature kerogen holding onto ⁷Li, as is the case in the Alum shale (Williams et al., 2013).

Of interest is also the fact that the B and Li contents are correlated here. With the exception of the three Ptr47 fractions, all other B and Li contents, including the data points of the organics from hosting shales, plot along two positively correlated trends (Fig. 6). The higher yields a convincing correlation coefficient of 0.98 (for n=9), whereas the lower is slightly less defined with a correlation coefficient of 0.74 (for n=8). The two correlations suggest a combined uptake of Li and B by crystallizing and growing illite, which depends on the Li dilution coefficient in the interacting fluids. For Higgins (1988), such correlations suggest that B and Li behave similarly during diagenesis and hydrothermal alteration of vitreous volcanics, as were the bentonite beds right after deposition. Furthermore, B appears to have been stable, whereas Li was mobile after primary devitrification of the ash beds. Still after Higgins (1988), B may have been stabilized by the crystallization of B-bearing submicroscopic phenocrysts in the lavas, or may have been less soluble than Li in relatively cool groundwater. To be remembered in this context are the three outliers here with very high Li contents that are yet to be explained.

## The NMR data

The three nano-sized fractions of the Ptr13, Ptr47 and Sen21 samples were selected for NMR analysis. On the basis of the SIMS data, the fractions of sample Ptr47 yield 137 to 158 µg/g B, those of sample Ptr13 slightly more at 198 to 343 µg/g, and those of sample Sen21 even more from 296 to 310 µg/g. As mentioned above, their K-Ar ages are slightly decreasing
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from 10.2 to 9.0 Ma for two fractions of Ptr47, from 8.0 to 7.2 Ma for those of Ptr13, and
from 5.5 to 4.0 Ma for those of Sen21 (Table 2).

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The ¹¹B NMR spectra outline two characteristic peaks for B that correspond to its trigonal and tetrahedral coordination in the crystal structure (Fig. 7). These peaks confirm the occurrence of B in both coordination sites of the illite-enriched crystals, as previously suggested by Wunder et al. (2005) and Williams et al. (2007). The relative amounts of the trigonal and tetrahedral B are similar in the three size fractions of the two Ptr samples (Fig. 7A and B), while not in those of sample Sen21. The superimposable peaks of the tetrahedral B in the two Ptr samples indicate similar crystallographic environments for B, suggesting in turn a similar nucleation and growth process. Alternatively, the bulky peak shape for BO₃ suggests that it is incorporated in an environment that is not fully compatible with a very-well defined crystallographic site. The lack of significant changes in the trigonal B line shapes diagnoses a highly disorganized crystallographic environment, which is consistent with an incorporation of this trigonal B into the smectite interlayers where less geometrical constraints are required by the crystallographic structure. In Sen21 sample, the highest peak of the trigonal B is, again, in the intermediate size, in fact far higher than in the two Ptr samples (Fig. 7C).

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DISCUSSION

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The B and Li contents and isotope fractionation of nucleating and growing nano-sized illite crystals studied here are characterized by: (1) the absence of specific mineral fractionation for B isotopes, like is the case for oxygen or hydrogen, on the basis of a coordination dependence, and (2) B (and Li if taken up) isotope compositions of the interacting fluids that are fairly indicative of their origin, considering that ionic strength may affect substitution (Dickson, 1990). Therefore, correlating contents and isotopic signatures of B and Li are suggestive of chemistry changes in the interactive fluids and/or in the crystallization temperature during crystal growth. The challenge becomes then to unravel the B and Li contents and the corresponding δ¹¹B and δ⁷Li values of successive crystal generations that altogether appear complex and unconstrained at a first glance. The two independent identified parameters that impact theoretically the behavior of B and Li are both prime parameters for illitization, while the relative duration of the process and the timing of the elemental supplies add a further challenge.
A preliminary finding is that the B and Li contents are positively correlated for most of the studied nano-sized fractions (Fig. 6), together with those of the organics from nearby host shales. This correlation suggests that both B and Li were released by maturating organics into the interstitial fluids. Additionally, isotopic differences may also reflect the influx of saline brines from nearby salt beds. Indeed, it has already been demonstrated that pore fluids of the shales migrated into the bentonite beds, on the basis of K incorporation into smectite crystals modifying into illite. As these shales also contain variable amounts of organic matter and trace elements that migrated with the soluble organics, it is reasonable to assume that the interacting fluids picked up some of the labile elements from organic components and carried them as solutes into the bentonite beds where they were potentially incorporated into newly precipitated illite crystals. The Sol'na Bana Formation of the same region as the bentonite beds studied here contains significant amounts of evaporites consisting mainly of halite (Karoli et al., 1997). Analysis of brine inclusions from this halite indicates that the stability field of the brines that are of the Na-K-Mg-Cl-SO₄ type was at the limit between kainite and sylvite (Galamay and Karoli, 1997). In this saline context, the regional Badenian brines consisted mainly of halite, gypsum and anhydrite, on the basis of the analysis of the fluid inclusions from halite crystals (Galamay et al., 2004). Honty (1999) also reported anhydrite grains in Banovce drill cores located next to the Trhoviste 1/37 drilling studied here, and NaCl-type waters in the Senne boreholes, both confirming that the salt domes consist mainly of halite, which is still mined in the region.

However, the whole saline contribution was obviously not only of marine origin, as the sedimentation included sub-aqueous environments to saltpans with desiccation features, and finally a normal, open-marine regime towards the end of the salt episode. Therefore, contribution of fluids that originated in nearby evaporite beds is plausible with higher K amounts and probably heavier B and Li signatures (Vengosh et al., 1991; Chan et al., 2002). Of course, a mixed contribution of B and Li from both the shale organics and nearby salty beds could have contributed, successively or not, in the present situation. The obtained scattered data raise another difficulty in the geochemical reconstruction of the illitization process, namely that either the temperature was not continuously increasing over time, and/or the chemistry of the fluid incursions was not repetitively identical. For example, the two smallest Ptr47 fractions and the coarser Cic1/15 fraction are way off the common correlation, both with abnormally high Li contents.

## Estimates of the illite crystallization temperature
Present-day temperatures measured in the drill-holes at the depth of each collected sample fit a unique trend including all available data (Fig. 8A). In fact, this trend records only the present-day geothermal gradient of the sedimentary sequence, not that (those) during the successive crystallization episodes detected by the K-Ar ages of illite at the different sampling locations. Another potential tracer of mineral crystallization temperatures is the δ¹⁸O, which allows only precise temperature estimations if the pore fluid δ¹⁸O is known at the temperature of illitization (e.g., Clauer et al., 2014). In fact, crystallization temperatures seem not to have changed significantly here during crystal growth on the basis of the measured δ¹⁸O. As reported by Šucha et al. (1993), precisely in the East Slovak Basin, a further approach to determine the crystallization temperatures of illite crystals can be based on the SₓRD, that is to say, the amount of smectite layers in the illite crystals as a record of illite nucleation. However, the fact that clay fractions of the Cic1/15, Ptr13, Sen13 and Sen21 samples contain far too many illite layers, on the basis of their SₓRD data, relative to the temperatures measured at their respective depth locations, is to be stressed (Fig. 8B). In fact, the present-day temperature at about 165°C for the Trh1 sample seems to correspond to the maximum crystallization temperature. The same can be stated for the Cic1/11 illite with a temperature of about 95°C, and by extension for the Cic1/18 illite that nucleated probably at about 130°C. Therefore, one may postulate that illite of the samples Ptr13, Sen13 and Sen21 crystallized at significantly higher temperatures than those measured presently in the drill holes. On the basis of a graphical extrapolation, these three illite separates could have nucleated at a similar 175-180°C temperature as that estimated for sample Trh1, whereas that of sample Cic1/15 may have occurred at 160-165°C (Fig. 8B).

In fact, as the illite δ¹¹B values are clearly not coincident with the present-day temperatures measured in the drill holes (Table 2), it can be inferred that negative δ¹¹B values were obtained for the Cic1/11 fractions that crystallized at about 95°C, whereas positive δ¹¹B values were obtained for the Sen21 size fractions that could have crystallized at about 165°C on the basis of the SₓRD data (Fig. 8B). However, illite crystals of sample Trh1 that precipitated also at about 165°C and those of samples Sen13 and Ptr13 for which crystallization temperature was estimated to be in the same range yield variably negative δ¹¹B of -10.5, -15.4 and -0.8‰, respectively. Therefore, illite having crystallized at about 165°C outlines variably interacting ¹¹B-enriched waters with δ¹¹B ranging widely from -15.4 to +6.8‰.

In summary, one cannot suggest a direct connection between crystallization temperature and δ¹¹B values of the authigenic illite crystals. In fact, illite crystallization
673 temperature appears to be much less critical than that inducing B and Li releases from
674 organics. If illite crystals precipitated at a lower temperature than that at which B and Li were
675 released from organics, the latter contribution was probably not recorded in the isotopic
676 signature of authigenic illite. It will only relate to changing B isotope compositions of the
677 fluids, if B or Li were released from organic matter into the fluids at approximately the time
678 illite nucleated. Indeed, temperatures at which B and Li are released from organics coincide
679 approximately with the release of NH₄, which peaks at maximum oil generation and R1
680 ordering of I-S (Williams et al., 2013).
681

681

682 **Impact of the fluid chemistry on the B signature of illite**

683 Most illite δ¹¹B values plot within a range between -4 and -14‰ that is independent of
684 the corresponding widely scattered B contents from 40 to almost 500 µg/g. This overall
685 dataset defines a common field framed by three of the four measured δ¹¹B values of the
686 organics from nearby hosting shales (Fig. 4). Two trends suggest the supply of higher B
687 amounts with either a significantly higher δ¹¹B, or a significantly lower δ¹¹B in the finest illite
688 crystals of sample Cic1/11. Obviously, the coarser 0.05-0.1 µm particles of samples Ptr13,
689 Sen21 and Sen13 have interacted with fluids with heavier δ¹¹B signatures than the
690 corresponding finer crystals. The B isotope composition of the fluids that interacted with the
691 younger crystals seems to have been abruptly modified. In the case of sample Cic1/11, the
692 trend towards lower δ¹¹B and higher B contents suggests also a change in the chemical
693 composition of the interacting pore fluids.
694

694 The high B contents and δ¹¹B values could correspond to a “reservoir” consisting of
695 B-rich minerals, for instance of marine evaporitic origin (Swihart et al., 1986; Vengosh et al.,
696 1992; Bottomly and Clark, 2003) that were actually described in the East Slovak Basin
697 (Karoli et al., 1997; Honty et al., 2004). These evaporites could have been dissolved during a
698 burial-induced diagenetic alteration, producing fluids with heavy ¹¹B compositions. An
699 alternative interpretation is based on an aqueous-pyrolysis experiment of an oil-bearing shale
700 from California, which shows that a progressive heating to 365°C of the <2 µm fraction did
701 not produce any detectable alteration of the clay fraction and the other silicates (Clauer et al.,
702 2014b). Therefore, the released B and Li contents were necessarily extracted from
703 progressively altered organic matter of the shale. A simple differential graphic calculation
704 based on the B and Li contents and the δ¹¹B and δ⁷Li of the untreated fraction and of the same
705 fraction after hydrous-pyrolysis heating at 310, 349 and 365°C during 72h and at 365°C
during 216h, shows that the B released at each heating step was nearly constant at 140 µg/g regardless of the applied temperature. However, this released B yields changing δ¹¹B values that increased from +2‰ at 310°C to +9.5‰ at 365°C (Fig. 10A; Clauer, Williams et al., unpublished data). The progressively higher δ¹¹B obtained for the organics can then explain the high δ¹¹B values at +16‰ obtained by calculation for those of the host shale nearby the bentonite beds. Compared to the lower δ¹¹B of the organics, between -4 and -12‰, the high δ¹¹B of the organics from shales suggests that it resulted from a high-temperature maturation, yet difficult to evaluate relative to the pyrolysis temperatures. In summary, the high δ¹¹B that gave the trend of the coarser fractions of samples Sen13, Sen21 and Ptr13 (Fig. 4) could result from supplies by fluids that either interacted during illitization with saline marine units, or carried B and Li solutes that were released by maturing organics of the host shales. This latter case is well supported by the large δ¹¹B scatter of illite studied here that seems not to be correlated with the crystallization temperatures, at least for the fractions from samples of the “salt-bearing” beds (Fig. 8). Indeed, the δ¹¹B values are widely scattered from -17.6 to +6.8‰ in illite from “salty” bentonite levels, while the temperature estimation is not changing notably, from 165 to 170°C (Fig. 9). Similar δ¹¹B values between -30.1 and +10.2‰ were reported for non-marine evaporite minerals (Barth, 1993). In fact, the narrow temperature spread is not expected to impact the δ¹¹B of illite, neither that of the interacting fluids, which B contents outline a similar theoretical pattern with an insignificant shift from +20.4‰ at 165°C to +20.1‰ at 170°C, obtained by mass-balance calculation. In summary, both interpretations favor local origins for the B incorporated into the authigenic illite crystals, rather than from allochthonous fluids migrating over long distances from foreign source areas.

When bentonites beds are interlayered with sediments, the pore fluids of the latters being loaded with organic-derived solutes characterized by varied elemental and isotopic signatures are expected to diffuse along grain boundaries into the formers (Williams et al., 2013). Also, as reported by Środoń (2010), if such transfer of organic derived B into bentonite beds was continuous during illitization, the coarser grown nano-sized illite crystals characterized by K incorporation should also be enriched in B relative to the initial nuclei (Fig. 11A). At low K concentrations (<3% K₂O), the B contents are low (<50 µg/g) in the Bid1 and Cic1/18 size fractions. Beyond 5.5% K₂O, they increase to as much as about 350 µg/g in the two coarser Ptr13 size fractions. Exceptions of this correlation are the three younger Cic1/11 fractions with extremely high B contents for very low (< 1%) K₂O contents and the intermediate Cic1/15 fraction that yields also an unexpected low K₂O content,
considering that its B content implies a different fluid source. A confirmation of B substituting into tetrahedral positions of the illite crystallographic structure is given by the negative correlation between B and SiO₂ contents of illite (Fig. 11B). Again, integration of B into the illite crystal structure is exponential with low B contents at about 50 µg/g in illite crystals with the SiO₂ content higher than 63% and when the SiO₂ content decreases, the B content increases to 350 µg/g. The few fractions outside the trend are the intermediate Ptr47 and the two coarse Cic1/15 fractions together, again, with the three Cic1/11 separates. Based on several aspects already mentioned, sample Cic1/11 yields specific characteristics that were not detected in the other illite crystals: for instance increasing B contents with progressively lighter δ¹¹B values when crystal size increases. As it represents the youngest (<1 Ma) illite generation studied here, it might have precipitated during the most recent regional volcanic activity. Illitization obviously proceeded much faster than for the other bentonite beds, within a different thermal and chemical context characterized by lower K contents and relatively higher B contents (246 to 491 µg/g) of very light δ¹¹B from -13.5 to -26.9‰. No B isotopic data are available for the glass materials of this volcanic activity, to the best of our knowledge, but continental arc volcanic rocks yield systematically light δ¹¹B that can reach -15.0‰, for instance in Western Anatolia (Tonarini et al., 2005). The cases of the intermediate sized fractions of samples Ptr47, Cic1/15(3) and Cic1/15(2) that plot outside the trend are different. Their displacements on the graphics could be due to a dilution effect by increased amounts of quartz that were not detected in some fractions (Table 2 of Honty et al., 2008). Indeed, small amounts as determined here (< 4% for the highest quartz occurrence in the intermediate Cic1/15 fractions) are difficult to quantify routinely in clay-rich separates.

While most coarse illite crystals yield high B contents, it is not the case for those of the samples Cic1/18, Bid1 and Ptr47 (Fig. 4). When growing, the nano-sized illite fractions (numbered 1 to 3 from finest to coarsest) may or may not trend toward a specific fluid reservoir, and thus record fluid mixtures. In the case of crystal growth, changes in the δ¹¹B values may also reflect changes in temperature. A regional δ¹¹B decrease of ~3‰ can be explained by the isotopic fractionation (1000 lnα) decrease between mineral and water, due to progressive burial temperature increase (a 50°C temperature increase can induce a δ¹¹B decrease of 3‰). Temperature increase during organic maturation could then have induced an increased release of ¹⁰B from kerogen into the associated pore fluids, still locally with a limited migration, as expected during progressive burial diagenesis. Calculation of the δ¹¹B of fluids relative to theoretical temperatures has been tested in the above section on the basis of
the few known parameters, but its impact has not been very significant. The temperature increase in the East Slovak Basin has probably been too limited to modify significantly the $δ^{11}B$ of either the solids or the fluids. The same seems to have happened for the fluid $δ^{7}Li$ that shifted systematically from +18.3‰ at 98°C to +14.6‰ at 165°C, on the basis of the illite $δ^{7}Li$.

## Respective impacts by changing fluid temperature and chemical composition

According to Michalícek (1965, 1970), pore waters infiltrated the salt deposits of the East Slovak Basin during the Middle-Miocene period (Burdigalian = Badenian) and, according to Magyar and Ocovský (1981), these saline to hyper-saline pore waters interacted for a long time with nearby bentonite beds identified as “salt-bearing”. In the same context, brackish waters were found in Middle-Miocene beds (Langhian = Upper Badenian and Serravallian = Lower Sarmatian), but they seem not to have affected the bentonite layers (Michalícek, 1970). Apparently, the salinity of the fluids increased the degree of illitization, which can be estimated by the Na and Cl contents of the nucleating and growing illite. Enhanced illitization in the “salt-bearing” samples resulted then from a local increase in the crystallization temperature due to an increased conductivity in the proximity of the salt domes. In this context, relatively high fluid $δ^{18}O$ estimates were obtained for the size fractions of samples Sen13 and Ptr13 associated with salt occurrences, whereas the Cic1/11 and Cic1/15 salt-free illite fractions yield lower fluid $δ^{18}O$ estimates (Clauer et al., 2014a). In fact, the temperature increase was probably not that intense as the illitization degree of nano-sized illite from salt-free bentonites overlaps that of the salt-bearing illite (Honty et al., 2004). The different illitization degree may, then, outline a distinction resulting from water evaporation until brine formation with an increased $δ^{18}O$. This could have happened prior to overprint by fluid-crystal interactions. Interestingly the illitization degree had no impact on the K-Ar age of both types of illite (Claur et al., 2014a), which is a strong argument in favor of the K-Ar ages.

The samples Sen13, Sen21, Ptr13 and Ptr47 belong to salt-enriched environments, but their B contents and $δ^{11}B$ values are apparently not abnormal, except for the intermediate and coarser fractions that yield positive $δ^{11}B$ values and B contents higher than 300 µg/g, which possibly reflects a local saline supply (Vengosh et al., 1992; Fig. 4). The fractions of these samples yield also higher $δ^{18}O$ values, which could result in more illite layers in the I-S structures, as reported by Honty et al. (2004). Illitization started soon after deposition (about 10 Ma) in sample Ptr47, but did only last about 1 million years. The B contents and $δ^{11}B$ of its
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a chemically constant fluid.

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Experimental determination of an equilibrated isotopic fractionation between illite and water as a function of temperature allows interpretation of the water δ¹¹B if the temperature and mineral δ¹¹B are known (Williams et al., 2001), preferred incorporation of ¹⁰B from fluids into clays having been reported (Spivack et al., 1987; Palmer and Swihart, 1996). Also, where fluids are already ¹⁰B-enriched, as is the case of oilfield brines (Moldovanyi et al., 1993), the δ¹¹B of illite records anomalously low δ¹¹B where temperature is known (Williams et al., 2012). In the East Slovak Basin, the input of saline fluids associated with evaporites could also have changed the fluid δ¹¹B. Outside of this basic picture are the B trends from <1-Ma old nano-sized fractions of sample Cic1/11, which tend towards high B contents and low δ¹¹B. As the B trend of the Cic1/11 fractions is obviously not representative of the “diagenetic trend”, crystallization of this illite cannot result from a long-term interaction with fluids but could be connected with the recent local volcanic activity. The Bid1 size fractions yield also a decreasing δ¹¹B value when grain size increases, while the K-Ar ages suggest episodic illitization about 10 Ma ago for the coarse fraction and 6-7 Ma ago for the two finer fractions, implying two successive episodes with a probable change in the fluid chemistry (Fig. 12). The same seems to have occurred for sample Cic1/18, for which the temperature increase could have been partly obscured by the concomitant changing fluid chemistry. This could be well illustrated by an interaction of samples with saline brines carrying a high δ¹¹B. In fact, given the pyrolysis data that outline consistent B-content increase in fluids over an experimental temperature range of 310-365°C, the diagenetic temperatures seem to have less control over the B-contents than the evaporite source.

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The δ¹⁸O values determined previously (Clauer et al., 2014a) on some of the illite fractions studied here range very narrowly at an uncertainty of ±0.2 to ±0.5‰ (2σ), which suggests relatively constant temperature and fluid-chemistry conditions during crystal nucleation and growth. However, O-isotope variations are not as sensitive to fluid compositional changes as are traces of B, because O is a component of all silicates and fluids, whereas B is not. In summary, while crystallization temperatures might have been different in each bentonite bed, they appear to have remained nearly constant during each illitization episode. Therefore, the variations in δ¹¹B obtained here appear to reflect dominantly variations in the water chemistry.

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Potential impacts of varied temperature and fluid chemistry on illite δ¹¹B
Boron contents and δ¹¹B values of the Ptr47, Bid1, Cic1/18 and Trh1 fractions plot narrowly within analytical uncertainty (Fig. 4). A straightforward interpretation is that neither temperature nor fluid chemistry changed significantly during illitization in those samples. In fact, beyond the limiting analytical uncertainty, it can be seen that if the Ptr47 fractions collected near a salt dome, are within narrow B contents and isotopic values, the Bid1 fractions yield similar B contents at about 50 µg/g with slightly lower δ¹¹B values when illitization progresses by crystal growth, also suggesting constant crystallization temperature and fluid chemistry. Sample Cic1/18 provides similar results with similar B contents of about 30-50 µg/g and δ¹¹B values at about -5‰ with increasing grain size. The two slightly coarser Trh1 fractions yield the same pattern with B contents at about 375 µg/g and an increasing δ¹¹B when grain size increases.

These results match the expected “diagenetic trend” of Williams et al. (2001a, b; 2002), considering that nucleation and growth induce basically preferential incorporation of ¹⁰B into illite-type structures (Spivack et al., 1987). In fact, it looks like interaction of nucleating illite crystals and pore fluids depends on specific local B supplies, for instance from evaporate beds or organic material, both of these being suspected here. For instance, organic matter modifies indirectly the B isotopic composition of pore waters that affects in turn the δ¹¹B of the resulting authigenic illite crystals (Williams et al., 2001a, b). Later, Williams and Hervig (2006) showed also that the first nucleated illite crystals retain the δ¹¹B of the water from which they equilibrated, while growing crystals gradually approach equilibrium with the later fluids. At constant temperature, these would only be of the same δ¹¹B if the fluid composition did not change. Reported in the sediments around the bentonite beds of the East Slovak Basin, but not in the pyroclastic beds, gas occurrences can also cause isotopic anomalies in δ⁷Li of authigenic clays (Williams et al., 2015). It is therefore obvious that timing of light elemental releases from salty minerals and/or organic matter as solutes into fluids, and of their interactions with nucleating and growing illite-type crystals becomes a prime factor to address. In the present case, such fluids could have carried B from evaporites and/or organic sources at varied temperatures and at successive episodes of the regional evolution. The K-Ar ages support the fact that B trends of the nano-sized illite crystals from bentonites studied here did not occur during a continuous process, some of the Ptr47 and Cic1/18 crystals being, for instance, significantly delayed relative to deposition of the ash layers.

If burial-induced temperature is the effective driving parameter of illitization in bentonite beds, back-stripping of some of these volcano-clastic layers indicates a general
rapid temperature increase soon after deposition, from about 15 until 13 Ma for most samples, that is to say before illitization began. A slow temperature increase has been reported afterwards until 6 Ma, and almost no temperature increase is believed to have occurred since (e.g., Honty et al., 2004; Clauer et al., 2014a). In turn, illitization seems not to have occurred during significant temperature increases on the basis of the above information and the available K-Ar ages, while the medium-sized illite of sample Cic1/18 crystallized at an almost constant temperature. In summary, only a slight temperature increase seems to have existed during nucleation and growth of most illite crystals studied here, the question becoming then about the impact temperature did really have relative to changing chemical compositions of the interacting fluids.

884

## K-Ar ages of illite combined with B tracing

Correlating the δ¹¹B values and B contents of the analyzed illite crystals with their K-Ar ages outlines four illitization episodes more or less constrained between 11 and 9, 8 and 5, 4.5 and 2, and 1 and 0 Ma (Fig. 12A). During the initial 11-9 Ma period, the Trh1/37, Ptr47 and Bid1 illite crystals integrated B with a narrow δ¹¹B of -10.4 ± 1.4‰. Interestingly, the amounts of B picked up at that time by the crystals were widely scattered from about 376 µg/g for the Trh1 fractions, to 152 µg/g for the Ptr47 crystals, and to 50 µg/g for those of Bid1 sample. This combination can result from a B-content reflecting the salinity by ionic strength of the fluid due to an increase of the B solubility (from evaporites) and availability to be incorporated into authigenic phases (Dickson, 1990), while the δ¹¹B is controlled by the coordination change from fluid to mineral.

During the second illitization episode from 8 to 5 Ma, the δ¹¹B appears systematically higher than during the previous episode, reaching up to +4.8‰ in the Ptr47, Bid1, Cic1/18, Ptr13, Sen21 and Sen13 fractions. The B contents are again within the brackets of the older crystals, however somewhat lower: the highest B contents range from 349 to 296 µg/g for the Ptr13, Sen21 and Sen13 illite, whereas intermediate contents, from 198 to 137 µg/g, were determined for the Ptr13 and Ptr47 fractions, and lower amounts, from 49 to 30 µg/g, for the Bid1 and Cic1/18 fractions. Authigenic illite of the third crystallization episode between 4.5 and 2 Ma contains B with δ¹¹B values in a similar range as those of the previous generation, from -9.9 to +4.1‰ with two significantly lower values of -17.6 and -15.4‰. The contents range again from as high as 310 to 261 µg/g, to intermediate values of 157 to 105 µg/g, the lowest content being at 30 µg/g. Only sample Cic1/11 of the youngest <1 Ma episode yields

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abnormally low δ¹¹B values from -13.5 to -26.9‰ and abnormally high B contents between 246 and 491 µg/g (Fig. 12A and B).

On the basis of potential interactions with either saline fluids from nearby evaporite layers, or with fluids carrying B released from maturing organics of the host shales, the δ¹¹B incorporated by illite was constant at -10.4 ± 1.4‰ during initial illitization between 11 and 9 Ma ago. This low and constant value excludes normal groundwater, rather suggesting that there was no significant interaction with fluids carrying saline or organic components. During the second and third illitization episodes, the δ¹¹B values were higher, ranging up to +4.8‰, which suggests an impact of either saline fluids when the B contents are high, or of plain ground waters when the B contents remain low (Vengosh, 1994). Abnormally low δ¹¹B values occurring during the third illitization episode are confirmed by the δ¹¹B of the youngest Cic1/11 sample (-26.9‰). Such very low δ¹¹B values were reported for hydrothermal fluids associated with volcanic activity in the Yellowstone National Park (USA; Palmer and Sturchio, 1990), and were attributed to leaching of nearby sediments or hydrothermally altered minerals. Given the evidence for recent regional volcanic activity in the East Slovak Basin (Szabo et al., 1992), this is a viable explanation for these δ¹¹B values.

Combining the K-Ar ages with the B contents and isotope compositions of the analyzed illite crystals outlines a series of inferences, which raise comments about the potential chemistry of the migrating fluids and the illite crystallization temperatures during the evolution of the East Slovak bentonite beds after the initial 12-Ma old step. In the case of the youngest Cic1/11 illite generation probably less than 1-Ma old, the crystallization temperature appears to have decreased, probably due to decreasing volcanic activity, and the fluid chemistry probably changed. For the fractions of samples Trh1, Ptr47 and Cic1/18, neither the temperature nor the fluid chemistry seems to have changed between 11 and 2.5 Ma ago. For the Bid1 and Sen21 illite crystals, the situation is different: the fluids remained similar, whereas the temperature decreased between 11 Ma and 6 Ma ago for the former, and between 5 to 4 Ma for the later. The data of the samples Sen13, Ptr13 and Cic1/15 appear slightly more complex: in the case of sample Sen13, the fluid chemistry seems to have changed between nucleation/growth of the coarser and that of the two finer fractions. For sample Cic1/15, temperature seems to have decreased during crystallization of the intermediate size, the fluid appearing modified for the coarser fraction between 5 and 3 Ma ago. If B contents and δ¹¹B of authigenic illite result from a combined changing of the fluid chemistry and temperature, one might attempt to compare the B contents and isotope compositions of the size fractions with their K-Ar ages, and also to question some of the data,
as they could have been slightly biased by either the extraction technique or the analytical
methods. For instance, the intermediate fraction of sample Cic1/18 seems abnormally young
(Fig. 12A and B) relative to the two other fractions, as is the case for the same intermediate
fraction of sample Ptr47. On the other hand, the B content of the intermediate fraction of
sample Cic1/15 appears slightly high. This three decreasing trends can be visualized for the B
contents of illite relative to their crystallization timing. During initial illitization, saline fluids
seem to have interacted with sample Ptr47 and the coarser Bid1 fractions with variable B
contents, the highest B contents being for the "non-salted" Trh1 illite. At the second
illitization event, illite of the "salted" Bid1 and "non-salted" Cic1/15 samples resulted from
interaction of fluids with low B contents, whereas the three fractions of the "salted" Ptr13,
two of the "salted" Sen13 and Sen21 and one of the Ptr47 samples yield higher contents. If
considering the intermediate fractions of samples Cic15 and Cic18 as suspicious, two
fractions of the "salted" Sen13 and the "non-salted" Cic1/15 samples outline intermediate B
contents of about 150 µg/g (Fig. 12A). If now the δ¹¹B values of illite are evaluated relative to
the crystallization ages, and considering the same restrictions as above for the B contents, the
oldest illitization episode is characterized by a mean δ¹¹B of -10‰. During the second
episode, δ¹¹B values at -10‰ are still occurring, but most are between -5 and +5‰,
suggesting varied crystallization temperatures among the illite crystals of the different
bentonite beds, unless the chemistry of the fluids changed at the same time. This data scatter
is maintained during the third illitization event, with a couple δ¹¹B data at about -16‰.

While the latest, most recent illitization can be assigned to a different event/source
than the older ones, it can be stipulated from data of the older illite that temperature and fluid
chemistry varied concomitantly for most of the sediments in the basin, except during the
event between 11.5 and 9 Ma. There were no specific areas of the basin that outlined specific
thermal and/or chemical characteristics; it looks rather that the diagenetic process influenced
the illite chemistry of each bentonite bed at each location. About the meaning of the generic
term used for "fluid chemistry", it is needed to recall that the interacting fluids were
apparently supplied here by two identified pools: (1) B and Li released from maturing organic
matter that the fluids picked up and carried into the bentonite beds together with other
elements, especially K to drive episodic illitization, and (2) brines from saline beds that
contributed also to the changing chemistry of the fluids. Combining B contents, δ¹¹B and K-
Ar ages of nano-sized illite shows that illitization did not act as a generalized process over the
whole basin, but rather locally in response to specific thermal and chemical conditions (e.g.,
supply of reactants). Of course, the end product of the reactions that was extracted from
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sediments yields overall similar-to-identical characteristics of the basic illite mineral, but the
intimate process leading to this “average” illite obviously resulted from variable thermal and
chemical interactions that are specific for different places in the basin. In summary,
combination of all contributing parameters examined here shows that B is an extremely
refined tracer to identify interactions of nucleating and growing illite crystals with their fluids,
if temperature and chemistry changes of the fluids can be sorted out fairly.

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δ<sup>7</sup>Li variations in illite with changing fluid temperature and chemistry

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Fractionation of Li between illite and pore fluid, as for all minerals, is smaller at high
than at low temperature (Chan et al., 1998; Hervig et al., 2002). For instance, the source
mineral/kerogen δ<sup>7</sup>Li becomes heavier when temperature increases, <sup>6</sup>Li being preferentially
released into the fluid. The low Li content in the samples studied here could then either reflect
a different source of Li or a loss of Li from diagenetic illite. If the latter is true the host shales
may have contained organic matter that was altered progressively by the pore fluids under
increasing temperature, releasing its <sup>6</sup>Li into the fluids as solutes and leaving the shale
material <sup>7</sup>Li-enriched. Alternatively, heavier illite δ<sup>7</sup>Li might have precipitated in the
bentonite beds due to the input of fluids of evaporite composition that originated in nearby
salt beds. However, abnormal values such as those of the coarser Ptr47 and Cic1/15 fractions
incline one to consider that a change of fluid chemistry rather than temperature could have
induced the different δ<sup>7</sup>Li values during episodic crystallization, as the estimated 50°C
temperature increase lowers the δ<sup>7</sup>Li by only a few ‰. The very limited changes in the
crystallization temperatures, as already discussed, make it challenging to favor a temperature
impact, especially during the 10-Ma long evolution of the whole basin. In fact, the diagram
δ<sup>7</sup>Li vs. K-Ar ages supports the idea of a change relative to time between ambient
temperature, crystallization age and δ<sup>7</sup>Li composition, with a positive correlation for the
Cic1/11, Sen13 and Cic1/18 fractions. The δ<sup>7</sup>Li values of -15 to -20‰ in samples that are at a
present-day temperature of 130°C (Fig. 5) indicate that precipitation occurred in equilibrium
with a fluid δ<sup>7</sup>Li composition of +1.5 to -3.6‰. These values are lighter than those of most
aquifer ground waters (Tomascak et al., 2003), implying a significant source of <sup>6</sup>Li (Fig. 7B).
While the present-day temperatures (at about 130°C) were most probably not the temperature
during illitization, as already discussed in an earlier section, Li released from organics at as
low as 80°C (Williams et al., 2015), the fractionation factor (1000 lnα) for Li changing by
only -3‰ between 80°C and 130°C (Williams and Hervig, 2006), which appears to not
account for the observed changes of up to about -20‰ in the δ<sup>7</sup>Li of illite. However, similar

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1009 $\delta^7$Li compositions of water have been calculated from those of illite in bentonites from
1010 Wattenburg gas field in the Denver Basin (Williams et al., 2015).

1011 At higher temperatures (>130°C), most illite-rich fractions yield positive $\delta^7$Li values
1012 (up to about +11‰), while the Li contents are systematically low, below 20 µg/g. At lower
1013 temperatures (<115°C), the size fractions yield systematically negative $\delta^7$Li values and most
1014 have higher Li concentrations (up to 140 µg/g; Fig. 5). Increasing temperature could increase
1015 the fluid $\delta^7$Li if evaporites dissolved at higher temperature. Otherwise, the implied source of
1016 $^6$Li-enrichment in the nucleating illite results from maturing organic matter that produced the
1017 -14 to -22‰ $\delta^7$Li and 115-140 µg/g Li content of the finer Ptr47 fractions. The impact of
1018 saline fluids was also detected in the host sample, which in fact does not seem to have
1019 modified the Li contents and isotopic composition, as the Li signatures of the Sen21 fractions
1020 do not outline the same low $\delta^7$Li and high Li contents. The overall comparison of the Li
1021 contents and $\delta^7$Li from studied nano-sized fractions shows major correlations (Fig. 5). The
1022 less thermally mature bentonites contain crystals that are characterized by high smectite
1023 contents, low Li contents of 10-40 µg/g with low $\delta^7$Li of -13 to -16‰, or a Li content of about
1024 140 µg/g with even lighter $\delta^7$Li of -22‰. Calculation of the $\delta^7$Li of the fluids relative to
1025 temperature change is probably only of a few ‰ by taking the average burial temperature for
1026 illitization in the bentonite beds. This decreasing $\delta^7$Li with increasing Li content in illite
1027 suggests equilibration with fluids that contained higher $^6$Li. Possibly the organic-rich
1028 sediments preferentially released $^6$Li, because clay minerals in the shale source rock retained
1029 $^6$Li in the silicate structure up to much higher temperatures of >350°C (e.g., Chan et al.,
1030 2002). As for B, complementary information is provided by Li results of organics from our
1031 experiment of progressive aqueous pyrolysis of a <2 µm clay fraction from a hydrocarbon
1032 source shale of California. As no mineral phase was affected by progressive heating (Clauer
1033 et al., 2014b), the released Li amounts with increasing $\delta^7$Li relate to the progressively
1034 matured organics. The data show that the released Li amounts almost doubled, from about 20
1035 to 37 µg/g, when heating increased from 310°C during 72h to 365°C during 216h (Fig. 10B),
1036 while the $\delta^7$Li of the <1 µm fraction decreased slightly from -4 to -8‰. This confirms that
1037 the kerogen fraction left in the source rock was $^7$Li-enriched, as it appears to have
1038 preferentially released $^6$Li to fluids during thermal maturation (Williams et al., 2013).

1039

1040 The $\delta^7$Li and Li contents of illite relative to its K-Ar ages
The δ<sup>7</sup>Li values appear widely scattered, while the Li contents are constrained relative to crystallization timing. The Li contents plot along a lower trend from about 20 µg/g at 10-11 Ma ago to about 10 µg/g at 2 Ma ago (Fig. 13A). A few data are scattered above, except again, those of the three Cic1/11 fractions. Concerning the δ<sup>7</sup>Li, sample Cic1/18 yields the highest values at about +10‰ during the second illitization episode. Two additional trends can be suggested: a lower one at about -15‰ with a slight decreasing slope at younger ages, the second trend yielding a steeper slope from +7‰ at about 11 Ma ago to about -8‰ at 2.5 Ma ago (Table 2; Fig. 13B).

1049

1050 **NMR contribution to the illitization process during crystal growth**

Given that B is incorporated in a tetrahedral coordination of authigenic illite crystals from diagenetic environments (Spivack et al., 1987; Williams et al., 2001; 2007), and that trigonal B is mainly present in smectite interlayers (Williams et al., 2007), the presence of trigonal B in the intermediate sized fractions may be considered as a preliminary diagnostic for the occurrence of residual smectite remnants during growing of illite crystals. This conclusion is supported by the relationships between the [BO₄]/[BO₃] and illite/smectite ratios, as well as the K₂O contents. Therefore, it can be assumed that the relative increase of tetrahedral B detected in the size fractions by NMR relates to increasingly ordered-illite crystals relative to initially disordered smectite crystals. The lack of crystallographic ordering in the smectite crystals is confirmed, in turn, by the absence of “tooth-like” shaped BO₃ peaks expected for B in any ordered trigonal coordination (Stebbins and Xue, 2014).

The NMR data obtained here, still on a limited number of size fractions, may then help to improve the schematic process of illite nucleation and growth. Indeed, a disordered crystal environment is clearly identified by the occurrence of trigonal B. These disordered crystal sites seem to be independent at least of the initial nucleation as they occur mostly in the intermediate 0.02-0.05 µm fractions where most smectite has been detected by XRD, while either absent or of limited amount in the finer and coarser fractions. This result suggests that illite growth is spoiled by addition of “external” detrital smectite-type crystals at the 0.05-0.2 µm size. These fractions yield systematically the highest relative content of trigonal B, as well as the highest occurrence of remnant smectite layers. Therefore, the smectite crystals have to be of a slightly larger size than the nucleating illite crystals, representing in fact the authigenic mineral phase characteristic of the initial alteration of the volcanic glass. Alternatively, the lack or lower contents of smectite layers in the coarser fractions could then result from a progressive cannibalization of the smectite layers by the growing illite crystals, unless
remaining inert and being progressively scavenged by the overwhelming well-ordered illite layers.

Based on these preliminary NMR data, the smectite layers seem not to interfere in the nucleation of the illite crystals, as they are only detected at a slightly coarser crystal size than the initial illite nuclei. Disappearance of the trigonal B in the coarsest size fractions suggests then a dissolution process of the smectite layers, even if a progressive disappearance due to an overwhelming occurrence of illite layers cannot yet be discarded. In turn, the intimate aspect of illitization is still not clarified, as it could include, as suggested by Eberl et al. (1998b), larger crystals (of smectite-type) with defects on which new illite material crystallized. In summary, the preliminary NMR analyses that were obtained here on diagenetic illite open a promising potential to decipher further intimate details of the illitization process, as well as of the plausible alteration of associated detrital clay crystals. Of course, more investigations are necessary to describe more precisely these steps and quantify more precisely the B contents and types of the initial smectite layers and of that incorporated into the illite crystals. However, the limited available database points also to a somewhat variable chemical environment of B (in terms of the second neighbor's nature) that needs also to be addressed more precisely.

## CONCLUSIONS

Boron and Li contents and isotopic compositions were determined in nucleating and growing nanometer-sized crystals (<0.02 to 0.2 µm) of mixed-layered illite-smectite from bentonite beds of the East Slovak Basin. A few solid-state Nuclear Magnetic Resonance (NMR) spectroscopic determinations of three size fractions from three bentonite beds were performed for a preliminary evaluation of these data relative to the B isotope study. The B and Li contents are positively correlated in most of the analyzed nano-sized fractions, meaning that their substitution during crystal growth is probably coincident, and that illitization progressed mostly via direct precipitation from interacting pore fluids and not by solid-state transformation. Preliminary NMR <sup>11</sup>B data outline an exciting potential to discriminate between the different illitization processes accompanied or not by contemporaneous alteration of earlier clay crystals.

Basically, both the temperature and chemical composition of fluids impact the B and Li isotopic fractionation with the interacting minerals at equilibrium. As B and Li are released from their source materials into the pore fluids, the observed trends result from their initial
interaction with the nucleating and growing illite crystals. The results call for a concept combining B contents of the nano-sized illite crystals resulting from chemical variations of the interacting fluids, and δ¹¹B changes due to variable crystallization temperatures. In fact, temperature was often quite constrained during crystallization in the present case, whereas ⁷Li and ¹¹B were varying considerably, suggesting that the changing fluid chemistry during the successive illitization episodes was the major factor impacting the B and Li isotope compositions of the nucleating illite crystals.

General trends combining B amounts and δ¹¹B values show that the bentonites beds collected near salt-bearing sediments contain illite with high B contents (475 µg/g) and heavy δ¹¹B values (+7‰). This combination calls for precipitation from saline brines of the illite crystals older than 10 Ma. The younger illite populations yield lower B contents (30-100 µg/g) with δ¹¹B decreasing from -5 to -17‰ as the B content increased, with one outlier of sample Cic1/11 sample. These combined changes suggest interactions between organic materials of the shale host rocks and the pore fluids that carried the elements released by the maturing organics (e.g., K, B, Li) into the bentonite beds after about 6 Ma ago. The initial saline fluids were apparently not enriched in Li, as the coarse and older illite fractions yield low Li content (<20 µg/g) and high δ⁷Li (+7‰), which reflects a water δ⁷Li composition near 23-25‰ similar to normal ground waters. Younger illite crystals yield higher Li contents (up to 140 µg/g) with decreasing δ⁷Li (-7 to -22‰).

The light element trends for illite crystals younger than 6 Ma show decreasing δ¹¹B and δ⁷Li as crystal size increases. This correlates with an increase of δ⁷Li and δ¹¹B in nearby kerogen as burial temperature increased, suggesting release of isotopically light B and Li from kerogen to the pore fluids during hydrocarbon generation when illitization was concomitant. Equilibration between the pore fluid isotopic composition and authigenic illite records changes in the fluid chemistry during crystal growth, and K-Ar dating of the illite size fractions indicates the timing of these important paleo-fluid changes related either to early dissolution of salt beds or to late hydrocarbon maturation and migration. Of significance is the obvious organic source of these important light elements (B and Li) that have not been considered in many studies of their global geochemical cycles.

1138

## 1139 ACKNOWLEDGEMENTS

1140
1141 We would like to sincerely thank R.L. Hervig (ASU) for supportive discussions during
1142 data acquisition and conceptual presentation of the preliminary results. Renewed sincere
1143 thanks are also due to Mrs. M. Zielinska for her determining contribution in completing the
1144 nanometer-size fractionation of the selected samples, and to R. Adriaens (KUL) for Sybilla
1145 modeling of the XRD patterns. We also like to acknowledge use of the Arizona State
1146 University National SIMS Facility supported by the US National Science Foundation (EAR-
1147 0948878), as well as the financial support by the TGIR-RMN-THC Fr3050 CNRS for
1148 completion of the NMR analyses. We also thank the two anonymous reviewers for help to
1149 improve a previous version of the draft.

1150

1151 REFERENCES

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1508 Williams L.B., Elliott W.C. and Hervig R.L. (2015) Tracing hydrocarbons in gas shale using lithium and boron isotopes: *Denver Basin USA, Wattenberg Gas Field. Chemical Geology*,
1510 417, 403-413.

1511 Wilson T.P. and Long D.T. (1993) Geochemistry and isotope chemistry of Michigan Basin
1512 brines: Devonian formations. Applied Geochemistry, 8, 81-100.

1513 Wunder B., Meixner A., Romer R.L., Wirth R. and Heinrich W. (2005) The geochemical
1514 cycle of boron: Constraints from boron isotope partitioning experiments between mica and
1515 fluid. Lithos, 84, 206-216.

1516 You C.F. and Chan L.H. (1996) Precise determination of lithium isotopic composition in low
1517 concentration natural samples. Geochimica et Cosmochimica Acta, 60, 909-915.

1518

1519

1520 FIGURE AND TABLE CAPTION

1521

1522 **Figure 1:** (A) Location of the East Slovak Basin in the Pannonian system of Neogene basins
1523 (after Kováč et al., 1995); (B) Enlargement of the square in map (A) with location of the
1524 studied wells (after Honty et al., 2004).

1525 **Figure 2:** XRD of oriented slides of three different separated clay fractions in (A) air-dried
1526 natural state and (B) ethylene-glycolated state of sample Ptr47 (from Honty et al., 2004).
1527 NS stands for not saturated and EG for ethylene-glycol saturation.

1528 **Figure 3:** The δ<sup>11</sup>B and δ<sup>7</sup>Li with the B and Li contents of the <2 µm components of the host
1529 shales of the bentonite beds. The gray symbols represent the bulk <2 µm fractions, the
1530 open symbols the clay component and the full symbols the organic components.

1531 **Figure 4:** The δ<sup>11</sup>B values relative to the B contents of the studied nano-sized fractions. The
1532 numbers from 1 to 3 refer to the grain size: 1 = <0.02, 2 = 0.02-0.05 and 3 = 0.05-0.1 µm
1533 (from Clauer et al., 2009).

1534 **Figure 5:** The δ<sup>7</sup>Li values relative to the Li contents of the studied nanometer-sized fractions.

1535 **Figure 6:** The B contents vs. Li contents of the studied nano-sized fractions.

1536 **Figure 7:** The NMR spectra of three size fractions from sample Ptr13 (A), sample Ptr47 (B)
1537 and sample Sen21 (C). The size of the fractions, and the BO<sub>3</sub> contents are outlined in the
1538 graphs.

1539 **Figure 8:** The δ<sup>11</sup>B (A) and burial depth (B) relative to the present-day temperature in the
1540 boreholes.

1541 **Figure 9:** The δ<sup>11</sup>B in the “salt-bearing” and “salt-free” illite separates relative to their
1542 crystallization temperatures based on their S<sub>XRD</sub> contents and their re-calculated
1543 temperatures of figure 7B.
1544 **Figure 10:** The B (A) and Li (B) data from aqueous-pyrolysis experiments of an oil-bearing
1545 shale from California (data from Clauer, Williams et al., unpublished data).
1546 Complementary information is available in Clauer et al. (2014b). The data of the removed
1547 B and Li were calculated graphically by simple deduction from untreated and heated
1548 fractions at the different pyrolysis temperatures.

1549 **Figure 11:** Correlations between (A) the B and K₂O contents, and (B) those of B and SiO₂
1550 from studied separates.

1551 **Figure 12:** The B contents (A) and δ¹¹B values (B) of the nanometer-sized fractions of the
1552 studied samples. The grey areas represent the different crystallization episodes of illite.
1553 The numbers from 1 to 3 refer to the grain sizes 1 = <0.02, 2 = 0.02-0.05, and 3 = 0.05-0.1
1554 µm.

1555 **Figure 13:** The Li contents (A) and δ⁷Li values (B) of the nanometer-sized fractions of the
1556 studied samples. The grey areas represent the different crystallization episodes of illite.

1557
1558 **Table 1:** Identification, stratigraphic age, burial depth and expandability in % of smectite
1559 (labeled %Sm) of the studied nanometer-sized fractions.

1560 **Table 2:** Boron and Li contents, and δ¹¹B and δ⁷Li values of the studied nanometer-sized
1561 fractions. The K-Ar ages and δ¹⁸O values are also given (of Clauer et al., 2014a).
<figure><img src="image_1.png" /></figure>

A

<figure><img src="image_2.png" /></figure>

Neovolcanic areas

<figure><img src="image_3.png" /></figure>

pre - Neogene of inner orogenic zones

<figure><img src="image_4.png" /></figure>

East Slovak Basin

<figure><img src="image_5.png" /></figure>

Tertiary Molasse Zone

<figure><img src="image_6.png" /></figure>

Flysch Belt

<figure><img src="image_7.png" /></figure>

Neogene basins

<figure><img src="image_8.png" /></figure>

B

N

10 km

<figure><img src="image_9.png" /></figure>

Neogene sediments

<figure><img src="image_10.png" /></figure>

Neovolcanites
Ptruksa 47

<figure><img src="image_11.png" /><figcaption>B</figcaption></figure>

<figure><img src="image_12.png" /></figure>

2 theta degrees
<figure><img src="image_13.png" /></figure>

<figure><img src="image_14.png" /></figure>
<figure><img src="image_15.png" /></figure>

O Ptr47

▲ Cic1/15

□ Bid1

● Trh1

□ Cic1/11

■ Cic1/18

△ Ptr13

○ Sen13

△ Sen21

X organics
<figure><img src="image_16.png" /></figure>

○ Ptr47

▲ Cic1/15

□ Bid1

● Trh1

■ Cic1/11

■ Cic1/18

△ Ptr13

○ Sen13

▲ Sen21

X organics
Ptr13

Fraction <0.02 µm

%BO₃ = 7%

Fraction 0.02-0.05 µm

%BO₃ = 20%

Fraction 0.05-0.2 µm

%BO₃ = 9%

(A)

Ptr47

Fraction <0.02 µm

%BO₃ = 8%

Fraction 0.02-0.05 µm

%BO₃ = 25%

Fraction 0.05-0.2 µm

%BO₃ = 8%

(B)

Sen21

Fraction <0.02 µm

%BO₃ = 21%

Fraction 0.02-0.05 µm

%BO₃ = 35%

Fraction 0.05-0.2 µm

%BO₃ = 10%

(C)

$^{11}$B frequency (ppm)
Temperature (° C)

<figure><img src="image_17.png" /></figure>

Smectite in I/S (%)

<figure><img src="image_18.png" /></figure>
<figure><img src="image_19.png" /></figure>

Ptr13

Cic1/15

Trh1

Cic1/11

Cic1/18

Sen13

Sen21
<figure><img src="image_20.png" /></figure>

<figure><img src="image_21.png" /></figure>
<figure><img src="image_22.png" /></figure>

<figure><img src="image_23.png" /></figure>
<figure><img src="image_24.png" /></figure>

<figure><img src="image_25.png" /></figure>

B
<figure><img src="image_26.png" /></figure>

<figure><img src="image_27.png" /></figure>
<table><caption>Table 1</caption><thead><tr><th>Sample identification<br>(<2 mm)</th><th>Stratigraphic age<br>(in Ma)</th><th>Burial depth<br>(m)</th><th>Meas. temp<br>(°C)</th><th>SXRD<br>(%)</th></tr></thead><tbody><tr><td colspan="5"><strong>Salt-free samples</strong></td></tr><tr><td>Cicarovce 1/11</td><td>Lower Sarmatian (12.5-11.5)</td><td>1614</td><td>97</td><td>90</td></tr><tr><td>Cicarovce 1/15</td><td>Lower Sarmatian (12.5-11.5)</td><td>2002</td><td>116</td><td>28</td></tr><tr><td>Cicarovce 1/18</td><td>Lower Sarmatian (12.5-11.5)</td><td>2296</td><td>131</td><td>51</td></tr><tr><td>Trhoviste 1/37</td><td>L. Badenian-U. Carpathian (17-15.5)</td><td>3015</td><td>165</td><td>18</td></tr><tr><td colspan="5"><strong>Salt-bearing samples</strong></td></tr><tr><td>Bidovce 1 3/1</td><td>Lower Badenian (16.5-15.5)</td><td>1557</td><td>n.d.</td><td>70</td></tr><tr><td>Ptruksa 47 9/7</td><td>Upper Badenian (14.5-13.5)</td><td>2106-2115</td><td>n.d.</td><td>25</td></tr><tr><td>Ptruksa 13 14</td><td>Upper Badenian (14.5-13.5)</td><td>2095-2103</td><td>115</td><td>13</td></tr><tr><td>Senné 13 5A</td><td>Middle Badenian (15.5-14.5)</td><td>2402</td><td>132</td><td>15</td></tr><tr><td>Senné 21 4/1</td><td>Middle Badenian (15.5-14.5)</td><td>2311</td><td>134</td><td>10</td></tr></tbody></table>

Meas. temp. stands for measured temperature, SXRD for expandability measured by XRD
<table><thead><tr><th>Sample description</th><th>Size (µm)</th><th>B (µg/g)</th><th>δ11B (±1σ)</th><th>Li (µg/g)</th><th>δ7Li (±1σ)</th><th>K-Ar (Ma ±2σ)</th><th>δ18O*</th></tr></thead><tbody><tr><td colspan="8"><strong>Bentonites</strong></td></tr><tr><td>Ptruksa 47 9/7 "salted"</td><td>&lt;0.02</td><td>146</td><td>-10.5 (0.4)</td><td>115</td><td>-13.6 (1.0)</td><td>9.0 (0.2)</td><td></td></tr><tr><td>2106 m</td><td>0.02-0.05</td><td>137</td><td>-12.0 (0.6)</td><td>139</td><td>-21.6 (1.2)</td><td>5.9 (0.4)</td><td></td></tr><tr><td></td><td>0.05-0.2</td><td>158</td><td>-11.7 (2.0)</td><td>5</td><td>1.8 (3.2)</td><td>10.2 (0.2)</td><td></td></tr><tr><td>Cicarovce 1/15</td><td>&lt;0.02</td><td>133</td><td>-2.1 (0.6)</td><td>18</td><td>-16.6 (1.8)</td><td>3.2 (0.1)</td><td>12.6</td></tr><tr><td>2002 m</td><td>0.02-0.05</td><td>261</td><td>-3.2 (0.6)</td><td>8</td><td>-7.9 (0.2)</td><td>4.3 (0.1)</td><td>12.2</td></tr><tr><td></td><td>0.05-0.2</td><td>157</td><td>-9.9 (0.8)</td><td>106</td><td>-17.4 (1.4)</td><td>4.6 (0.1)</td><td>12.2</td></tr><tr><td>Senné 13 5A "salted"</td><td>&lt;0.02</td><td>108</td><td>-15.4 (1.4)</td><td>9</td><td>-8.6 (.8)</td><td>2.2 (0.1)</td><td>18.7</td></tr><tr><td>2402 m</td><td>0.02-0.05</td><td>105</td><td>-17.6 (0.6)</td><td>7</td><td>-2.8 (0.6)</td><td>3.0 (0.1)</td><td>17.6</td></tr><tr><td></td><td>0.05-0.2</td><td>281</td><td>1.6 (1.2)</td><td>7</td><td>1.6 (0.8)</td><td>5.1 (0.1)</td><td>18.4</td></tr><tr><td>Bidovce 1 3/1</td><td>&lt;0.02</td><td>50</td><td>-5.9 (1.8)</td><td>12</td><td>-13.8 (1.0)</td><td>7.4 (0.9)</td><td></td></tr><tr><td>1565 m</td><td>0.02-0.05</td><td>52</td><td>-8.0 (1.0)</td><td>8</td><td>-16.0 (0.6)</td><td>6.3 (0.4)</td><td></td></tr><tr><td></td><td>0.05-0.2</td><td>40</td><td>-10.7 (0.6)</td><td>7</td><td>-14.5 (1.6)</td><td>10.2 (0.5)</td><td></td></tr><tr><td>Senné 21 4/1 "salted"</td><td>&lt;0.02</td><td>n.d.</td><td>6.8 (1.8)</td><td></td><td></td><td>4.4 (0.1)</td><td></td></tr><tr><td>2311 m</td><td>0.02-0.05</td><td>296</td><td>-1.5 (2.0)</td><td></td><td></td><td>5.5 (0.1)</td><td></td></tr><tr><td></td><td>0.05-0.2</td><td>310</td><td>4.1 (1.4)</td><td></td><td></td><td>4.0 (0.1)</td><td></td></tr><tr><td>Cicarovce 1/11</td><td>&lt;0.02</td><td>246-415</td><td>-13.5 (1.0)</td><td>42-43</td><td>-13.2 (1.4)</td><td>0.6 (0.8)</td><td>14.9</td></tr><tr><td>1614 m</td><td>0.02-0.05</td><td>324</td><td>-20.1 (1.6)</td><td>38</td><td>-13.1 (1.0)</td><td>0</td><td>15.1</td></tr><tr><td></td><td>0.05-0.2</td><td>471-491</td><td>-26.9 (0.6)</td><td>44-60</td><td>-7.1 (0.6)</td><td>1.0 (0.6)</td><td>14.6</td></tr><tr><td>Ptruksa 13 14 "salted"</td><td>&lt;0.02</td><td>198</td><td>-0.8 (1.2)</td><td></td><td></td><td>7.2 (0.1)</td><td>15.1</td></tr><tr><td>2100 m</td><td>0.02-0.05</td><td>349</td><td>-0.6 (0.8)</td><td></td><td></td><td>7.6 (0.2)</td><td>15.5</td></tr><tr><td></td><td>0.05-0.2</td><td>343</td><td>4.8 (0.4)</td><td></td><td></td><td>8.0 (0.1)</td><td>14.8</td></tr><tr><td>Cicarovce 1/18</td><td>&lt;0.02</td><td>30-29</td><td>-4.0 (0.4)</td><td>3</td><td>8.3 (1.0)</td><td>7.4 (0.4)</td><td></td></tr><tr><td>2296 m</td><td>0.02-0.05</td><td>30</td><td>-4.7 (1.0)</td><td>3</td><td>8.6 (1.0)</td><td>2.6 (0.9)</td><td></td></tr><tr><td></td><td>0.05-0.2</td><td>49</td><td>-5.7 (1.0)</td><td>12</td><td>10.9 (0.2)</td><td>7.1 (0.3)</td><td></td></tr><tr><td>Trhoviste 1/37</td><td>0.05-0.1</td><td>366</td><td>-10.5 (0.8)</td><td>31</td><td>-0.1 (1.6)</td><td>10.2 (0.5)</td><td></td></tr><tr><td>3015 m</td><td>0.3-0.5</td><td>386</td><td>-8.8 (1.6)</td><td>12</td><td>7.0 (0.4)</td><td>11.3 (0.5)</td><td></td></tr><tr><td colspan="8"><strong>Shales</strong></td></tr><tr><td>Cicarovce 1/20</td><td>bulk &lt;2</td><td>172</td><td>-6</td><td>51</td><td>3</td><td></td><td></td></tr><tr><td></td><td>clay</td><td>109</td><td>-8</td><td>49</td><td>-1</td><td></td><td></td></tr><tr><td></td><td>organic</td><td>63</td><td>-4</td><td>2</td><td>XX</td><td></td><td></td></tr><tr><td>Trhoviste 1/37</td><td>bulk &lt;2</td><td>243</td><td>-8</td><td>75</td><td>-1</td><td></td><td></td></tr><tr><td></td><td>clay</td><td>185</td><td>-9</td><td>64</td><td>-2</td><td></td><td></td></tr><tr><td></td><td>organic</td><td>58</td><td>-6</td><td>11</td><td>5</td><td></td><td></td></tr><tr><td>Senné 13 5</td><td>bulk &lt;2</td><td>217</td><td>-3</td><td>61</td><td>-1</td><td></td><td></td></tr><tr><td></td><td>clay</td><td>177</td><td>-7</td><td>55</td><td>-2</td><td></td><td></td></tr><tr><td></td><td>organic</td><td>40</td><td>16</td><td>6</td><td>7</td><td></td><td></td></tr><tr><td>Senné 21</td><td>bulk &lt;2</td><td>276</td><td>-6</td><td>60</td><td>-5</td><td></td><td></td></tr><tr><td></td><td>clay</td><td>226</td><td>-4</td><td>50</td><td>-3</td><td></td><td></td></tr><tr><td></td><td>organic</td><td>50</td><td>-12</td><td>10</td><td>-17</td><td></td><td></td></tr></tbody></table>

XX stands for un-realistic value and * for the data by Clauer et al. (2014)