Jan Środoń

Oksana Kuzmenkova

Jan J Stanek

Sabine Petit

Daniel Beaufort

H Albert Gilg

Sirle Liivamägi

Magdalena Goryl

Leszek Marynowski

Marek Szczerba

Hydrothermal alteration of the Ediacaran Volyn-Brest volcanics on the western margin of the East European Craton

Keywords: aluminoceladonite, chlorophaeite, cyanobacteria, griffithite, ferrosaponite, Martian smectite, palagonite

The Ediacaran continental flood basalts and associated tuffs were studied to identify and quantify alteration processes by means of XRD and chemistry, supplemented by Mössbauer and FTIR spectroscopies, petrography, oxygen and iron isotopes, K-Ar dating, and organic geochemistry. Two superimposed alteration processes were identified: the Ediacaran hydrothermal alteration, induced by meteoric waters, heated and put in motion by the cooling basalt, and the Caledonian and/or Variscan potassic alteration. The degree of basalt alteration was quantified using as an index the sum of primary minerals in the bulk rock. The sequence of minerals dissolved and crystallized during the hydrothermal alteration was established. The alteration resulted in the loss of Ca (dissolution of plagioclases), compensated by the gain of water and Mg (crystallization of clays), and proceeded from the edges of the basalt flows in an oxidizing environment, evidenced by the increasing amount of hematite and Fe 3+ /Fe 2+ ratio of the bulk rock. Cyanobacteria were active in the hydrothermal system, most probably responsible for the measured negative  56 Fe values and more reducing conditions at the stage of intense alteration. Chlorophaeite (palagonite), following quartz as the earliest petrographically identifiable basalt alteration product was found to vary systematically from fully isotropic to birefringent. The chlorophaeite was identified as a mixture of Femontmorillonite and Fe-saponite, identical with griffithite and oxysmectites, probably preceded by a finer-grained ferrosaponite at the isotropic stage. REE content of chlorophaeite indicates basaltic volcanic glass (sideromelane) as the major source of material. REE in clays are contained mostly in the dioctahedral smectite, while in the bulk rock mostly in phosphates. The smectite characteristics and Mg enrichment are indicative of the hydrothermal basalt alteration process, which perhaps was dominant also on Mars. Both mineral and chemical composition of tuffs vary continuously from basaltic to felsic, the latter close to the measured rhyodacite composition, dominated by quartz and feldspars. The basaltic tuffs resemble the most altered basalts but contain also abundant albite and chlorite, indicative of higher alteration temperatures, up to 220 o C. Tuff composition indicates stronger component of felsic volcanism in the trap formation than evidenced by the preserved bodies of effusive rocks.

Introduction

Sub-surface (non-related to weathering) alteration of basalts is a widespread and intensely studied phenomenon, known from three major Earth environments: submarine hydrothermal, burial diagenetic/metamorphic, and continental hydrothermal.

A good example of submarine hydrothermal alteration within a 60-165 o C measured temperature range is the 1300 m profile of DSDP Hole 504B south of the Costa Rica Rift in the Pacific [START_REF] Alt | Hydrothermal alteration of a 1 km section through the upper oceanic crust, DSDP Hole 504B: mineralogy, chemistry, and evolution of seawater-basalt interactions[END_REF][START_REF] Shau | Phyllosilicates in hydrothermally altered basalts from DSDP Hole 504B, Leg 83a TEM and AEM study[END_REF]. A clear alteration sequence was observed: iron oxides, saponite, celadonite, phillipsite, anhydrite, and aragonite formed below 150 o C, and mixed-layer saponite-chlorite, corrensite, corrensite/chlorite, and finally chlorite, minor talc, quartz, pyrite, epidote, laumontite, heulandite, prehnite, actinolite, and albite formed in deeper, hotter parts. Altered basalt becomes a sink for Mg and K from seawater, affecting the element mass balance of the oceans. From the saponitic to chloritic end of the trioctahedral clay sequence, Si/(Si+Al) ratio decreases, while Fe/(Mg+Fe) increases. In another submarine site (801), [START_REF] Alt | Low-temperature hydrothermal alteration of Jurassic ocean crust, Site 801[END_REF] observed a range of alteration products from pure saponite to almost pure dioctahedral smectite, accompanied by pyrite and calcite. Other samples contain celadonite, glauconite, and iron oxides. With progressing alteration (measured by H2O + ), an increase in   O, iron oxidation and K2O is observed, while CaO decreases, but MgO does not show a clear trend.

A particularly clear sequence of burial metamorphic alteration zones was documented [START_REF] Schmidt | Regional and local patterns of low-grade metamorphism in the North Shore Volcanic Group, Minnesota, USA[END_REF][START_REF] Schmidt | Metamorphic grade and porosity and permeability controls on mafic phyllosilicate distributions in a regional zeolite to greenschist facies transition of the North Shore Volcanic Group, Minnesota[END_REF] from the 8 km thick pile of the Keweenawan Middle Proterozoic continental subaerial rift basalts, outcropping on the shore of Lake Superior, and known for native copper deposits. At the top of sequence, saponite is dominant, and going down section, random chlorite-smectite and corrensite appear, accompanied by laumontite and albite, then chlorite with prehnite and pumpellyite, and finally epidote and actinolite (greenshist facies). Albite is the most abundant alteration product of Ca-rich plagioclase. Fe oxides were not reported. Celadonite occurs sporadically. Like in the submarine alteration, from the saponitic to chloritic end of the trioctahedral clay sequence Si/(Si+Al) ratio decreases, while Fe/(Mg+Fe) increases. Strong control of porosity and permeability over the extent of alteration is evident: massive flow centres have less advanced alteration than porous, thus permeable tops and bottoms.

Subaerial hydrothermal alteration of basalts has been most thoroughly studied in Iceland, where data measured in active hydrothermal systems were used to assess temperature ranges of subsequent alteration zones, analogous to burial and marine hydrothermal [START_REF] Franzson | Chemical transport in geothermal systems in Iceland: Evidence from hydrothermal alteration[END_REF] geothermal well in Hawaii [START_REF] Stone | Hydrothermal alteration of basalts from Hawaii Geothermal Project Well-A, Kilauea, Hawaii[END_REF]. This review demonstrates that the alteration of basalts in three different environments proceeds via a very similar sequence of mineral assemblages, which indicates dominant control of the rock chemistry over the alteration mineralogy, with porosity and permeability affecting the reaction progress, and in particular the pathway of smectite to chlorite transition [START_REF] Robinson | Reaction pathways and reaction progress for the smectite-to-chlorite transformation: evidence from hydrothermally altered metabasites[END_REF].

Similar qualitative mineralogical data (e.g. [START_REF] Juskowiakowa | Upper Precambrian sediments[END_REF][START_REF] Kuzmenkova | Zeolites of the Vendian basalts of Belarus[END_REF][START_REF] Derevska | Geological genetic model of ore-forming sysytem andprospecting criteria of native copper in the Volynian flood basalts[END_REF][START_REF] Emetz | Geology, mineralogy and origin of Zhyrychi native copper deposit (North-Western Ukraine)[END_REF]Kuzmenkova et al., 2006;[START_REF] Melnychuk | Hydrothermal mineralogical zonality and metamorphism in copperbearing Vendian flood basalts of the Volyno-Podolian plate[END_REF][START_REF] Shumlyanskyy | On the character of secondary alteration of volcanogenic rocks of the Volynian flood basalt province[END_REF][START_REF] Skakun | Formation of zeolite mineralization in the hydrothermal rocks of of the flood basalts of Volynian series[END_REF] are available also for the Ediacaran Volyn-Brest volcanics of the East European Craton (EEC; Figure 1) and they inspired the present study. Our study was focused on adding a quantitative mineralogical dimension to the knowledge of hydrothermal basalt alteration, on understanding the conditions of this alteration, and on discriminating the effects of younger alteration phenomena. The clay alteration products were studied in most detail to characterize the substrate on which the Ediacaran weathering of basalts took place [START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF]. It was also hoped that better understanding of the Ediacaran basalt hydrothermal alteration may be useful as a model for the origin of clays on Mars. They have been detected both by remote and in-situ methods as the main products of basalt alteration in the Noachian age (e.g., Ehlmann et al., 2011a,b;[START_REF] Carter | Hydrous minerals on Mars as seen by the CRISM and OMEGA imaging spectrometers: Updated global view[END_REF][START_REF] Chemtob | Synthesis and structural characterization of ferrous trioctahedral smectites: Implications for clay mineral genesis and detectability on Mars[END_REF]. The origin of these clay minerals is currently debated, and potential formation environments include subsurface hydrothermal systems, subaerial weathering profiles, lacustrine environments, and steam or a supercritical atmosphere of water and CO2 (e.g., Ehlmann et al., 2011a,b;[START_REF] Vaniman | Mineralogy of a mudstone at Yellowknife Bay, Gale crater[END_REF][START_REF] Bishop | Surface clay formation during short-term warmer and wetter conditions on a largely cold ancient Mars[END_REF]Cannon et al., 2018). Here, we identify mineralogical and geochemical signatures of subsurface hydrothermal clay deposits that can be used to constrain clay formation environments on Mars.

Study area

The Volyn-Brest Ediacaran flood basalts, up to 500 m of lavas and volcanoclastics, cover 140, 000 km 2 of the western margin of the East European Craton. The original volcanic cover extended much further east and north, as evidenced by isolated erosional remnants in these areas (Figure 1). Towards the southeast, the volcanic cover continues to an unknown distance, deeply buried below Paleozoic sediments of the Teisseyre-Tornquist Zone [START_REF] Pacześna | The evolution of late Ediacaran riverine-estuarine system in the Lublin-Podlasie slope of the East European Craton, southeasten Poland[END_REF], which separates the Precambrian from the Paleozoic platform. In Volyn (northwest Ukraine north of Lvov), the Ediacaran basalts occur over ca. 80 km 2 at the surface or under a thin cover of the Cretaceous rocks and have been used as building stones since XVII century [START_REF] Rajchel | Antique cobblestone pavement in Krakow[END_REF]. In other regions, they are covered by the Late Ediacaran clastic sediments and younger deposits.

The Volyn-Brest trap region is a continental flood basalt province, analogous to the well-known Columbia River, Deccan, Paraná, or Siberian traps [START_REF] Kuzmenkova | Comparison of Neoproterozoic Volyn-Brest magmatic province with large world continental plate basalt provinces, the nature of low-and high-Ti basite magmatism[END_REF]. It is considered a product of the supercontinent Rodinia rifting, which detached Amazonia from Baltica by opening of the Tornquist Ocean in the Late Ediacaran [START_REF] Poprawa | Late Vendian-Early Palaeozoic tectonic evolution of the Baltic Basin: regional tectonic implications from subsidence analysis[END_REF][START_REF] Nawrocki | Development of Trans-European Suture Zone in Poland: from Ediacaran rifting to Early Palaeozoic accretion[END_REF]. Deep root bodies of the basalt province were identified recently by geophysical methods in the Teisseyre-Tornquist Zone, SE of Brest (Mazur et al., 2018a).

Volyn basalts have been studied since the mid-19th century, and particular interest was generated by finding a native copper mineralization [START_REF] Małkowski | On discovery of the native copper deposit in Volyn[END_REF]. These early studies (conf. Białowolska et al., 2002 for references) were carried out on rocks from the Volyn outcrops, later also on the core materials from Belarus, Poland and Ukraine [START_REF] Ushakova | Lower Paleozoic trap formation in western part of the Russian platform[END_REF][START_REF] Juskowiak | Volcanic and accompanying sedimentary rocks from boreholes Kruszyniany and Mielnik[END_REF][START_REF] Маkhnach | Volcanogenic formation of the Upper Proterozoic (Vendian) of Belarus[END_REF][START_REF] Juskowiakowa | Basalts of eastern Poland[END_REF][START_REF] Compston | Numerical ages of volcanic rocks and the earliest faunal zone within the Late Precambrian of east Poland[END_REF][START_REF] Białowolska | Neoproterozoic flood basalts of the upper beds of the Volhynian Series (East European Craton)[END_REF][START_REF] Bakun-Czubarow | Neoproterozoic flood basalts of Zabolottya and Babino Beds of the volcanogenic Volhynian Series and Polesie Series dolerites in the western margin of the East European Craton[END_REF][START_REF] Emetz | Geological framework of the Volhyn copper fields with a review of the Volhyn flood basalt province (western margin of the East-European Craton)[END_REF][START_REF] Kuzmenkova | Petrography of magmatic rocks of the trap formation of Belarus[END_REF][START_REF] Nosova | Neoproterozoic Volhynia-Brest magmatic province in the western East European Craton: withing-plate magmatism in ancient suture zone[END_REF][START_REF] Kuzmenkova | Material composition of the Khotislavsky cooper occurrence[END_REF]and many others).

Recent investigations, summarized in a monograph (Shumlyanskyy, ed., 2006), were centered on the genesis of native copper mineralization. All of these studies were devoted mainly to geochronology, geochemistry, and petrographic observations of mineral successions. [START_REF] Shumlyanskyy | The U-Pb zircon and baddeleyite ages of the Neoproterozoic Volyn Large Igneous Province: implication for the age of the magmatism and the nature of a crustal contaminant[END_REF] reviewed the geochronological literature. Early K-Ar measurements of the whole rock gave a broad range of 690-540 Ma. The 40 Ar/ 39 Ar whole rock method yielded two age groups: 590-560 Ma (magma crystallization) and 393-369 Ma (potassic alteration). The Rb-Sr isochron method applied to four whole rock samples gave an imprecise age of 552±59 Ma, while U-Pb zircon ages range from 551±4 Ma to 573±14 Ma.

Geochemical studies identified the trap rocks as within plate continental basalts and documented vertical zonation of the trap: picrobasalts and subalkaline olivine basalts in the lower basalt flows are covered locally (Figure 1) by felsic rocks (andesite, dacite, rhyodacite, trachyandesite), and tholeiites dominate the upper flows, all accompanied by more or less abundant tuffs [START_REF] Bakun-Czubarow | Neoproterozoic flood basalts of Zabolottya and Babino Beds of the volcanogenic Volhynian Series and Polesie Series dolerites in the western margin of the East European Craton[END_REF][START_REF] Kuzmenkova | Mineralogy and petrogenesis of Vendian basalts and dolerites of Belarus[END_REF][START_REF] Nosova | Neoproterozoic Volhynia-Brest magmatic province in the western East European Craton: withing-plate magmatism in ancient suture zone[END_REF].

The Ediacaran paleosols are developed both on the lower [START_REF] Kuzmenkova | Petrology and correlation of trap formations of the Vendian in the adjacent areas of Belarus and Ukraine[END_REF] and on the upper basalts [START_REF] Levykh | Weathering crusts of the western part of the East European platform[END_REF][START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF], documenting the continental, subaerial origin of the volcanic sequence.

Since the findings of native copper [START_REF] Małkowski | On discovery of the native copper deposit in Volyn[END_REF] and celadonite [START_REF] Kamieński | Volyn basalts[END_REF], later confirmed by [START_REF] Lazarenko | About seladonite from Volyn basalts[END_REF], it was realized that Volyn basalts underwent widespread hydrothermal alteration of variable intensity. [START_REF] Shumlyanskyy | On the character of secondary alteration of volcanogenic rocks of the Volynian flood basalt province[END_REF] recognized this variable degree of alteration and characterized geochemically the alteration sequence: oxidation of iron; increase of MgO, K2O, LOI, Rb, Ba, Zr, Nb, Ta, Th, U; decrease of CaO and Sr. They observed more intense alteration of tuffs compared to the flood basalts, but they interpreted the exceptionally high MgO content of some tuffs as reflecting the primary volcanic composition. [START_REF] Melnychuk | Hydrothermal mineralogical zonality and metamorphism in copperbearing Vendian flood basalts of the Volyno-Podolian plate[END_REF] documented carbonate, zeolite, and prehnite-pumpellyite zones of hydrothermal alteration in the flood basalts and volcanic tuffs.

In Volyn and Belarus, i.e. outside the Teisseyre-Tornquist Zone, the basalts have never been deeply buried, as indicated by regional organic geochemistry data for overlying sediments (Nehring-Lefeld et al., 1977) and confirmed recently by studying the Ediacaran paleosols developed on basalts [START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF] and hopane biomarkers in the overlying Ediacaran sediments [START_REF] Goryl | Exceptional preservation of hopanoid and steroid biomarkers in Ediacaran sedimentary rocks of the East European Craton[END_REF]. Initially, a "fresh" appearance of the unaltered sections was used as evidence of the lack of regional metamorphic or advanced diagenetic changes [START_REF] Ushakova | Lower Paleozoic trap formation in western part of the Russian platform[END_REF][START_REF] Juskowiak | Volcanic and accompanying sedimentary rocks from boreholes Kruszyniany and Mielnik[END_REF][START_REF] Маkhnach | Volcanogenic formation of the Upper Proterozoic (Vendian) of Belarus[END_REF]. [START_REF] Juskowiakowa | Upper Precambrian sediments[END_REF] observed that the alteration develops from the upper and lower surfaces of the lava flows, it may be complete in thin flows and tuffs or tectonic zones, and it involves the development of chlorite, illite and celadonite. The alteration sequence was characterized by the microscopic studies of Volyn native Cu deposits [START_REF] Derevska | Geological genetic model of ore-forming sysytem andprospecting criteria of native copper in the Volynian flood basalts[END_REF][START_REF] Emetz | Geology, mineralogy and origin of Zhyrychi native copper deposit (North-Western Ukraine)[END_REF][START_REF] Skakun | Formation of zeolite mineralization in the hydrothermal rocks of of the flood basalts of Volynian series[END_REF]. Plagioclase (andesineanorthite), clinopyroxene, magnetite, ilmenite, titanomagnetite, volcanic glass, and rare olivine (often recognized only as prehnite pseudomorphs) were identified as primary basaltic components. Clinochlore, hematite, albite, and pumpellyite were identified as pre-ore minerals; prehnite, laumontite, wairakite, quartz, and K-spar as syn-ore minerals; and quartz, stilbite, mordenite, heulandite, thomsonite, tri-and dioctahedral smectite, chlorite, vermiculite, calcite, analcime, barite, and kaolinite as post-ore minerals. Similar composition, including a wider range of zeolite minerals, was reported by [START_REF] Kuzmenkova | Zeolites of the Vendian basalts of Belarus[END_REF]. Kuzmenkova et al. (2006) characterized in detail glassy-looking accumulations of dark material, described from various basalts under the name "chlorophaeite" and present even in very fresh-looking basaltsthe equivalent of palagonite (conf. review by [START_REF] Stroncik | Palagonitea review[END_REF], of the late-magmatic lithogel of [START_REF] Simanovich | Smectites and isotropic phases in basalts of northern Timan[END_REF], and probably of glass of [START_REF] Emetz | Geology, mineralogy and origin of Zhyrychi native copper deposit (North-Western Ukraine)[END_REF]. Smectite, chlorite and analcime were identified in chlorophaeite accumulations. [START_REF] Mysiak | Unusual association of native metals in the platobasalts of Volyn and conditions of their formation[END_REF] reported native Cu and Ag occurring along with hydrothermal minerals, while native Fe, Sn, Al, chlorides of Sn and Fe, and a FeCrNiTi alloy in pneumatolytic veins cutting the hydrothermal alteration zones. Native Fe was reported also by [START_REF] Kvasnytsya | Native iron from Volyn Vendian volcanogenic rocks[END_REF]. [START_REF] Kuzmenkova | Petrography of magmatic rocks of the trap formation of Belarus[END_REF] investigated the felsic rocks (rhyodacites) and found large amounts of quartz, K-spar, albitized plagioclase, and chlorite+biotite pseudomorphs after pyroxene. The chemical composition of tuffs is also much more variable than flood basalts, demonstrating higher K and Na and lower Ca. Saponite, mixed-layer chlorite-smectite, hydromica, analcime, other zeolites, and hematite were identified as alteration minerals in the tuffs.

To summarize: the Volyn flood basalt mineral and chemical composition, and their alteration processes have been well recognized by numerous studies. This study aims at quantifying the alteration process, investigating its relationship to the conditions on the Ediacaran land, and investigating in more detail the iron-rich smectite, which is the main alteration product.

Materials

All investigated samples (locations in Figure 1) are core materials, except for the Tashki samples, collected from a quarry. Present depth, macroscopic characteristics, and location of a sample in the trap profile (lower, middle, and upper volcanogenic strata) are given if available in Inline Supplementary Tables 1 and2, listing separately flood basalts, clays from veins and cavities in basalts, felsic rocks, and tuffs, according to the macroscopic identification. Several chlorophaeite accumulations, studied earlier by Kuzmenkova et al. (2006) were also investigated, some of them only by FTIR because of small amount of the available material, along with the reference montmorillonite samples 4529-2 and 4517-1 from the paleosol, developed on the Volynian basalt [START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF]. In three cases (Mokrany-1, 2, and 3), both basalt sample and a chlorophaeite accumulation separated from this basalt sample were available. Additional paleosol samples were used in the organic geochemistry study.

Methods

The main techniques used in this study are: quantitative X-ray diffraction of the bulk rock (QXRD) using in-house Q-Min software (by Dr. M. Szczerba, Institute of Geological Sciences PAS), XRD of the clay fractions separated from the bulk rock, chemical analysis (major elements by X-Ray Fluorescence (XRF) and trace elements (including REE) by Laser Ablation Inductively Coupled Plasma Mass Spectrometry (LA-ICP-MS) on the same lithium borate fused disc), Mössbauer spectroscopy and iron isotope composition of the bulk rock and of the clay fractions, and the K-Ar dating of clay fractions. These methods are identical to those used by [START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF] in the study of paleosols developed on these basalts, and their descriptions are reproduced with some modifications in the Supplementary Materials, along with additional techniques applied in this study: oxygen isotope and FTIR measurements of the bulk rock and the clay fraction, organic geochemistry, and petrographic observations of the basalt alteration products.

Results

XRD of the bulk rock

Flood basalts and chlorophaeites

Quantitative XRD data for basalts (in wt%) are given in Inline Supplementary Table 1, and representative XRD patterns, illustrating the encountered variability of mineral composition, are presented in Figure 2. Most common major components are calcic plagioclase, pyroxene, magnetite, dioctahedral smectite, trioctahedral smectite and chlorite or mixed-layer chloritesmectite. Quantities of these major components vary broadly. Both di-and trioctahedral smectites are present in most samples, as evidenced by separate 02 and 06 peaks (Figure 2).

Common minor components are: quartz, potassium feldspar of orthoclase/sanidine/sanidine group (most often more abundant than quartz), and hematite. Orthoclase cannot be distinguished from sanidine in the bulk rock XRD, so they are reported as a group. Rare minor components are ilmenite, anatase, analcime, pseudobrookite, clinoptilolite, laumontite, calcite, and siderite. Olivine, reported in published petrographic studies, was not detected.

The mineral compositions of two available gabbro samples fall within the limits characteristic of basalts (Inline Supplementary Table 1). Three chlorophaeite accumulations were investigated by XRD along with the host basalts, but only qualitatively in glass capillaries, because of the small amount of sample available. Di-and trioctahedral smectites, quartz and traces of calcic plagioclase, calcite and dolomite were detected in these samples. The smectitic clays are identified as separate di-and trioctahedral phases based on well resolved 02 reflections, by analogy to the bulk basalt XRD patterns, where the 06 reflections could also be recorded (compare Figures 2 and3).

Felsic rocks

Three rhyodacite samples were available. Their mineral composition is totally different from basalts: quartz, orthoclase/sanidine/sanidine, dioctahedral clay (illite-smectite), trioctahedral mica, hematite, chlorite, and a trace of kaolinite in one sample (Inline Supplementary Table 1, Inline Supplementary Figure 1).

Tuffs

Only purevolcanic materials were considered (Inline Supplementary Table 2, Inline Supplementary Figure 1). The presence of minerals absent from pyroclastic materials: microcline and muscovite was regarded as evidence of non-volcanic contamination (tuffites) and such samples were not included in the analysis. Mineral composition of tuffs is extremely variable. Most common components are: quartz, orthoclase/sanidine, Ca-plagioclase, dioctahedral expandable clay (illite-smectite or smectite), aluminoceladonite (identified by 1Mtv polytype and d001 close to 9.88 Ǻ, characteristic of this mineral, and distinguishing it from celadonite of d001=9.97 Ǻ), and chlorite or mixed-layer chlorite-smectite. Less abundant are trioctahedral smectite, albite, anatase, hematite, analcime, minor and rare are pyroxene and amphibole. Locally, tuffs contain dolomite, Mn-Fe dolomite and calcite (Pinsk-26), or only calcite (Tashki).

Tuff-basalt alternation

The contrasting characteristics of flood basalts and tuffs, emerging from the entire data set, can be observed also in a single vertical profile (Kobryn), composed of three basalt flows interbedded with four tuff layers free of non-volcanic contamination (Inline Supplementary Table 2). Flood basalts contain trace of quartz and hematite, abundant Ca-plagioclase, pyroxene, and trioctahedral smectite, and occasional minor chlorite or chlorite-smectite. Clay abundances increase towards the top surfaces of the basalt flows. Tuffs contain abundant quartz, orthoclase/sanidine, Ca-plagioclase and hematite, chlorite plus minor chloritesmectite, dioctahedral illite-smectite, and aluminoceladonite. Top samples from two tuff layers contain kaolinite.

XRD identification of clay fractions

Mineral composition of <0.2 m clay fractions, including percent smectite (%S) in mixedlayer illite-smectite, measured from the diagnostic peak positions on the glycolated patterns, using the regressions based on data of Środoń et al. (2009), is presented in Inline Supplementary Table 3.

Flood basalts

Attempts to separate pure clay fractions were unsuccessful for basalts with >60% of primary minerals, despite long grinding in McCrone mill and sonification. The <0.2 m separates contain abundant plagioclase and exhibit poor orientation (Mokrany-1B in Inline Supplementary Figure 2), indicating intimate intergrowth of smectite with host plagioclase crystals. More altered samples allow clean separation of clay fractions (e.g. 4504-1A in Inline

The chemical data were obtained for selected bulk samples and clay fractions (Inline Supplementary Table 4). The bulk rock data are grouped according to lithology into five categories: flood basalts, rhyodacites, tuffs, chlorophaeite, and vein and cavity clays. When average values are compared, tuffs are intermediate between flood basalts and rhyodacites, except Mg is similar to that of basalt and tuffs have elevated LOI, B, Cr, Cs, Cu, Ni, and Pb.

The chlorophaeite sample (Mokrany-3A), when compared to its host basalt (Mokrany-3B), is depleted in most elements, including HREE and Fe, Na, Ca, Ba, Sr, Nb, Sc, Y, Zn, U, V, and Zr, but is strongly elevated in Si and moderately elevated in Mg, LOI, Cr, Cs, Cu, Ni, Rb, and LREE. The vein clays, when compared to average basalt, have strongly elevated Mg, LOI, and Co, elevated Fe, Ga and Ni, but most of the elements are depleted. REE are depleted particularly strongly and several times stronger in chlorite (Kob-23) than in the trioctahedral smectite sample (4600-1). The clay fraction of the latter is depleted even stronger. When clay fractions separated from basalts are compared to their host rocks, the common characteristics are the clay enrichment in Mg, LOI, Be, Co, Cr, Mo, and Ni, and depletion in B, K, Ca, Sr, Ba, Ti, Cu, Nb, and V. REE contents of the basalt clay fractions do not differ significantly from the bulk rock contents.

Mössbauer data

The Mössbauer data (Table 1) were collected for a range of bulk rock basalt, vein clay, and tuff samples, selected clay fractions <0.2 m, and two pyroxenes separated from the basalts.

The spectra of pure minerals separated from basalts were used as starting parameters for fitting polymineral spectra of rocks (Table 1, Figure 4) and include: trioctahedral smectite (Inline Supplementary Figure 5, 4600-1), dioctahedral smectite (4529-2, cf. [START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF], chlorite (Figure 4, Kob-23), and pyroxenes (4519-0 and 4603-1). Where applicable, the well known sextet of hematite and two sextets of magnetite were included in the fit.

Chlorite was found to contain only Fe 2+ , dioctahedral smectite only Fe 3+ , and trioctahedral smectite both Fe 2+ and Fe 3+ . Two pyroxenes contained Fe 2+ distributed over two possible crystallographic positions in different proportions, visible in the Mössbauer spectra as two well resolved quadrupole doublets. Finally, the rocks were analyzed assuming up to 7 different components with 9 subspectra (Table 1, Figure 4). In sample Mokrany-1B one additional component was detected with parameters characteristic of the amorphous wüstite, FeO.

The measured bulk rock percent Fe 2+ in the total Fe vary from 65 to 23% for flood basalts and from 49 to 11% for tuffs. The corresponding Fe 2+ values, calculated for the clay fraction from the bulk rock data, are lower than the bulk rock values for less altered flood basalts but higher than the bulk rock values for more altered basalts and tuffs. The % Fe 2+ values of clays, measured for the <0.2 m fractions and calculated from the bulk rock analysis are essentially identical, if no chlorite or chlorite-smectite is present (Mokrany-3B and 4504-1 in Table 1), indicating that the separation procedure did not affect the oxidation status of Fe in smectites. The % Fe 2+ values are lower in the bulk rock in other cases (4417-1 and 4600-1), which is explained by relative concentration of smectite with respect to chlorite in the <0.2 m fraction.

FTIR data

Identification of clay fractions

The mid-infrared (MIR) spectra of the <0.2 um fractions of XRD-identified end-member components (dioctahedral smectite from paleosol 4517-1, trioctahedral smectite from 4600-1, and chlorite from Kob-23) were measured to serve as references (Figure 5). Spectra appear similar in the vOH region, exhibiting two broad bands centered at 3624 cm -1 due to smectite or chlorite structural OH, and at 3434 cm -1 , mainly due to water. For chlorite, the band at 3624 cm -1 is characteristically broader than for the other samples and is typical of trioctahedral chlorite (clinochlore type). For the dioctahedral smectite, the 3624 cm -1 is characteristically narrower than for the other samples and is typical of montmorillonite (e.g. [START_REF] Madejová | IR Spectra of Clay Minerals[END_REF]. For the trioctahedral smectite (4600-1), a shoulder at about 3670 cm -1 is observed in addition to the 3624 cm -1 band. Such bands are found in griffithite [START_REF] Madejová | IR Spectra of Clay Minerals[END_REF], which was previously identified as a trioctahedral iron-rich saponite [START_REF] Komadel | Reduction of Fe(III) in griffithite[END_REF]. Note that for saponite depleted in iron, the OH is narrow and centered at 3678 cm - 1 . The spectra of the two other samples Mokrany-3B and 4504-1A are similar to 4600-1.

The 1200-400 cm -1 region was more discriminant for sample characterization. The dioctahedral smectite could be unambiguously identified as montmorillonite through the occurrence of δAl2OH and δAlMgOH specific vibration bands at 920 and 837 cm -1 , respectively (e.g. [START_REF] Madejová | IR Spectra of Clay Minerals[END_REF]. A shoulder at 878 cm -1 due to δAlFe 3+ OH indicated that the montmorillonite contains some octahedral trivalent iron. The spectra of the 4600-1 and 4504-1A samples corresponded clearly to griffithite with the characteristic bands at 1016, 678 and 452 cm -1 and shoulders at 740 and 520 cm -1 [START_REF] Madejová | IR Spectra of Clay Minerals[END_REF]. The two strong bands at 1016 and 452 cm -1 are respectively due to δSi-O and δSi-O, whereas the 678 cm -1 band is mainly due to δ(Fe 2+ ,Mg)3OH and indicates clearly the trioctahedral nature of the smectite. The Mokrany-3B spectrum revealed quartz in mixture with griffithite. The small shoulders at 876 and 912 cm -1 indicated the occurrence of octahedral aluminium and trivalent iron in dioctahedral configuration.

Comparison between clay component in chlorophaeite and the surrounding basalt

As seen in Figure 6, the clay components in chlorophaeite are the same as those in the surrounding basalt. Indeed, for Mokrany-3A and Mokrany-4A, the spectra revealed the same features as for Mokrany-3B described above. The clay component is then griffithite (see above). The slightly more pronounced dioctahedral features (shoulders at 878 and 912 cm -1 due to of δAl2OH and δAlFe 3+ OH) in Mokrany-4A indicated a higher content of octahedral aluminium and trivalent iron in dioctahedral configuration than in Mokrany-3A. For Mokrany-1A, the difference between the two spectra was the occurrence of feldspars in the surrounding basalt that hide the clay component features. However, in both samples griffithite could be identified and no dioctahedral features could be observed. The variability of chloropheite composition detected in MIR was also well expressed in NIR (near infrared), where the clay features could be observed without overlapping by quartz or feldspars.

Organic geochemistry data

All the rock samples investigated in this study contain organic matter (OM), but total organic carbon (TOC) content did not exceed 0.2%. The main constituents of OM in silylated total extracts (Figure 7A) were n-alkanes and fatty acids (FAs). The distribution of these biomarkers in the basalt sample 4600-3B differs significantly from other samples (Figure 7B): long-chain n-alkanes and n-alkanoic acids with the maxima between C23 and C29, and C22 and C30 carbon atoms, respectively, are present in relatively high concentrations. The values of short-chain to long-chain ratio (SCh/LCh) for n-alkanes and FAs for 4600-3B are 2.64 and 1.27, respectively, whereas they are 9.12 and 29.81 for 4600-3A. Moreover, OM in sample 4600-3B is characterized by the predominance of high molecular weight odd-carbonnumbered n-alkanes and even-carbon-numbered FAs, with the carbon preference index (CPI) values of 1.70 (CPI(25-31)) and 1.30 (CPItotal) for n-alkanes, and 7.43 (CPIf.a.(20-30)) for FAs (for CPI formulas see [START_REF] Rybicki | Molecular tracers preserved in Lower Jurassic "Blanowice brown coals" from southern Poland at the onset of coalification: Organic geochemical and petrological characteristics[END_REF]. In the clay sample 4600-3A such predominance is not evident (CPI(25-31)=1.01 and CPItotal=0.83).

Alteration petrography of chlorophaeite-bearing flood basalts

The primary components of basaltic rocks identified by the petrographic observations are pyroxenes (ortho-and clinopyroxene), Ca-Na plagioclases, Fe,Ti oxides (titanomagnetite and ilmenite) and a homogeneous orange-brown isotropic substance presenting all the characteristics of chlorophaeite, as identified in several basaltic fields [START_REF] Peacock | Chlorophaeite, sideromelane and palagonite from the Columbia River Plateau[END_REF]. Such isotropic material occurs in several petrographic settings: (1) areas of intersertal mesostasis through the body of the basaltic rocks (Figure 8A), (2) filling or lining of circular cavities (Figure 8B), (3) veinlets and associated infilling of dixitaxitic voids in the basaltic rock (Figure 8C), (4) pseudomorphs after orthopyroxene (Figure 8D).

In the least altered basalts (sample 4522-1), chlorophaeite is totally isotropic at a microscopic scale and it contains fine networks of minute cracks, which tend to be more developed when observed under vacuum (evidence of water content). In the cavities, chlorophaeite occurs as irregular films showing frequently collomorphous texture typical of precipitation of a gel material (Figure 8B). Chemical microanalysis of the isotropic material observed in thin sections confirms its chemical homogeneity (Table 2). The sum of oxide weight percent, which does not exceed 85%, indicates a high water content of this isotropic material. Silicon and iron ( oxidation state not determined from EDS) are dominant, with subordinate Mg, Al and Ca, and minor amounts of alkali elements (Na, K). Such a composition reasonably agrees with the formula given by [START_REF] Peacock | Chlorophaeite, sideromelane and palagonite from the Columbia River Plateau[END_REF] for chlorophaeite: (Fe,Al)2O3 . 2(Mg,Fe,Ca)O . 4SiO2 + 10H2O.

Observation of secondary mineral infillings indicates that chlorophaeite was deposited after the crystallization of quartz on the walls of vesicles and before the crystallization of chlorite, which cements the central parts of the vesicles. In the absence of phyllosilicates, minerals of the zeolite group are commonly observed as the late mineral cement of the residual open spaces of the basaltic rocks.

Smectitic clayswere observed systematically as alteration products of the chlorophaeite bearing basaltic rocks. In the less altered basalts (sample 4522-1), the occurrences of smectitic clays are limited to minute spherules replacing chlorophaeite close to the contacts with the primary minerals of basalt (Figure 8E). Detailed observations with SEM indicate that such spherules are alteration features of chlorophaeite with a typical concentric structure. Smectitic clays, rimmed by a thin iron rich layer, were observed in the inner part of the spherule. The transition to unaltered chlorophaeite is marked by an external alteration halo, due to increasing microporosity and/or hydration of the initial isotropic material (Figure 8F). Chemical microanalysis of smectitic clays formed in the spherules indicates a close chemical similarity with that of the chlorophaeite starting material, except for lower iron and higher Mg and Al in the smectite (Table 2). The thin iron rich layer of the spherules likely accommodates the excess iron released during the alteration process of chlorophaeite.

In moderately altered samples (sample 4519-0), most of chlorophaeite lost its isotropic optical characteristics because of replacement by spherulitic and ribbon-like smectitic clays. Furthermore, primary or the previously formed hydrothermal minerals present features of partial dissolution. This is particularly the case for the iron rich phyllosilicates such as Fechlorite (Figure 8G). It should be noted that quartz presents also some evidence of dissolution at the contact with the smectitic clays (Figure 8H). In strongly argillized samples (4603-1, and 4603-4), chlorophaeite is totally replaced by smectitic clays and the basaltic primary minerals are partially replaced by smectitic clays and minerals of the zeolite group.

O and Fe isotope data

Oxygen isotope data were collected for a range of bulk basalts showing variable alteration intensity, for a few clay vein and tuff samples, and for selected clay fractions <0.2 m (Table 1). The least altered basaltic rocks have oxygen isotope values of 5.6 to 8. A few of these <0.2 m clay fractions were also analyzed for iron isotope compositions (Table 1). In moderately altered basalts, the measured  56 Fe values are positive (0.028 for Mokrany-3B), but become negative with increasing alteration (down to -0.574 for Kob-23).

K-Ar dating of aluminoceladonite and illite-smectite

Selected 2-0.2 and <0.2 m fractions of tuffs containing aluminoceladonite and illite-smectite were dated, at least twice, in order to establish the age of potassic alteration (Table 3). The K2O contents vary from 1.785 % for the sample Tashki-10, with minor amounts of aluminoceladonite and illite-smectite, to 6.71% for Kob-9B, dominated by aluminoceladonite (comp. Inline Supplementary Figure 4). The measured ages range from 447.5±2.5 to 296.0±5 Ma.

Interpretation

Alteration of flood basalts

Mineral changes

No single sample was identified by XRD to be composed exclusively of primary basalt minerals: Ca-plagioclase, pyroxene, magnetite, ilmenite, and pseudobrookite. All samples contain hydrated components (mostly clays), but in variable quantities. If the sum of primary basalt minerals is accepted as an inverse index of basalt alteration, and if the degree of alteration is interpreted as a chronosequence (stronger alteration considered as a subsequent step of the process), a clear pattern of basalt alteration emerges, which can be described by five steps (Inline Supplementary Table 1), leading from almost fresh basalt to completely altered rock, though still preserving the basalt texture.

In the investigated samples, the earliest alteration products are quartz with trioctahedral smectite, or with a mixture of di-and trioctahedral smectite, all together < 10 wt%. This is also the mineral composition identified in chlorophaeite accumulations, found in basalts with more advanced alteration (compare Mokrany samples in Inline Supplementary Table 1). Based on this identity and on the microscope observations, we assume that the early quartz did not crystallize from melt but it is a product of basalt alteration (mostly of volcanic glass), deposited in crystal interstices or in larger chlorophaeite accumulations. Thus, step I can be identified tentatively as predominantly the alteration of volcanic glass. [START_REF] Emetz | Geology, mineralogy and origin of Zhyrychi native copper deposit (North-Western Ukraine)[END_REF] also identified quartz as an alteration mineral based on petrographic observations. In step II (90-80 % primary minerals), orthoclase/sanidine and hematite appear and the amount of smectites increases compared to step I. In step III (80-60 % primary minerals), chlorite appears, hematite, quartz and smectite abundances increase, pyroxene decreases, and ilmenite is absent. In step IV (60-25% primary minerals), anatase, analcime, zeolites and calcite appear, smectite, quartz and orthoclase/sanidine increase further, plagioclase starts decreasing, and pyroxene and pseudobrookite disappear. In step V (<25% primary minerals), plagioclase and magnetite disappear, and dioctahedral smectite plus chlorite prevail over trioctahedral smectite. The mineral composition of green clay veins and cavities in basalts is similar to the most altered basalts of step V. Ubiquitous hematite indicates an oxidizing alteration environment. In general, the lower basalts appear more altered than the upper basalts. Petrographic study confirmed the crystallization of smectite both before and after chlorite.

Chemical changes

The mineral alteration of basalt is paralleled by characteristic chemical changes, which are surprisingly limited (Inline Supplementary Table 4) considering the extent of alteration observed in some samples. The most pronounced is the decrease of CaO from 10% almost to 0% (except of samples with abundant calcite). The progressing dissolution of Ca-plagioclase results in the removal of Ca from the system (Inline Supplementary Figure 6). Sr decreases along with Ca. In steps IV and V, when plagioclase decreases strongly and analcime and other zeolites appear, Na seems to undergo redistribution: it is depleted from rocks without zeolites and enriched in remaining rocks.

The most pronounced gain during basalt alteration is the sum of OH and H2O in clays, represented by the loss on ignition (LOI), which increases from almost 0% to 6.5% in samples free of calcite. The correlation of sum of clay with LOI is very high and extrapolates to 0 for unaltered basalt and to 9% for completely altered basalt. The trend line runs between values characterizing smectitic and chloritic green clay veins in basalt (Figure 9), in agreement with the intermediate, smectite+chlorite composition of clays in strongly altered basalt. LOI can serve as a simple measure of the degree of basalt alteration.

In the course of alteration, basalts also gain MgO, which increases about two times: from 5 to 10% and is positively correlated with the sum of clays (Figure 9), except of the samples with carbonates (not shown in the figure). Boron increases even more, in particular in steps IV and V, and good positive correlation indicates that B is located in the dioctahedral clay (Inline Supplementary Figure 7). Some increase in K, Cs, and Rb is also observed, and these trends can be best explained by combined contributions from K-spar and dioctahedral clay (Inline Supplementary Figure 8), which is partially illitized during much younger Paleozoic events (see below), particularly in highly altered basalts (Inline Supplementary Table 3).

The remaining elements vary, but no trends related to the degree of alteration can be observed (Inline Supplementary Table 4). The mass of basalt during alteration essentially does not change, because the additions and subtractions of elements quantitatively compensate. This mass stability is best expressed by very stable SiO2 and Al2O3 contents. The collected data do not indicate measurable chemical changes between the lower and upper flood basalts. REE patterns of most basalt samples are very similar and characteristic of the continental flood basalts [START_REF] Farmer | Continental Basaltic Rocks[END_REF]: enriched in LREE and without an Eu anomaly (Figure 10). REE do not differentiate the lower from the upper basalts. LREE seem slightly depleted at the final alteration step, particularly in the rock with high chlorite content (Kob-22), but otherwise basalt alteration does not affect the REE pattern, despite sometimes total destruction of major primary minerals. Inline Supplementary Figure 9 indicates that LREE in basalts seem mostly contained in phosphates, of too low content to be detectable mineralogically (Inline Supplementary Table 1).

In chlorophaeite, LREE are strongly concentrated, compared to regular basalts, while HREE are depleted, and a strong negative Eu anomaly is present (Mokrany-3A in Figure 10 and Inline Supplementary Figure 9). This composition confirms the origin of chlorophaeite from the volcanic glass, enriched in incompatible LREE, while depleted in Eu contained in plagioclase [START_REF] Weill | Europium anomaly in plagioclase feldspar: experimental results and semiquantitative model[END_REF]) and depleted in HREE, which form isomorphic substitutions in pyroxene [START_REF] Rollinson | Using Geochemical Data: Evaluation, Presentation, Interpretation[END_REF][START_REF] Kuzmenkova | Geochemistry of the trappean formation of the Vendian of Belarus: PhD thesis, Republican Unitary Enterprise[END_REF]. The basalt sample hosting chlorophaeite concentrations (Mokrany 3B in Figure 10 and Inline Supplementary Figure 9) displays analogous characteristics. Both have much higher LREE/P2O5 ratio compared to regular basalts (Inline Supplementary Figure 9). These characteristics are similar to rhyodacite samples, which also differ from basalts by elevated LREE, negative Eu anomaly (Figure 10), and much higher LREE/P2O5 ratio (Inline Supplementary Figure 9).

Oxygen isotopes and Fe 2+ /Fe 3+ ratio

The progress of basalt alteration can be traced also by oxygen isotope compositions of the bulk rock samples (Table 1, Figure 11). They evolve linearly and extrapolate to a value close to 5‰ for the unaltered protolith, consistent with the oxygen isotope composition of unaltered continental flood basalts that show little to no contamination by continental crustal material (e.g. [START_REF] Harmon | Oxygen isotope heterogeneity of the mantle deduced from global 18 O systematics of basalts from different geotectonic settings[END_REF][START_REF] Eiler | Oxygen isotope variations of basaltic lavas and upper mantle rocks[END_REF], and to the value of about 22‰ for complete alteration, similar to the isotopic compositions of smectite from the weathering zone (4529-2 and 4517-1) developed on basalts. A similar value of 23‰ characterizes the measured chlorophaeite sample, composed exclusively of secondary minerals (Mokrany-3A) and the clay fraction separated from the host basalt (Mokrany-3B<0.2). In Figure 12, the equilibrium relationships between oxygen isotope composition of smectite, of water, and temperature, based on the fractionation equation of [START_REF] Sheppard | Stable isotope geochemistry of clay minerals[END_REF] are depicted. The smectiterich paleosol samples are in equilibrium with waters of about -5 ‰ at ambient temperatures.

Using this isotope value of meteoric water, temperatures of dioctahedral smectite alteration in the basalts is estimated at less than 40°C. On the other extreme, if only magmatic water with a heavy oxygen isotope composition of +5‰ is assumed as the clay-forming fluid, then temperatures of about 100 to 130°C are indicated for the dioctahedral smectite formation.

Another parameter correlating well with the applied index of basalt alteration is % Fe 2+ of the bulk rock measured by Mössbauer spectroscopy (Table 1, Figure 13). The only measured chlorophaeite sample also fits this trend (Mokrany-3A). The linear trend extrapolates to ca. 67% for pure basalts and falls to 30% at ca. 40% of primary minerals, where those containing Fe almost disappear. Fe 2+ was lost due to oxidation, and not removal under anoxic conditions, as Fe content of the rock remained stable (Inline Supplementary Table 4). This trend implies that basalt alteration from the very beginning involved water carrying oxygen, in agreement with the increasing content of hematite.

The most altered samples (Kob-15 and 4417-1) and green clays from veins and cavities (4600-1 and Kob-23) fall off the general trend, as their % Fe 2+ is much higher (Table 1, Figure 13). This implies higher % Fe 2+ of clays at this stage of alteration, which agrees well with the Mössbauer estimates of % Fe 2+ in clay fractions of rocks (Figure 13), the Mössbauer data for tuffs, and the abundance of chlorite in these rocks. These data lead to the conclusion that at the most advanced stage of alteration, the conditions became more reducing.

Alteration of organic compounds

The primary distribution of particular organic compounds depends on the source of OM and can be modified by a wide range of post-depositional processes. Algae and bacteria generate short-chain n-alkanes with preponderance of C17 and C19 carbon atoms in the molecule, while terrestrial higher plants (especially their leaf epicuticular waxes) are characterized by the occurrence of long-chain n-alkanes and n-fatty acids (n-FAs). Our studies evidenced a clear predominance of long-chain n-alkanes and n-FAs (excluding C16 and C18) with majority of an odd and even number carbon atoms in the molecules, respectively, in most of the Ediacaran sedimentary rocks from the EEC [START_REF] Goryl | Pattern of n-alkanes distribution characteristic for terrestrial organic matter from the Ediacaran sedimentary rocks of the East European Platform[END_REF]. [START_REF] Bobrovskiy | Molecular fossils from organically preserved Ediacara biota reveal cyanobacterial origin for Beltanelliformis[END_REF] suggested that a possible source of long-chain n-alkanes in the Ediacaran rocks are spherical cyanobacterial colonies. The level of thermal maturity of the Ediacaran OM from the EEC is unusually low (the estimated equivalent of vitrinite reflectance did not exceed 0.57% Rr) which promoted preservation of the primary character of distribution of organic compounds [START_REF] Pehr | Ediacara biota flourished in oligotrophic and bacterially dominated marine environments across Baltica[END_REF][START_REF] Goryl | Exceptional preservation of hopanoid and steroid biomarkers in Ediacaran sedimentary rocks of the East European Craton[END_REF].

We believe that OM from the basalt sample 4600-3B is the most primary, nonoxidized or only slightly oxidized remnant of the Ediacaran cyanobacteria (and other microorganisms?) activity in the basalt hydrothermal system. Decay of long-chain n-alkanes and n-FAs and decreasing value of CPI in all other samples indicate more intense oxidation of OM in paleosols and in clays from cavities. Similar effect of OM alteration was presented by [START_REF] Elie | Natural and laboratory oxidation of low-organic-carbon-content sediments: comparison of chemical changes in hydrocarbons[END_REF][START_REF] Elie | Effects of water-cement solutions on the composition of organic compounds leached from oxidized Callovo-Oxfordian argillaceous sediment[END_REF]: gradual degradation of long-chain n-alkanes and n-alkanoic acids during artificial oxidation of argillaceous rocks. Also [START_REF] Marynowski | Organic geochemical evidences of early diagenetic oxidation of the terrestrial organic matter during the Triassic arid and semi-arid climatic conditions[END_REF] presented significant influence of early diagenetic oxidation on n-alkanes and other organic compounds distribution on an example of green and red Triassic clays. It seems that in the basaltic rock, primary OM was better protected against oxidation by the Ediacaran hydrothermal processes and further Caledonian and/or Variscan alterations. Such cyanobacteria activity offers a feasible explanation of more reducing conditions observed at the final stages of basalt alteration.

Characteristics of the basaltic smectite

For three available pure smectitic samples (<0.2 m fractions) with variable ratios of di-and trioctahedral components, the chemical data (Inline Supplementary Table 4), including Fe 2+ /Fe 3+ ratio from the Mössbauer spectroscopy (Table 1), were cast into structural formulae (Table 4). The amount of potassium is very low, consistent with the XRD characteristics (Inline Supplementary Table 3, Inline Supplementary Figure 2).

The chemical compositions of the three smectitic samples evolve in a regular manner, consistent with the XRD evaluation of di/tri-smectite ratio (Table 4). The calculated numbers of octahedral cations are between 2 and 3 per half unit cell and they can be used as a measure of di/tri-smectite ratio. Plots of other data with respect to these numbers are linear (R 2 =0.92-0.99) and allow extrapolating compositions of the end member components (Table 5). The dioctahedral smectite end member is a perfect montmorillonite without tetrahedral substitution, high Fe and Mg contents, and no divalent iron, as confirmed by Mössbauer spectra of the dioctahedral smectite component (Table 1) and dioctahedral smectite identified from basalt weathering zones by [START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF]. Its oxygen isotope characteristics and its positive d 56 Fe are also similar to the montmorillonites from paleosols, which are however much more rich in Al (op. cit.). The dioctahedral smectite has been partially illitized during the Paleozoic thermal events (see below). Almost all REE is contained in the dioctahedral clay, which is consistent with their trivalent character (the dioctahedral sheet accepts mostly trivalent cations). Exceptionally low REE content of chlorite Kob-23 (Figure 10), which contains 95% of divalent iron (Table 1) is consistent with this interpretation of REE segregation. Thus dioctahedral clays are partially responsible, along with phosphates, for the REE budget of the investigated rocks. This conclusion is confirmed by P2O5 depletion in clay fractions compared to the bulk rock (Inline Supplementary Table 4), parallel to the LREE enrichment in the clay fraction with dominant dioctahedral component (Mokrany-3B). REE seem to be incorporated in the dioctahedral clay structure and not adsorbed like in weathering clays [START_REF] Bau | Rare-earth element mobility during hydrothermal and metamorphic fluid-rock interaction and the significance of the oxidation state of europium[END_REF], as the clay fractions were analyzed after Na-exchange. The location of REE in clay structure indicates that they became available at the time of clay crystallization and not later. Otherwise they would occupy the exchange sites.

The trioctahedral smectite end member is almost ideal iron-rich saponite, with 0.84 Al IV , no Al VI , very high Mg, and Fe not much higher than in the dioctahedral clay, but ferrous in 42%. Its  18 O value is much lower than in di-smectite and its  56 Fe is negative, like in chloritic trioctahedral clays (Table 1).

XRD, chemistry and FTIR data are consistent and detect variable di/tri ratio of the investigated basaltic smectites. Based on the XRD evidence (separate 02 and 06 reflections) and very different REE contents and oxygen and iron isotope ratios, the smectitic clay component of Volyn basalts is interpreted here not as a mineral but as a mixture of iron-rich montmorillonite and iron-rich saponite. This mineral mixture occurs both in basalts and in the chlorophaeite accumulations (Figure 3 and Inline Supplementary Figure 3). A similar material was identified by [START_REF] Faust | Thermal analysis and X-ray studies of griffithite[END_REF], [START_REF] Ross | Review of the relationships in the montmorillonite group of clay minerals[END_REF], and [START_REF] Komadel | Reduction of Fe(III) in griffithite[END_REF] as iron ditrioctahedral saponite, and named griffithite. Formulae of griffithite calculated from [START_REF] Ross | Review of the relationships in the montmorillonite group of clay minerals[END_REF] and [START_REF] Komadel | Reduction of Fe(III) in griffithite[END_REF] data (Table 4) and the XRD data, obtained for [START_REF] Komadel | Reduction of Fe(III) in griffithite[END_REF] sample (Inline Supplementary Figure 3), suggest that it is also a mineral mixture, close to our sample 4600-1. The reported griffithite formulae differ strongly in the degree of iron oxidation, which may reflect sample heterogeneity and/or the measurement technique.

Both samples come from the same original griffithite locality at Griffith Park in California, but may represent slightly different material. Other samples from this area are end-member Fe-saponites with high Fe 3+ /Fe 2+ ratios [START_REF] Treiman | Ferrian saponite from the Santa Monica Mountains (California, USA, Earth): Characterization as an analog for clay minerals on Mars with application to Yellowknife Bay in Gale Crater[END_REF].

Very similar materials, named oxysmectites, were described by [START_REF] Dainyak | New variety of trioctahedral smectite from effusive basalts of Tunguskaya Sineclise[END_REF] and [START_REF] Simanovich | Trap magmatism and mobilization of hydrocarbon fluids (Western Siberia)[END_REF] from continental flood basalts of Siberia (Table 4) and from an oceanic basalt sample. The continental sample has di/tri ratio close to our 4504-1 but it has more Fe than Mg. A pure trioctahedral analogue of our sample 4600-1, but with Fe dominant over Mg (Table 4), was identified in a Siberian hydrothermal deposit and called ferrosaponite (Chukanov et al., 2004), the name accepted by IMA [START_REF] Jambor | New mineral names[END_REF]. Numerous occurrences of similar materials were reviewed by [START_REF] Sarkar | Alteration of volcanic glass to well-crystallized ferrosaponite in the vesicles of the Deccan trap basalts at Bhuleshwar Ghat section, Pune district, Maharashtra[END_REF], who identified ferrosaponite in continental basalts.

The relative enrichment of montmorillonite in the heavy Fe isotope compared to Fesaponite is consistent with the general rule that the oxidized form of Fe is enriched in heavy isotopes (Teutsch et al., 2009 and literature cited). Negative  56 Fe values of chloritic samples (Table 1), which contain only Fe 2+ , support this interpretation. Comparable or even higher fractionations of iron have been observed during oceanic basalt alteration and interpreted as abiotic [START_REF] Rouxel | Iron isotope fractionation during oceanic crust alteration[END_REF], though a biotic involvement cannot be excluded in our case, as the dissimilatory iron reduction is regarded as the most efficient mechanism for generating isotopically light Fe 2+ (Wu et al., 2010), and the presence of unaltered organic matter in basalt has been documented in previous sections.

Major differences in  18 O values between di-and trioctahedral smectite could be related to formation from the same fluid at different temperatures with trioctahedral smectite forming at ca. 50 o C higher temperatures if fractionation factors are similar (Figure 12). However, theoretical considerations [START_REF] Savin | Isotopic studies of phyllosilicates[END_REF][START_REF] Kohn | Oxygen isotope geochemistry of the amphiboles: Isotope effects of cation substitutions in minerals[END_REF] and empirical data [START_REF] Craw | Stable isotopic signatures of authigenic minerals in a Holocene ophiolitic debris flow, Southland, New Zealand[END_REF] suggest that oxygen isotope fractionation between smectite and water is about 5 ‰ smaller for Mg-rich smectite than for Al-rich smectite. Thus, di-and trioctahdral smectites may have formed at similar temperatures and from the same fluid. On the other hand, the temperature difference may have been higher, if trioctahedral smectites formed from a fluid with magmatic component (isotopically heavier).

Petrographic observations document the appearance of early smectite as the alteration product of isotropic material, which was described earlier in many publications under the name "chlorophaeite". It was deposited during hydrothermal processes, cementing successively the residual open spaces of the basaltic rocks. Under the microscope chlorophaeite appears isotropic only in the least altered rocks, and during progressing alteration it gradually evolves into birefringent smectite. This alteration is almost isochemical, but some iron is lost (Table 2). In XRD, smectite from the chlorophaeite accumulation in least altered basalt differs from others by more trioctahedral character and much smaller crystal thickness, evidenced by more diffuse 001 reflection (Figure 3). NIR spectra indicate that it is richer in Fe, without detectable Al contribution. Combining these observations, we suggest that the isotropic substance is a trioctahedral smectite close to ferrosaponite, with extremely fine crystal size, which results in the isotropic characteristics. It may have been deposited originally as an amorphous material, but no clear evidence of such material was found in the investigated samples. Smectite crystallizing from an amorphous material is known to form originally extremely thin, even monolayer crystals (Derkowski et al., 2015 and references therein). In the course of further alteration, ferrosaponite alters into a mixture of Femontmorillonite and Fe-saponite with well-developed multi-layer crystallites. At early stages, the quantities of both smectites grow in parallel in the course of progressing alteration, but trioctahedral smectites grow at twice higher rate (Inline Supplementary Figure 10). Latter, the content of dioctahedral smectites continue to grow steadily, while the content of trioctahedral smectites levels off and starts decreasing, when the hematite content starts increasing. We interpret these relationships as indicative of the redox control over the process: as long as the conditions are relatively reducing, crystallization of trioctahedral smectite is favored, both by abundant Mg 2+ and Fe 2+ : the major constituents of the trioctahedral sheet. When Fe 3+ becomes dominant, the crystallization of dioctahedral smectite is favored (trivalent ions dominant in the dioctahedral sheet), and the decrease of trioctahedral smectite indicates that it is altered into chlorite and/or dioctahedral smectite at these advanced stages of basalt alteration.

Alteration of tuffs

Mineral changes

Mineralogically, tuffs (volcano-clastic rocks) differ strongly from flood basalts by their much broader range of quartz content, which, if elevated above basalt values, is always higher than K-feldspar content. The opposite relation, characteristic of flood basalts, is common in low quartz tuffs. If % quartz is plotted against % Kspar for all tuff samples (Inline Supplementary Figure 11) it becomes clear that the tuffs have a broad range of intermediate compositions, extending from basalts to the felsic rocks. Quartz content was used then as a proxy for this primary tuff mineral variability (Inline Supplementary Table 2), and tuffs were divided into samples close to basaltic composition (quartz <14%) and more felsic samples. In general, tuffs with >20% quartz are competent rocks and with lower content of quartz are friable.

The basaltic tuffs resemble by mineral composition the most altered flood basalts: some plagioclase still present in some samples, pyroxene and other primary minerals are absent, anatase is present, hematite, di-and trioctahedral smectite plus chlorite are abundant, and analcime is occasionally present. The apparent differences are the presence of albite and aluminoceladonite.

The felsic tuffs seem less altered because their calcic plagioclase content is much higher. They contain less anatase, hematite, trioctahedral smectite and chlorite. Only dioctahedral clay is as abundant as in basaltic tuffs, but more illitized. They also contain aluminoceladonite, but no albite or analcime. Kaolinite detected in the top samples of two Kobryn tuffs are indicative of subaerial weathering and soil development before deposition of subsequent lava flows.

Chemical changes

The arrangement of tuff samples by % quartz reveals very clear trends in their chemical composition (Inline Supplementary Table 4). Towards the basaltic end, Si and K contents decrease to a level typical of basalt, while Fe, Al, Mn, Mg, Ti and LOI increase. Mg and LOI increase to the level characteristic for the most altered basalts, which is consistent with the mineralogical evidence. Among trace elements REE, Be, Ba, Cs, Nb, Rb, Sr, Ta, Th, U, Y, and Zr decrease, while Co, Cr, Cu, Ni, Sc, V and Zn increase. No measurable control of the trap stratigraphy (lower vs. upper basalts) over chemistry can be observed.

In accordance with these trends, average contents of most elements for tuffs are intermediate between values for flood basalts and rhyodacites, indicating that, like flood basalts, also tuffs were altered without a major change of mass. For a few elements, the averages for tuffs are higher than the end member values: B, Mg, LOI, Cr, Cs, Cu, Ni, and Pb. Such elevated values imply the import of these elements during alteration, similar to the detected in basalts (B, Mg, LOI, Cs) but including metals. Also in basalts metals are mobile, as recorded by the metals enrichment in chlorophaeite and clay fractions, as compared to their bulk basalt samples (Inline Supplementary Table 4). This conclusion is consistent with the observation of native copper and silver occurrences in association with chlorophaeite (Kuzmenkova et al., 2006a).

Elements commonly most immobile in rocks: Hf, Nb, Pr, and Zr, used as the mass balance references in weathering studies (e.g. [START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF] are much less abundant in basalts than in felsic rocks. They differentiate the studied rocks similarly to quartz content (compare Inline Supplementary Figure 11and 12), confirming the classification of tuffs based on mineralogy into more basaltic and more felsic. In general, also this approach does not differentiate lower from upper volcanics. Locally, subsequent levels of flood basalts and tuffs can be differentiated, as illustrated in Inline Supplementary Figure 13for Kobryn profile, where three levels of tuffs and two levels of flood basalts have been sampled over ca. 150 m of depth. The tuff composition is clearly more felsic. On the other hand, four tuff samples from Tashki plot in Inline Supplementary Figure 9 outside the basalt field, indicating an even more basic composition than the available basalt samples. They represent bottom part of the volcanic sequence, where picrites and picrobasalts were identified [START_REF] Kuzmenkova | Petrology and correlation of trap formations of the Vendian in the adjacent areas of Belarus and Ukraine[END_REF][START_REF] Shumlyanskyy | The U-Pb zircon and baddeleyite ages of the Neoproterozoic Volyn Large Igneous Province: implication for the age of the magmatism and the nature of a crustal contaminant[END_REF].

The tuffs vary in color in most profiles, alternating from green to brown and even red.

Green color is more common among more felsic tuffs. When compared with other tuffs of similar quartz content, green tuffs differ from brown tuffs only by less hematite and more U, which is consistent with more reducing local environment (evacuation of Fe, enrichment in U) but does not allow us to conclude whether this is a hydrothermal alteration feature or the result of a younger process. The lack of systematic difference in clay composition favors the latter explanation.

Oxygen isotopes and Fe 2+ /Fe 3+ ratio

Both sets of data collected for tuffs of basaltic composition do not follow trends characteristic of flood basalts and resemble data obtained for clay veins in basalts: d 18 O values are lower ( Figure 11) and % Fe 2+ values are higher than extrapolations of basaltic trends to the zone of complete alteration (comp. Table 1 with Figure 13). These relations can be explained by higher abundance of chlorite in the analyzed tuff samples compared to the flood basalts (Supplementary Table 1 and2), based on the % Fe 2+ and the oxygen isotope data for Kob-23 chlorite (Table 1).

K-Ar dating of clay fractions

Dates of replicate analyses and dates of two fractions all are within the error range for a given sample, thus they can be regarded as representing unique illitization episodes (Środoń et al., 2002). In the south (Tashki), the Carboniferous date of ca. 300 Ma was obtained, while in the north, dates range from 417 to 447 Ma. A very similar geographical distribution of ages was detected by dating aluminoceladonite and illite-smectite in paleosols developed on basalts [START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF]. Clearly, the potassic alteration episode affected tuffs and paleosols developed on flood basalts long after the formation of these rocks. Basalts were also affected, though to a lesser extent: aluminoceladonite was not formed, but smectite was slightly illitized (Supplementary Table 3).

Discussion and conclusions

Flood basalts and tuffs of the Volyn-Brest province were affected to a very different extent by two subsequent alteration episodes, the first pre-dating and the second post-dating the Ediacaran weathering of top surface.

The first alteration episode involved sequential destruction of volcanic glass and primary basaltic minerals and crystallization of secondary phases: tri-and dioctahedral smectite, chlorite and/or mixed-layer chlorite-smectite, hematite, quartz, anatase, and zeolites.

Basalts which underwent Ediacaran weathering studied by [START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF] contained such secondary mineral assemblage.

The basaltic smectites are of unique composition: montmorillonite with high Fe 3+ and saponite also high in Fe, but with significant Fe 2+ content. Similar clays have been described earlier from various basalt provinces under the names "chlorophaeite", "griffithite" or "oxysmectite", and "ferrosaponite" if Fe>Mg.

The term "chlorophaeite" is used also in macroscopic descriptions to name accumulations of dark glassy matter in basalts. Such chlorophaeite accumulations do not contain primary basaltic minerals, but mainly secondary minerals characteristic of the first step of basalt alteration (quartz, tri-and dioctahedral smectite), also if chloreophaeite accumulations are contained in basalts of higher alteration step (Supplementary Table 1). In particular, they do not contain hematite, they are strongly enriched in Si and enriched in Mg and metals compared to the host basalt. These observations, along with the REE pattern, suggest that chlorophaeite accumulations are products of alteration of the amorphous material (solidified leftover melt, i.e. volcanic glass), in agreement with conclusions of many earlier researchers (Kuzmenkova et al., 2006b and the references cited), and that the alteration took place at the earliest stage, in non-oxidizing environment, but it is related to the same solutions that altered flood basalts and tuffs.

The subaerial location, the evolution of bulk rock Fe 2+ /Fe 3+ measured by Mössbauer spectroscopy, the evolution of bulk rock oxygen isotope composition,the presence of hematite, and the lack of depletion in Fe and Mn indicate that the basalt alteration involved oxygenated meteoric waters, probably heated and put in motion by the cooling basalt flows.

Such heat-driven convection of meteoric water is a very effective mechanism of alteration of cooling volcanic bodies (e.g. [START_REF] Eberl | Sericite from the Silverton caldera, Colorado: Correlation among structure, composition, origin, and particle thickness[END_REF]. A similar mechanism of alteration of Volyn flood basalts was proposed by [START_REF] Melnychuk | Hydrothermal mineralogical zonality and metamorphism in copperbearing Vendian flood basalts of the Volyno-Podolian plate[END_REF], who envisioned the Ediacaran dolerite sills intruding basalts and the underlying sandstones as the source of heat. The heat-driven convection mechanism explains well the observed major enrichment in Mg, brought by upflow from deeper parts of the basalt body, where olivine and pyroxene are early subjects to alteration. The alteration sequence is analogous to these reported from other hydrothermal subaerial, hydrothermal submarine, and burial metamorphic alterations of basalts, including the difference in Si/(Si+Al) and Fe/(Mg+Fe) ratios between the smectitic and chloritic clays (Table 4). The alteration conditions were more oxidizing, which is manifested by the presence of hematite from early stages of the alteration process. A rare feature for the basalt alteration series is the abundant development of dioctahedral smectite, which continued later during the weathering stage, as documented by [START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF]. The alteration proceeded from outside into the basalt layers, as observed also earlier [START_REF] Juskowiakowa | Upper Precambrian sediments[END_REF], and in extreme cases it was complete: all basaltic minerals were altered but the basalt texture was preserved.

Our data indicate that microbial organisms participated in the hydrothermal alteration process, which is consistent with earlier evidence of the microbial life on the Ediacaran land in this area [START_REF] Kremer | Cyanobacterial-algal crusts from Late Ediacaran paleosols of the East European Craton[END_REF][START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF].

Compared to flood basalts, the alteration of tuffs of basaltic composition is much more advanced, which seems related to their high porosity and permeability. Mineralogically, they differ by much more abundant chloritization, and albitization not noted in basalts. The alteration of rhyodacites and tuffs of felsic composition seems less advanced than basaltic tuffs but more advanced than of the neighboring flood basalts, as best illustrated by the Kobryn profile (Supplementary Table 2). The abundance of felsic tuffs in the entire profile of the volcanic province indicates that felsic volcanism was active during much longer period than indicated by the known felsic lavas of the middle volcanogenic series (e.g. [START_REF] Kuzmenkova | Petrology and correlation of trap formations of the Vendian in the adjacent areas of Belarus and Ukraine[END_REF].

The detected mineral assemblage: from smectite to chlorite with zeolites, but without major albitization, epidote, actinolite, or amphibole puts the upper temperature limit of postmagmatic, pre-weathering alteration at ca. 220 o C and lower at 50 o C, based on the calibration in the Icelandic active hydrothermal systems (Franzson et al., 2008: cf. Introduction) and many other smectite-to-chlorite transformation series reviewed by [START_REF] Robinson | Reaction pathways and reaction progress for the smectite-to-chlorite transformation: evidence from hydrothermally altered metabasites[END_REF]. The evidence of such maximum temperature conditions comes from most tuffs and the most altered flood basalts. Low  18 O of chlorite (10.1‰ for Kob-23 sample) is consistent with a temperature of 190°C if a magmatic fluid with ca. 6‰ is assumed [START_REF] Cole | Oxygen isotope fractionation between chlorite and water from 170-350°C: A preliminary assessment based on partial exchange and fluid/rock experiments[END_REF]. In this scenario, chlorite is considered as an indicator of a higher temperature overprint, which altered older iron saponite, formed by the interaction of basalt with the heated meteoric waters. Locally, the temperatures of Ediacaran alteration may have been even higher, as albite, prehnite, and pumpellyite were reported [START_REF] Emetz | Geology, mineralogy and origin of Zhyrychi native copper deposit (North-Western Ukraine)[END_REF][START_REF] Melnychuk | Hydrothermal mineralogical zonality and metamorphism in copperbearing Vendian flood basalts of the Volyno-Podolian plate[END_REF], and epidote was identified in the olivine pseudomorphs in samples from west Belarus (Skveriki well: [START_REF] Kuzmenkova | Geochemistry of the trappean formation of the Vendian of Belarus: PhD thesis, Republican Unitary Enterprise[END_REF]. On the other hand, most flood basalts were altered by much cooler fluids, as evidenced by the lack of chlorite and by the preservation of fresh organic material.

Our data confirm earlier observations that the extent of basalt alteration is controlled by the temperature and permeability. Figure 14 presents a synthetic model of the basalt alteration during the hydrothermal episode.

The second alteration episode, evidenced by K-Ar dating (Table 1 and [START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF] and detected earlier by Ar-Ar of the whole rock samples [START_REF] Elming | Palaeomagnetism and 4 0 Ar/ 3 9 Ar age determinations of the Ediacaran traps from the southwestern margin of the East European Craton, Ukraine: relevance to the Rodinia break-up[END_REF], involved the crystallization of aluminoceladonite and potassium feldspar [START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF] and alteration of dioctahedral smectite into illite-smectite. K-Ar identifies two thermal events on the cratonic foreland of neighboring orogenic systems: in more northern locations (Belarus and Poland) related to the Caledonian orogeny, and in Volynto the Variscan. This late diagenetic alteration affected porous tuffs to a greater extent than flood basalts, and within the flood basalts, it preferentially affected samples more intensely altered by the first alteration episode, i.e. with higher secondary porosity. The temperature of this late alteration was not higher than 85 o C, according to illite-smectite data from the overlying sediments [START_REF] Liivamägi | Paleosols on the Ediacaran basalts of the East European Craton: a unique record of paleoweathering with minimum diagenetic overprint[END_REF] or even lower, according to the organic geochemistry data [START_REF] Goryl | Exceptional preservation of hopanoid and steroid biomarkers in Ediacaran sedimentary rocks of the East European Craton[END_REF]. Apparently, in a high-Mg environment, aluminoceladonite can crystallize at relatively low temperatures.

Inline Supplementary Figure 7 demonstrates that potassium was imported into basalt during the diagenetic alteration. In addition to potassium, enrichments in Cs, Rb, and B in basalts can be also explained by this diagenetic event, based on their observed positive correlations with % K2O and % dioctahedral clay, and established knowledge of their behavior (cf. [START_REF] Derkowski | Cesium, a water-incompatible, siloxane-complexed cation in Earth's upper crust[END_REF] for Rb and Cs, and Środoń, 2010 for B). For tuffs, such reasoning cannot be made, as their original K2O content is variable. Thus it remains unclear whether K2O, Cs, Rb, and B imported into flood basalts were derived locally from felsic tuffs, or if these elements result from a more long-range transport.

Hydrothermal alteration of the Volyn-Brest basalt trap, in continental settings and driven by heat of the cooling traps, can be considered as a good Earth analogue of the Noachian basalt alteration on Mars. The Noachian alteration also produced Fe-rich smectites and hematite (e.g. Ehlmann et al., 2011;Carter et al., 2103;[START_REF] Chemtob | Synthesis and structural characterization of ferrous trioctahedral smectites: Implications for clay mineral genesis and detectability on Mars[END_REF], which are also interpreted as a result of subsurface basalt interaction with heated ground waters (Ehlmann et al., 2011b), and are investigated with the techniques used in this study (XRD: e.g. [START_REF] Blake | Characterization and calibration of the CheMin mineralogical instrument on Mars Science Laboratory[END_REF][START_REF] Downs | Determining mineralogy on Mars with the CheMin X-ray diffractometer[END_REF]; Mössbauer: e.g. [START_REF] Klingelhöfer | Jarosite and hematite at Meridiani Planum from Opportunities Mössbauer spectrometer[END_REF]; NIR: e.g. [START_REF] Carter | Hydrous minerals on Mars as seen by the CRISM and OMEGA imaging spectrometers: Updated global view[END_REF][START_REF] Bishop | Remote detection of clay minerals[END_REF]. Particularly useful for Martian studies should be (1) detail characterization of the Fe-smectite alteration products, which are also dominant on Marsand in particular the role of redox in controlling the di/tri-smectite ratio; (2) detecting Mg enrichment during the alteration, considered indicative for the proposed alteration mechanism, and inconsistent with weathering models (e.g. Le [START_REF] Deit | Extensive surface pedogenic alteration of the Martian Noachian crust suggested by plateau phyllosilicates around Valles Marineris[END_REF]; (3) limited chemical changes despite radical changes in mineralogy, indicative for the proposed alteration mechanism and analogous to the observations of Martian basaltic sediments (e.g., [START_REF] Bristow | Clay mineral diversity and abundance in sedimentary rocks of Gale crater, Mars[END_REF]: and (4) the location of elevated boron contents in illitized dioctahedral smectitic clays, which may explain Martian observations (e.g., [START_REF] Gasda | In situ detection of boron by ChemCam on Mars[END_REF]. The analogy may not be perfect, as the involvement of microorganisms in the alteration process of the Ediacaran Volynian flood basalts is indicated by our data, which provide another piece of evidence for the Ediacaran life on land. Wu, L., Beard, B.L., Roden, E.E., Kennedy, C.B, Johnson, C.M., 2010. Stable Fe isotope fractionations produced by aqueous Fe(II)-hematite surface interactions. Geochimica et Cosmochimica Acta 74, 4249-4265. Table 1. Decomposition of Mössbauer spectra of rocks and selected clay fractions, % Fe 2+ in total Fe calculated from the decomposition data for bulk rock and clay fraction, oxygen and iron isotope data, along with % primary minerals from Supplementary Table 1.

Tables

Table 2. Chemical composition (wt% SEM/EDS microchemical analyses) of isotropic chlorophaeite and replacing birefringent smectititic clays in different sites of the basaltic rock affected by incipient alteration (sample 4522-1). Ferrous state has been assigned to iron in all analyses.

Table 3. K-Ar data for selected clay fractions containing aluminoceladonite and/or illitesmectite (cf. Supplementary Table 3). Table 4. Structural formulae of smectites from Volynian basalts, presented along with formulae derived from published chemical data of similar materials: griffithite from Ross (1960) and [START_REF] Komadel | Reduction of Fe(III) in griffithite[END_REF], oxysmectites from [START_REF] Dainyak | New variety of trioctahedral smectite from effusive basalts of Tunguskaya Sineclise[END_REF], and ferrosaponite from [START_REF] Jambor | New mineral names[END_REF]. Formula of chlorite (Kob-23) added to show tendency in chemical composition of clays.

Figure 1. Map of the Volyn-Brest volcanic province with sampling sites indicated (modified from [START_REF] Kuzmenkova | Comparison of Neoproterozoic Volyn-Brest magmatic province with large world continental plate basalt provinces, the nature of low-and high-Ti basite magmatism[END_REF], TTZ after Mazur et al., 2018b). 0,04 0,01 0,02 0,00 0,01 0,02 0,03 0,01 0,01 Mg 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 NH4 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 The fractions were collected by centrifugation, coagulation by adding concentrated NaCl solution, dialysis, and freeze-drying.

XRD of clay fraction

Clay mineral identification in the bulk rock was verified and made more accurate by studying 2-0.2 and <0.2 m fractions of selected samples. XRD patterns were recorded in airdried and glycolated state of the oriented slides, prepared by sedimentation of 10 mg clay/cm 2 from a sonified water suspension, with 4% MoS2 as the internal standard. The collected XRD patterns were first analyzed using Mering rules for the identification of clay minerals [START_REF] Moore | X-ray Diffraction and the Identification and Analysis of Clay Minerals[END_REF] and then percent smectite layers in mixed-layer illite-smectite (%S) was measured from the selected peak positions on glycolated patterns, using the regressions based on data of Środoń et al. (2009).

K-Ar dating of illite

Na-exchanged, 2.0-0.2 and <0.2 µm clay fractions containing illite-smectite and aluminoceladonite were dated by K-Ar method to establish the age of crystallization of these minerals. All the measurements were performed at the Institute of Geological Sciences, Polish Academy of Sciences in Kraków using Nu Instruments Noblesse multicollector noble-gas spectrometer. Four portions of samples: two 10 mg and two 50 mg, were weighted using Sartorius mass balance with precision of 0.01 mg. A small quantity of CuO (approximately 0.1 mg) was added to the two lightest portions. These samples were then wrapped in an Alfoil and proceeded to radiogenic argon measurements. Two heaviest portions were weighted directly on platinum evaporators and proceeded to potassium measurements. All the samples were equilibrated in the ambient atmosphere (air-dry state) and weighted one after another in order to minimize the errors due to relative humidity affecting clay hydration.

Radiogenic argon measurements were performed on Nu Instruments Noblesse multicollector noble-gas spectrometer (NG039). The spectrometer is equipped with a Niertype ion-source, 75° magnetic sector, two quadrupole lenses and a collector block. The collectors assembly consists of a Faraday detector (FAR) and three ETP ion-counting electron multipliers (IC0, IC1, IC2 Gas aliquot released from the sample was measured three times or more by opening and closing valves of the line in a certain sequence.

The amount of the original aliquot of 38 Ar spike was determined by measuring international standard GL-O (Odin, 1982: 24.85 ± 0.24 10 -6 cm 3 /g STP 40 Arrad). This standard is measured at least two times with every batch of the samples. Standard deviation of received results for GL-O is ± 0.09 (± 0.38%) 10 -6 cm 3 /g STP 40 Arrad for 19 measurements. Biotite LP-6 was also tested for comparison (Odin, 1982: 43.26 = 1583.5 [START_REF] Lee | A redetermination of the isotopic abundances of atmospheric Ar[END_REF]. The potassium contents were measured using Sherwood Model 420 flame photometer. The maximum error of this measurement was estimated to be equal 0.03% K. Based on 6 measurements, K content in GL-O glauconite was slightly higher: 6.61 ± 0.02%, than that of standard: 6.56 ± 0.06% [START_REF] Odin | Interlaboratory Standards for Dating Purposes[END_REF]. LP-6 biotite has slightly lower amount of K: 8.27 ± 0.022% (4 measurements) comparing to reference value: 8.33 ± 0.03% [START_REF] Odin | Interlaboratory Standards for Dating Purposes[END_REF]. Both measured standards are within reference ranges assuming 0.03% K error.

Age errors were calculated from the law of error propagation, taking into account uncertainties of:

spectrometric measurement of argon isotopes determined by spectrometer's software delivered by Nu Instruments -sum of uncertainty for background and sample, -weighting -assumed to be equal to ± 0.03 mg, -potassium measurementsequal to ± 0.03 % of absolute potassium content, -normalization of amount of 38 Ar in spike based on dating of GL-O standardassumed to be equal to ± 0.4 %, -assessment of 40 Ar/ 36 Ar and 40 Ar/ 38 Ar ratios, measured every day for air aliquotsequal to ± 0.017 % and ± 0.013 %, respectively.

Chemistry of the bulk rock and clay separates

Samples were ground in an agate mill and analyzed for the majority of elements at the Bureau Veritas Mineral Laboratories (ACME) in Perth, Australia using a combo package, which requires a single lithium borate fusion and provides major elements from X-Ray Fluorescence (XRF) and trace elements (including REE) from Laser Ablation Inductively Coupled Plasma Mass Spectrometry (LA-ICP-MS) on the same fused disc. Samples were pre oxidized using 66:34 flux containing 10% LiNO3 in order to convert sulfides into sulfates. All elements were measured on the oven dry (105°C) weight basis. Loss on Ignition (LOI) was determined from a separate 1g split using a Thermo Gravimetric Analysis (TGA) system with furnaces set to 110 and 1000 o C. Boron content was measured by the Prompt-Gamma Neutron Activation Analysis (PGNAA) with 0.5 ppm detection limit at the Activation Laboratories, Canada.

Mössbauer spectroscopy of the bulk rock and clay separates

The Mössbauer spectroscopy was used in order to quantify the oxidation state of iron (recent review of the method applied to minerals in [START_REF] Murad | Mossbauer Spectroscopy[END_REF]. The Mössbauer spectra, both for the bulk rock and for mineral separates, were recorded in the transmission geometry at room temperature using a Wissel spectrometer. In addition, both fresh and hydrothermally altered basalts were studied as reference. The samples were prepared in form of pellets, 12 mm in diameter, containing 100 mg specimen mixed with 200 mg lucide powder. The velocity scale calibration was determined from -Fe foil standard.

The spectra were evaluated with the use of WinNormos-for-Igor program. The spectra were decomposed into sets of Zeeman sextets and quadrupole doublets. For each subspectrum the hyperfine field B, isomer shit IS and quadrupole splitting QS were determined and used for the identification of the iron bearing fractions. For natural mineralogical samples, due to the distribution of the hyperfine parameters and overlapping of the subspectra, such identification may be sometimes ambiguous. However, the distinction between Fe 2+ and Fe 3+ states is obvious, due the significant difference in IS. The Fe 2+ /Fe 3+ ratio was determined as the area ratio of the corresponding subspectra. The thickness saturation effects and differences in the recoil free fraction of Fe in different states were not included into analysis. The typical accuracy of the relative contribution of each component is 1%. The % Fe 2+ values for clay fractions are less accurate, in particular if abundant pyroxene is present.

Fe isotopes of the bulk rock and clay separates

Fe isotope analyses ( 56 Fe/ 54 Fe and 57 Fe/ 54 Fe ratios, reported as δ 56 Fe and δ 57/54 Fe) were performed for bulk rock samples and clay separates to trace possible microbial influence (cf. [START_REF] Kappler | Evidence for equilibrium iron isotope fractionation by nitrate-reducing iron(II)-oxidizing bacteria[END_REF]. The measurements were conducted by ALS Scandinavia AB in Lulea, Sweden, by MC-ICP-MS technique of [START_REF] Malinovsky | Performance of high resolution MC-ICP-MS for Fe isotope ratio measurements in sedimentary geological materials[END_REF].

Oxygen isotopes

Oxygen isotopes were measured by Actlabs (Canada). Oxygen was extracted from 5 mg samples at 550-600°C according to the conventional BrF5 procedure of [START_REF] Clayton | The use of bromine pentafluoride in the extraction of oxygen from oxides and silicates for isotopic analysis[END_REF] and analyzed via dual inlet on a Thermo-Finnigan DeltaPlusXP Isotope-Ratio Mass Spectrometer (IRMS). The δ 18 O values are reported in permil (‰) relative to Vienna Standard Mean Ocean Water (VSMOW), with a precision of about 0.1‰.

FTIR measurements

Middle infrared (MIR) investigations (4000 -400 cm -1 ) were performed in transmission mode using a KBr beam-splitter and a DTGS/KBr detector on a Nicolet 6700 FTIR spectrometer. The spectra were obtained using KBr pressed pellets, 13 mm in diameter.

The pellets were prepared by mixing 1 mg of sample with 150 mg of well ground KBr. Then, the mixture was pressed for 5 min at 5 kbars and 4 min at 12 kbars. The pellets thus prepared were kept at 110 °C overnight in an oven before analyses. The spectra were recorded at 4 cm -1 resolution, from 100 co-added scans. Near infrared (NIR) spectra (7500 -4000-cm -1 ) were acquired with a Thermo-Scientific Integrating Sphere (diffuse reflectance) equipped with an internal InGaAs detector, set on a Thermo-Scientific Nicolet 6700 FTIR spectrometer. A white light source was used, associated with a CaF2 beam splitter. The resolution was set at 4 cm -1 with the co-additions of 100 scans. Samples were analyzed directly on the integrating sphere.

Organic geochemistry

Seven bulk rock samples (basalts or clays in basalt cavities: 4600-1, 4600-3A, 4600-3B, and paleosols on basalts: 4504-3, 4504-4, 4504-5 and 4517-1) were used in the organic geochemistry study. Total carbon (TC), total inorganic carbon (TIC), and total sulphur (TS) contents were measured using an Eltra CS-500 IR-analyzer with a TIC module. TC was determined using an infrared cell detector in CO2 gas evolved by combustion in an oxygen atmosphere. TIC contents were derived from a reaction with hot 15% hydrochloric acid; the resulting CO2 was determined using an infrared detector. Total organic carbon (TOC) was calculated as the difference between TC and TIC. Calibration was made according to the Eltra standards. The samples crushed to ca. 100 mesh were extracted using a dichloromethane (DCM)-methanol mixture (1:1 v:v) with an accelerated Dionex ASE 350 solvent extractor.

Elemental sulphur was removed using an activated copper mesh. Due to low TOC values and

  : smectite+zeolites correspond to the 50-200 o C range, trioctahedral mixed layer clays dominate over the 200-220 o C range, then chlorite takes over and albitization starts, epidote appears at ca. 240 o C, actinolite at 275 o C, and amphibole above 350 o C. Analogous basalt alteration sequences have been identified in other hydrothermal fields, e.g. in a

  4 ‰. Basalts with increasing amount of alteration as well as tuffs have elevated  18 O values ranging from 10.1 to 20.0 ‰, similar to the isotope range of clay separates (17.8 to 21.0 ‰), and green clays from veins and cavities (10.1 to 17.0 ‰). The paleosol samples have elevated oxygen isotope values of 21.3 to 21.7 ‰ and the chlorophaeite sample has the highest measured value of 23.0 ‰.

Figure 2 .

 2 Figure 2. Random powder XRD patterns of the bulk rock, with ZnO internal standard,

Figure 3 .

 3 Figure 3. Random powder XRD patterns of the chlorophaeite accumulations, registered in

Figure 4 .

 4 Figure 4. 57 Fe room temperature Mössbauer spectra of bulk rock basalts, chlorophaeite

Figure 5 .

 5 Figure 5. FTIR transmission KBr spectra (<0.2 m fraction) of basaltic clays: Mokrany-3B

Figure 6 .

 6 Figure 6. FTIR transmission KBr spectra of chlorophaeite accumulations (a) and surrounding

Figure 7 .

 7 Figure 7. Partial m/z 71 mass chromatogram for TMS derivatized basalt extract showing the

Figure 8 .

 8 Figure 8. Photomicrographs (optical microscopy) and SEM images (BSE mode) of isotropic

Figure 9 .

 9 Figure 9.

Figure 10 .

 10 Figure 10. Primitive mantle normalized REE spectra of basalts with variable degree of

Figure 11 .

 11 Figure 11. Relationship between % primary minerals from XRD and  18 OVSMOW in basalts of

Figure 12 .

 12 Figure 12. The oxygen isotope compositional ranges of smectites from paleosols, basalt, and

Figure 13 .

 13 Figure13. Relationship between % primary minerals from XRD and percent divalent iron in

Figure 14 .

 14 Figure 14. A synthetic model of basalt alteration during the hydrothermal episode, presenting

FAR

  detector (corresponding to 650000 cps on multiplier). Then, each sample was melted in a double-vacuum resistant crucible at above 1200 o C (assessed visually). CuO added to the samples enhances oxidation of organic matter during this step. Titanium sponge (>99.99% Ti) was always put to the place of the reactor (~0.5 g), where samples are heated. This was a first cleaning level of gases extracted from the samples. Most of the cleanup was achieved when temperature of the reactor was decreasing, which was monitored on Pirani gauge. Pure 38 Ar produced by the Institute for Inorganic and Physical Chemistry of Bern University was used as the spike. It was introduced to the extraction line after the sample extraction using a calibrated pipette, built at Heidelberg University. Impurities of 40 Ar and 36 Ar in spike are regularly monitored and the final results are corrected for these impurities. The final purification of argon was carried out in an isolated section of the line by a getter pump (D-100, SAES Getters), previously baked overnight to remove excess argon. The amount of gases poured to the spectrometer were optimized to keep 40 Ar at level below 0.15 V, because otherwise fractionation on Nier ion source can occur (Kellett and Joyce, 2014).

Table 1 .

 1 Decomposition of Mössbauer spectra of rocks and selected clay fractions, % Fe2+ in 1508 total Fe calculated from the decomposition data for bulk rock and clay fraction, oxygen and 1509 iron isotope data, along with % primary minerals from Supplementary Table1. 1510

	0,028 0,035 0,040	0,047

Table 2 .

 2 Chemical composition (wt% SEM/EDS microchemical analyses) of isotropic chlorophaeite and replacing birefringent smectititic clays in different sites of the basaltic rock affected by incipient alteration (sample 4522-1). Ferrous state has been assigned to iron in all analyses.

	Petrographic entity	Na2O	MgO	Al2O3	SiO2	K2O	CaO	TiO2	MnO	FeO	S
		0,49	2,96	5,61	48,41	0,56	1,75	0,06	0,14	22,63	8
	Isotropic
	chlorophaeite	0,45	2,84	5,89	49,39	0,54	1,79	0,11	0,15	22,54	8
		0,3	3,31	6,05	50,17	0,54	2	0,07	0,19	23,65	8
		0,32	4,54	8,23	47,69	0,4	1,96	0,16	0,13	19,34	8
	Smectite	0,09	4,25	9,73	46,23	0,87	1,81	0,34	0,13	16,45	7
		0,21	4,91	9,78	49,83	0,54	1,91	0,34	0,15	17,69	8

Table 3 .

 3 K-Ar data for selected clay fractions containing aluminoceladonite and/or illitesmectite (cf. Supplementary Table3).

		Mass					Error	Ave. age	Ave. error
	Sample and fraction	[mg]	% K2O	% K	% 40Ar* Age [Ma]	[Ma]	[Ma]	[Ma]
	Kob-9B 2-0.2m	10,50	7,395	6,139	95,7	446,7	11,7	448,5	8,3
		8,86			95,8	450,3	11,8
	Kob-9B <0.2m	10,91	7,355	6,105	91,4	452,2	12,0	451,1	8,4
		8,93			97,1	449,9	11,8
	Vilch19 <0.2m	10,91	4,115	3,416	93,6	442,7	13,1	439,4	9,3
		8,34			90,6	436,1	13,1
	Ochryny21 2-0.2m	6,97	2,885	2,395	84,5	430,6	15,8	432,6	11,1
		7,71			94,3	434,6	15,6
	Ochryny-21 <0.2m	8,45	3,74	3,105	96,3	420,2	13,1	419,2	9,4
		5,92			96,4	418,2	13,3
	Tashki1 2-0.2m	8,66	1,785	1,482	89,2	298,2	19,9	297,1	14,1
		10,07			76,2	296,1	19,9
	Tashki-1 <0.2m	10,81	2,07	1,718	92,9	310,9	16,3	309,0	11,4
		9,02			94,9	307,2	15,9

Table 4 .

 4 Structural formulae of smectites from Volynian basalts, presented along with formulae derived from published chemical data of similar materials: griffithite from Ross (1960) and[START_REF] Komadel | Reduction of Fe(III) in griffithite[END_REF], oxysmectites from[START_REF] Dainyak | New variety of trioctahedral smectite from effusive basalts of Tunguskaya Sineclise[END_REF], and ferrosaponite from[START_REF] Jambor | New mineral names[END_REF]. Formula of chlorite (Kob-23) added to show tendency in chemical composition of clays.

	Sample	Mokrany-3B < 0.2	4504-1A < 0.2	4600-1 < 0.2	Ross griffithite	Komadel griffithite	Oceanic oxysmectite	Continental oxysmectite	Ferrosaponite	Kob-23 bulk
	Fraction Fe 3+	0,87	0,75	0,56	0,46	0,91	0,54	0,83	0,25	0,00
	Si	3,90	3,47	3,21	3,17	3,34	3,40	3,23	2,91	3,57
	Al IV	0,10	0,53	0,75	0,83	0,66	0,59	0,47	1,03	0,43
	Fe IV	0,00	0,00	0,04	0,00	0,00	0,01	0,30	0,06	0,00
	Al VI	0,50	0,16	0,00	0,03	0,06	0,00	0,00	0,00	1,04
	Fe 3+	0,60	0,54	0,44	0,44	0,75	0,56	1,10	0,52	0,00
	Fe 2+	0,09	0,18	0,35	0,52	0,07	0,47	0,22	1,56	2,63
	Mg	1,11	1,81	2,18	1,88	1,82	1,78	1,27	0,87	1,96
	Layer charge	-0,38	-0,47	-0,42	-0,61	-0,46	-0,44	-0,50	-0,67	-0,13
	No. oct. cat.	2,31	2,68	2,97	2,87	2,70	2,80	2,59	2,95	5,63
	Ca	0,01	0,07	0,01	0,25	0,01	0,06	0,21	0,31	0,03
	Na	0,32	0,32	0,38	0,11	0,43	0,31	0,05	0,04	0,08
	K

  extrapolated using number of atoms in the octahedral sheet given in Table4.

	Smectite	di	tri
	Si	4,00	3,16
	Al VI	0,72	0,00
	Fe 3+	0,68	0,44
	Fe 2+	0,00	0,34
	Mg	0,64	2,27
	 18 0	24,0	14,5
	 56 Fe	0,27	-0,47
	ppm LREE 664,00	0,00
	ppm
	HREE	27,00	4,00
	Interlayer
	charge	0,38	0,47	0,42	0,61	0,46	0,44	0,50	0,67	0,15
	Si/(Si+Al)	0,87	0,83	0,81	0,79	0,82	0,85	0,87	0,74	0,71
	Fe/(Mg+Fe)	0,38	0,29	0,27	0,34	0,31	0,37	0,56	0,71	0,57

  ). The high mass isotope ( 40 Ar) is switchable between Faraday and IC0 multiplier, depending on the height of the measured signal. Due to high mass resolving power (MRP), which is ~5400, it is possible to resolve combined peaks of 36 Ar + H 35 Cl from peaks of 36 Ar + H 35 Cl + 12 C3. The signal of all argon isotopes was always calculated as a difference between measurement for sample and signal measured directly before, when ion pump was pumping the interior of the spectrometer. It was done to correct for background level of 36 Ar + H 35 Cl,38 Ar + H 37 Cl and also of 40 Ar that is always leaking in very small quantities to the spectrometer through gaskets (background signal was at around 60, 40 and 400 cps, respectively; while e.g. for GL-O it is at: ~2000, ~30000, ~5000000 cps, respectively). Amount of35 Cl was also always monitored prior to the sample and when the sample was being measured: differences are maximally of few percent. This indirectly indicates that level of H 35 Cl was constant prior and during measurements and, considering very low background level of 36 Ar + H 35 Cl peak, the used correction was reasonable.Extraction-purification line with all-metal valves is attached to the Noblesse spectrometer. The Al-foil wrapped portions were loaded into a glass container, which is a part of the line. The aliquots were subsequently evacuated to approximately 10 -10 mbar pressure.The line was heated overnight at 200 o C to remove all gases adsorbed on metal parts of the line. Samples were, however, not heated, to avoid any 40 Arrad removal. The reactor was running several blank measurements to achieve signal of 40 Ar lower than 0.01 V measured on

  ± 0.44 10 -6 cm 3 /g STP 40 Arrad) and have 43.08 ± 0.41 (± 0.96%) 10 -6 cm 3 /g STP 40 Arrad. Every day, the 40 Ar/ 36 Ar and 40 Ar/ 38 Ar ratios are measured for air sample aliquots, delivered from a calibrated air pipette. Based on this results 40 Ar/ 36 Ar and 40 Ar/ 38 Ar ratios were corrected for instrument mass fractionation and detector efficiencies assuming atmospheric ratios of ( 40 Ar/ 36 Ar)air = 298.57 and ( 40 Ar/ 38 Ar)air

Acknowledgements

Authors acknowledge financial support from the Polish National Science Centre MAESTRO grant 2013/10/A/ST10/00050, from the European Union (ERDF), and Région Nouvelle Aquitaine. We thank Anatoliy Makhnach for helpful discussions, Mariusz Paszkowski for the pyroxene separation, Małgorzata Zielińska for the clay separations, Zuzanna Ciesielska for collecting the XRD data, Michał Banaś for Ar isotope measurements, and Dorota Bakowska for K measurements. Many thanks to the State Enterprise "Ukrainian Geological Company" and Liubomyr Shymkiv personally, and the Republican Unitary Enterprise "Research and Production Center for Geology", and Alla Laptsevich and Sergei Mankevich personally for providing access to the core material and core descriptions. We also thank Valerian Ciobotaru for providing four Moldavian samples. Exceptionally thorough review by Elisabeth Rampe significantly improved quality of our presentation. Dennis D.

Eberl is thanked for final polishing our English.

PhiRhoz correction). Reproducibility of the standard analyses was close to 1.5% for all of the elements, except Na (3%).

Supplementary Figure 2), which contain mixtures of di-and trioctahedral smectites (Inline Supplementary Figure 3), as evidenced by separate 02 and 06 reflections. Sample 4417-1 (Inline Supplementary Figure 2) contains corrensite, chlorite and illite-smectite with ca. 40% S. Among clays filling cavities in basalts, pure trioctahedral minerals were identified: saponite (4600-1), and chlorite with minor corrensite (Kob-23: both shown in Inline Supplementary Figure 2). Sample 4600-1 contains a minor admixture of corrensite in 2-0.2 m fraction.

Tuffs

Clay fractions of tuffs differ from flood basalts by the presence of aluminoceladonite in many samples: from a trace to the dominant component. The trace amounts are associated with illite-smectite of >70%S (Inline Supplementary Figure 4: Pinsk-40), sometimes also with small amounts of chlorite-smectite close to corrensite (Inline Supplementary Figure 4: Tashki-10), and characterize rocks with low quartz content (<10%). In one sample a chloritesmectite close to chlorite was dominant (Inline Supplementary Figure 4: Pinsk-40C).

Samples with more abundant aluminoceladonite contain illite-smectite of <50 %S, chlorite-smectite close to corrensite, discrete chlorite, and represent rocks with higher quartz content (>10%, Inline Supplementary Figure 4: Pinsk-46). Such samples contain abundant chlorite in 0.2-2 m fraction. Most commonly, the 005 reflection of aluminoceladonite is at 1.978Ǻ (d001=9.90 Ǻ). In the extreme case, aluminoceladonite is the dominant component, (Inline Supplementary Figure 4: Kob-9B). In this sample, aluminoceladonite is accompanied by a regular illite, as 005 and 003 reflections are split into two peaks.

Chemical data

Table 5. Characteristics of end-member di-and trioctahedral smectite from Volynian basalts extrapolated using number of atoms in the octahedral sheet given in Table 4. Supplementary Table 1. Quantitative XRD mineral composition data for basalts, green clays from veins and cavities in basalt, and rhyodacites, plus XRD estimates of chlorophaeite composition. Basalts arranged by alteration steps I-V, based on sum of primary minerals. If data available, basalts are assigned to lower or upper volcanic series. Supplementary Table 2. Quantitative XRD mineral composition data for tuffs and for basalttuff alternation in section Kobryn. Tuffs arranged from felsic to basic, based on the content of quartz. If data available, samples are assigned to lower or upper volcanic series. Supplementary Table 3. Mineral composition of <0.2 m fractions, separated from basalts, green clays in veins and cavities, and tuffs. The peak position data, used in the evaluation of percent smectite layers (%S) in illite-smectite are given along with calculated %S and the average %S value. tr: trace, i-s: illite-smectite, ch-s: chlorite-smectite, 2: angular distance between i-s reflections in 42-49 o 2 range. Supplementary Table 4. Chemical composition of basalts, chlorophaeite, green clays from veins and cavities in basalt, rhyodacites, tuffs, and clay fractions. Major elements and loss on ignition (LOI) given in %, trace elements in ppm. Below average values for different lithologies and compositions of clay fractions along with their bulk rocks are listed. Inline Supplementary Figure 5. 57 Fe room temperature Mössbauer spectra of minerals, used for decomposing spectra of bulk rocks (results of decomposition in Table 1).

Figure captions

Inline Supplementary Figure 6. Relationship between % plagioclase from XRD and % CaO in basalts of variable degree of alteration. Three samples off the general trend (blue diamonds) contain calcite veinlets.

Inline Supplementary Figure 7. Relationship between % dioctahedral clays from XRD and ppm boron in basalts of variable degree of alteration.

Inline Supplementary Figure 8. Relationship between % K2O calculated from XRD contents of K-feldspar (assuming 12% K2O) and dioctahedral clay (6% K2O) and measured % K2O in basalts of variable degree of alteration.

Inline Supplementary Figure 9. Relationship between % P2O5 and ppm LREE in bulk rocks.

Inline Supplementary Figure 10. Quantitative evolution of secondary minerals in the course of basalt hydrothermal alteration traced by percent of primary minerals.

Inline Supplementary Figure 11. Differentiation of tuffs into felsic (close to rhyodacite) and basaltic, based on quartz and K-feldspar XRD contents.

Inline Supplementary Figure 12. Chemical variability of tuffs (felsic vs. basaltic) traced using the most immobile trace elements.

Inline Supplementary Figure 13. Chemical differences between basalts and tuffs illustrated for a single profile (Kobryn) using the most immobile trace elements.

Supplementary Materials

Experimental methods

QXRD of the bulk rock

The mineral composition of whole rock samples was studied by means of quantitative X-ray diffractometry (QXRD). A 3 g sample was spiked with 10% ZnO, ground in the

McCrone mill with methanol, and a randomly oriented preparation was made by side-loading (Środoń et al., 2001). The XRD patterns were recorded in 2-65 o 2CuK range with 0.02 o /sec step on Thermo X'tra diffractometer, the mineral phases were identified and then quantified using QMIN software (by Dr. Marek Szczerba), which matches the sum of the diffraction patterns of pure mineral standards with the pattern of a studied rock sample, allowing for manual corrections. Dioctahedral and trioctahedral clays were quantified as groups by careful 060 reflection fitting (Środoń et al., 2001) and split into species by fitting the peaks in the 17-35 o 2range. Chlorophaeite samples, available in quantities too small to conduct regular QXRD, were analyzed only qualitatively in glass capillaries in 4-39 o 2range, as random powders without the spike.

QMIN software does not provide errors of mineral phase determination. Uncertainty of the QXRD analysis can be however evaluated from the results of Reynolds Cupquantitative mineral analysis international round robin [START_REF] Raven | Outcomes of 12 years of the Reynolds Cup quantitative mineral analysis round robin[END_REF]. We got 4 th place in 2018 Reynolds Cup, with the absolute error per major components of 2.4%.

Considering much simpler mineralogy of the samples analyzed in this study than of the samples of Reynolds Cup, the uncertainty of QXRD measurements should be even smaller.

Clay fraction separation

The clay fractions 2-0.2 and <0.2 m were separated from ca. 50 g of the bulk rock after ultrasonic dispersion of the crushed rock in deionized water. Standard chemical treatments [START_REF] Jackson | Soil Chemical Analysisadvanced Course[END_REF] were used for samples prepared for K-Ar dating, but they were avoided in case of the Mössbauer and FTIR studies, in order to minimize the oxidation of structural iron.

low extract yields, samples were not separated into fractions. The n-hexane aliquots of the total extracts were converted to trimethylsilyl derivatives via a reaction with N,O-bis-(trimethylsilyl)trifluoroacetamide (BSTFA) and pyridine for 3 h at 70 ºC. A blank sample (silica gel) was analysed using the same procedure. Only trace amounts of phthalates and fatty acids (FAs) were detected. All solvents were spectroscopically pure and the n-hexane for derivatization was of super-dehydrated grade. The GC-MS analyses were carried out using an Agilent Technologies 7890A gas chromatograph and an Agilent 5975C mass spectrometer with a Triple-Axis Detector at the Faculty of Earth Sciences, Sosnowiec, Poland. Helium (6.0

Grade) was used as the carrier gas at a constant flow of 2.6 ml/min. Separation was performed in either of two fused silica columns:

(i) J&W HP5-MS (60 m × 0.32 mm i.d., 0.25 μm film thickness) coated with a chemically 120 bonded phase (5% phenyl, 95% methylsiloxane). The GC oven temperature was programmed from 45 ºC (1 min) to 100 ºC at 20 ºC/min, then to 300 ºC (held for 60 min) at 3 ºC/min. Solvent delay = 10 min.

(ii) J&W DB35-MS (60 m × 0.25 mm i.d., 0.25 μm film thickness) coated with a chemically bonded phase (35% phenyl, 65% methylsiloxane). The GC oven temperature was programmed from 50 ºC (1 min) to 120 ºC at 20 ºC/min, then to 300 ºC (held for 60 min) at 3 ºC/min. Solvent delay = 15 min. The GC column outlet was connected directly to the ion source of the MSD. The GC-MS interface was at 280 ºC, while the ion source and the quadrupole analyser were at 230 and 150 ºC, respectively. Spectra were recorded from m/z 45-550 (0-40 min) and m/z 50-700 (> 40 min). The mass spectrometer was operated in the electron impact mode (ionisation energy 70 eV).

Petrographic observations

Petrographic studies were performed on thin sections and rock fragments using a Inline Supplementary Figure 5. 57 Fe room temperature Mössbauer spectra of minerals, used for decomposing spectra of bulk rocks (results of decomposition in Table 1).

Inline Supplementary Figure 6. Relationship between % plagioclase from XRD and % CaO in basalts of variable degree of alteration. Three samples off the general trend (blue diamonds) contain calcite veinlets. Inline Supplementary Figure 7. Relationship between % dioctahedral clays from XRD and ppm boron in basalts of variable degree of alteration.

Inline Supplementary Figure 8. Relationship between % K2O calculated from XRD contents of K-feldspar (assuming 12% K2O) and dioctahedral clay (6% K2O) and measured % K2O in basalts of variable degree of alteration.

Inline Supplementary Figure 9. Relationship between % P2O5 and ppm LREE in bulk rocks.

Inline Supplementary Figure 10. Quantitative evolution of secondary minerals in the course of basalt hydrothermal alteration(traced by percent of primary minerals). Inline Supplementary Figure 12. Chemical variability of tuffs (felsic vs. basaltic) traced using the most immobile trace elements.

Inline Supplementary Figure 13. Chemical differences between basalts and tuffs illustrated for a single profile (Kobryn) using the most immobile trace elements.