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Article

The Fate of “Immobile” Ti in Hyaloclastites: An Evidence from Silica–Iron-Rich Sedimentary Rocks of the Urals Paleozoic Massive Sulfide Deposits

by
Nuriya R. Ayupova
*,
Valery V. Maslennikov
,
Irina Yu. Melekestseva
,
Dmitry A. Artemyev
and
Elena V. Belogub
South Urals Federal Research Center of Mineralogy and Geoecology UB RAS, Institute of Mineralogy, Miass 456317, Russia
*
Author to whom correspondence should be addressed.
Minerals 2024, 14(9), 939; https://doi.org/10.3390/min14090939
Submission received: 15 August 2024 / Revised: 10 September 2024 / Accepted: 12 September 2024 / Published: 13 September 2024
(This article belongs to the Special Issue Mineralization and Geochemistry of VMS Deposits)
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Abstract

:
The formation of Paleozoic silica–iron-rich sedimentary rocks in the Urals volcanic-hosted massive sulfide (VHMS) deposits is considered a result of seafloor alteration of hyaloclastites mixed with calcareous/organic or sulfide material. These rocks host various Ti mineral phases pointing to the transformation of precursor metacolloidal TiO2 phases to disordered anatase during seafloor alteration of hyaloclastites, which was later converted to globules and clusters and further to diagenetic rutile. The LA-ICP-MS analysis showed that the Ti content of hyaloclasts partly replaced by finely dispersed Si–Fe aggregates increases to 540–2950 ppm and decreases (<5 ppm) in full Si–Fe pseudomorphs after hyaloclasts. LA-ICP-MS element mapping reveals the enrichment in V, U, Cr, W, Nb, Pb, and Th of the anatase globules and the local accumulation of Zr, Y, and REE on their periphery. Corrosive biogenic textures in the outer zones of some hyaloclasts and biomorphic aggregates in rocks contain anatase particles in assemblage with apatite indicating the biophilic properties of Ti. This work fills the knowledge gaps about Ti mobilization during low-temperature seafloor alteration of hyaloclastites in VHMS deposits.

1. Introduction

Titanium is a lithophile element commonly regarded as chemically less mobile at low-temperature sedimentary environments [1,2,3,4]. The authigenic Ti minerals, however, are described in sedimentary rocks and metasediments from different world regions and are used for the interpreting their diagenetic evolution [5,6,7,8,9,10]. It is assumed that Ti in these rocks is incorporated in structure of Ti-bearing silicates (biotite, pyroxenes, amphiboles) and can be released during their diagenetic alteration with the formation of authigenic Ti oxides [11,12,13]. The authigenic nanocrystals of TiO2 polymorphs (brookite and anatase) are known in organic-rich sediments that is explained by the mobilization of Ti in low pH organic-rich fluids and further increase in pH results in preferential precipitation of anatase [14,15]. At the same time, ancient or modern submarine silica–iron-rich deposits contain low TiO2 amount (<0.02 wt%) and no data on their Ti minerals are available [16,17,18,19,20,21].
In volcanic-hosted massive sulfide (VHMS) regions worldwide, the silica–iron-rich sedimentary rocks with low TiO2 content are one of the main marking horizons for massive sulfide mineralization (e.g., [16,22,23,24,25]). In modern settings, many seafloor massive sulfide systems are also associated with silica-rich iron-oxyhydroxide deposits [26,27,28,29,30,31,32]. In most works, the ancient and modern silica–iron-rich rocks are interpreted to precipitate in form of silica–iron gels as a result of venting of low-temperature hydrothermal fluids on the seafloor. The high Si and Fe and low Ti contents and thus the high Fe/Ti or (Fe + Mn)/Ti ratios (>25) of rocks are considered a main indicator of their hydrothermal origin based on a general idea of Ti immobility under submarine hydrothermal processes [18,20,25,33,34,35,36]. On the other hand, the higher TiO2 (up to 0.58 wt%) content of silica–iron-rich rocks is consistent with the abundance of extensive alteration of tuffaceous [22] or volcaniclastic [16,20,21,25,37,38] fine-grained detritus at/below seafloor during their formation.
The Urals VHMS deposits contain various textural–structural types of silica–iron-rich rocks. In study of their formation, much attention is paid to halmyrolysis (i.e., seafloor weathering) of hyaloclastites mixed with calcareous/organic or sulfide material [24], as well as to microbial processes [39,40]. It was found that the silica–iron-rich rocks at some VHMS deposits are characterized by a highly variable TiO2 amount and contain single TiO2 oxides [24,41]. In this paper, we review the variations in Ti contents and the position, structure, morphology, paragenetic sequence and chemistry of extraordinary authigenic Ti minerals, which form during the evolution from hyaloclastites to silica–iron-rich sedimentary rocks on example of weakly metamorphosed Urals VHMS deposits. Considering Ti a sensitive indicator of potential halmyrolytic–diagenetic processes in close association with biogenic textures would unravel the Ti mobility during the formation of silica–iron-rich sedimentary rocks related to the VHMS areas worldwide.

2. Brief Geological Characteristic

The Ural Mountains stretch for more than 2500 km separating Europe and Asia in the Eurasian continent (Figure 1a). The Uralides are traditionally subdivided into a number of longitudinal zones that are mainly based on ages and paleogeodynamic setting of the dominant rock complexes and include (from west to east) the paleocontinental (Pre-Uralian, West Uralian and Central Uralian zones) and paleooceanic (Tagil–Magnitogorsk, East Uralian and Transuralian zones) sectors [42,43,44], which are divided by the Main Uralian Fault Zone (Figure 1a). Most well-preserved Silurian–Devonian VHMS deposits occur in the Tagil–Magnitogorsk Zone with the Paleozoic island arc volcanic and volcano-sedimentary complexes. The primary morphology of most VHMS deposits is relatively simple: they represent lenticular ore bodies, often arranged at several levels, with long aprons of clastic sulfide talus along the bedding of host rocks and host massive sulfide ores with fragments of smoker chimneys and sulfidized fauna [45,46,47,48,49]. The silica–iron-rich sedimentary rocks occur at different lithostratigraphic levels of the Urals VHMS deposits (Figure 1b): either within basaltic or dacitic lava clastites and hyaloclastites or in association with massive sulfide ore bodies often as intercalation with sulfide turbidites and chloritized hyaloclastites at the flanks of the deposits. Two main types of silica–iron-rich sedimentary rocks are recognized: jasperites and gossanites [24,50].
Jasperites are orange to red microbreccia-like analogs of jaspers, which are closely associated with hyaloclastites (Figure 2a–e). This type of silica–iron-rich rocks is characterized by abundant features indicative of the replacement of calcareous hyaloclastites by finely dispersed hematite–quartz aggregates [24]. Stratiform lenses, beds, and interlayers of jasperites in hyaloclastites are found in volcaniclastic units of most Urals VHMS deposits. Gossanites are lithified products of simultaneous halmyrolysis of clastic sulfide and hyaloclastic sediments [24,51], which are considered the ancient analogs of gossans or metalliferous sediments associated with modern seafloor sulfide mounds [26,27,29]. They are typically located on sulfide bodies or make up a sedimentary dispersion halo around eroded sulfide mounds in some VHMS deposits (Figure 2f,g). The thickness of the gossanite beds ranges from a few millimeters to 1.0–1.5 m. The contact between the gossanites and underlying massive sulfide ores is gradual due to erratic mixing of clastic sulfides and Fe oxides. At the flanks of the VHMS deposits, the gossanites are intercalated with beds of graded sulfide turbidites and chloritized hyaloclastites (Figure 2g–i).
The gossanites consist of variously oxidized clastic sulfides mixed with hyaloclastic material. Hyaloclasts are chloritized and some of them are transformed to finely dispersed hematite–quartz ± chlorite aggregates. Some gossanite beds intercalated with chloritized hyaloclastites no longer retain their typical purplish color and relic sulfides or other features related to sulfide oxidation and seem to be an intermediate member in a genetic range from gossanites to jasperites (Figure 2i). Authigenic mineral assemblages with rare Hg, Au and Ag tellurides, Cu and Ag sulfides, selenides, bismuthides and native gold occur in gossanites [51,52].
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Figure 1. Position of VHMS regions in scheme of geotectonic zonation of the Urals (a), after [44], and silica–iron-rich sedimentary rocks (b) in generalized cross-sections of key studied VHMS regions with the location of VHMS deposits on them, after [41]. (a), VHMS regions (numbers in yellow circles): 1—Ivdel, 2—Karpinsk, 3—Kaban, 4—Krasnouralsk, 5—Kirovgrad, 6—Rezh, 7—Degtyarsk, 8—Mauk, 9—Miass, 10—Kunashak, 11—Uchaly, 12—Uzelga, 13—Aleksandrinka, 14—Sibay, 15—Baymak, 16—Podol, 17—Buribay, 18—Mednogorsk, 19—Gay, 20—Terensay, 21—Dombarovka, 22—Sredneorsk, 23—Verkhneorsk, 24—Berchogur. (b), I–VII, VHMS regions and their deposits (1–16): I, Mednogorsk (1, Blyava, 2, Yaman-Kasy); II, Karpinsk (3, Valentorka); III, Buribai (4, Buribai; 5, Tashkula; 6, Makan; 7, Oktyabr’skoe), IV, Sibai (8, Novy Sibai); V, Aleksandrinka (9, Aleksandrinka; 10, Babaryk); VI, Dombarovka (11, Letnee; 12, Osennee; 13, Zimnee); VII, Sredne-Orskoe (14, Zharly-Asha; 15, Priorskoe; 16, Limannoe). Formations after [53]: bl, Ludlow Blyava; sh, Ludlow Shemur; bb, Lower Devonian Baymak–Buribai; ir, Eifelian Irendyk; kr, Eifelian–Givetian Karamalytash and Aleksandrinka; ul, Middle–Upper Devonian Ulutau; km, Lower Devonian Kiembai; kb, Lower Devonian Kukbukta; kur, Lower–Middle Devonian Kurkuduk; ml, Eifelian–Givetian Milyasha.
Figure 1. Position of VHMS regions in scheme of geotectonic zonation of the Urals (a), after [44], and silica–iron-rich sedimentary rocks (b) in generalized cross-sections of key studied VHMS regions with the location of VHMS deposits on them, after [41]. (a), VHMS regions (numbers in yellow circles): 1—Ivdel, 2—Karpinsk, 3—Kaban, 4—Krasnouralsk, 5—Kirovgrad, 6—Rezh, 7—Degtyarsk, 8—Mauk, 9—Miass, 10—Kunashak, 11—Uchaly, 12—Uzelga, 13—Aleksandrinka, 14—Sibay, 15—Baymak, 16—Podol, 17—Buribay, 18—Mednogorsk, 19—Gay, 20—Terensay, 21—Dombarovka, 22—Sredneorsk, 23—Verkhneorsk, 24—Berchogur. (b), I–VII, VHMS regions and their deposits (1–16): I, Mednogorsk (1, Blyava, 2, Yaman-Kasy); II, Karpinsk (3, Valentorka); III, Buribai (4, Buribai; 5, Tashkula; 6, Makan; 7, Oktyabr’skoe), IV, Sibai (8, Novy Sibai); V, Aleksandrinka (9, Aleksandrinka; 10, Babaryk); VI, Dombarovka (11, Letnee; 12, Osennee; 13, Zimnee); VII, Sredne-Orskoe (14, Zharly-Asha; 15, Priorskoe; 16, Limannoe). Formations after [53]: bl, Ludlow Blyava; sh, Ludlow Shemur; bb, Lower Devonian Baymak–Buribai; ir, Eifelian Irendyk; kr, Eifelian–Givetian Karamalytash and Aleksandrinka; ul, Middle–Upper Devonian Ulutau; km, Lower Devonian Kiembai; kb, Lower Devonian Kukbukta; kur, Lower–Middle Devonian Kurkuduk; ml, Eifelian–Givetian Milyasha.
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Figure 2. Silica–iron-rich rocks of the Urals VHMS deposits. (a–e), Jasperites (Jasp): (a), gradual transition of hyaloclastites (Hy) to silica–iron-rich rocks; (b,c), uneven replacement of hyaloclastite beds with siliceous-ferruginous material; (d,e), microbrecciated texture of jasperite in the top with hematite (Hem) patches; (f–i), gossanites (Gos): (f), replacement of sulfide ores (Sulf) by gossanite with pseudomorphic hematite clasts and relicts of hyaloclasts; g, gossanite beds intercalated with clastic sulfide and thin hyaloclastite beds; (h), hyaloclastite beds in gossanite; (i), intercalation of hyaloclastite, sulfide and gossanite beds. Polished samples from the Talgan (a,d), Molodezhnoe (b,g), Aleksandrinka (c,h,i), Babaryk (e) and Chebach’e (f) deposits. The scale size of all polished samples is consistent as in Figure 1a.
Figure 2. Silica–iron-rich rocks of the Urals VHMS deposits. (a–e), Jasperites (Jasp): (a), gradual transition of hyaloclastites (Hy) to silica–iron-rich rocks; (b,c), uneven replacement of hyaloclastite beds with siliceous-ferruginous material; (d,e), microbrecciated texture of jasperite in the top with hematite (Hem) patches; (f–i), gossanites (Gos): (f), replacement of sulfide ores (Sulf) by gossanite with pseudomorphic hematite clasts and relicts of hyaloclasts; g, gossanite beds intercalated with clastic sulfide and thin hyaloclastite beds; (h), hyaloclastite beds in gossanite; (i), intercalation of hyaloclastite, sulfide and gossanite beds. Polished samples from the Talgan (a,d), Molodezhnoe (b,g), Aleksandrinka (c,h,i), Babaryk (e) and Chebach’e (f) deposits. The scale size of all polished samples is consistent as in Figure 1a.
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3. Materials and Methods

The horizons of silica–iron-rich sedimentary rocks at the Urals VHMS deposits were sampled from open pits and mines during the detailed study of massive sulfide ores. The jasperite and gossanite samples were collected from more than ten weakly metamorphosed VHMS deposits: Talgan, XIX Parts’ezda, Uzelga, Molodezhnoe, Chebach’e, Uchaly, Aleksandrinka, Babaryk, Sibay, Yaman-Kasy, Blyava, Priorskoe, Shemur, and Novy Shemur. Optical microscopic, electron microscopic, and X-ray studies were conducted at the Institute of Mineralogy, South Urals Federal Research Center of Mineralogy and Geoecology UB RAS (IMin, Miass, Russia). Geochemical mapping and electron back-scatter diffraction (EBSD) were carried out at the Geomodel Resource Center of the Scientific Park of the St. Petersburg State University (SPbU, St. Petersburg, Russia). Prior to EBSD studies, samples were treated with a laser at the Nanophotonics Resource Center, Scientific Park, SPbU.
The content of major elements of bulk hyaloclastites and silica–iron rich rock samples was determined by photometry and titration (SiO2, TiO2, Al2O3, Fe2O3, FeO, MgO) and atomic absorption (CaO) in air-acetylene flame on a Perkin Elmer 3110 spectrometer (Waltham, MA, USA). Published data from [24] were also used for statistical calculations. The X-ray diffraction (XRD) patterns of chloritized hyaloclasts were obtained on a Shimadzu XRD-6000 diffractometer (Cu-Kα radiation with a graphite monochromator, 1°/min, interval 22 of 4–70°).
Samples of iron-silica rocks were polished and then studied on an Olympus BX51 optical microscope in reflected light (dark field regime) to characterize the mineral assemblages. The chemical composition of minerals was analyzed on a Tescan Vega 3 SBU scanning electron microscope (SEM) equipped with an Oxford Instruments X-act energy dispersive spectrometer (EDS). Operation conditions included 20 kV accelerating voltage and 15 nA beam current in a spot mode. The ED spectra were acquired, processed and quantified in the Inca 5.02 software. The EDS detection limit was 0.3 wt%. The analytical total close to 100 wt% (taking into account the calculated H2O content) was used as a criterion of the analytical correctness. The Fe and V valences of minerals were accepted as FeO and V2O5, respectively.
The contents of trace elements of hyaloclasts and silica–iron-rich pseudomorphs after hyaloclasts were determined by application of an Agilent 7700 quadrupole inductively coupled plasma mass spectrometer equipped with a New Wave 213-nm solid-state laser. The analyses were conducted by ablation of 40–60-μm spots at a laser repetition rate of 10 Hz and a laser beam energy of 3 to 4 J/cm2. The analysis time for each spot comprised a 30-s measurement of background and a 60-s measurement of elements. The anatase globules were mapped using their sequential linear burning at a laser beam of 12–30 µm across, a speed of 10–15 µm/s, an energy of 3–4 J/cm2, and a frequency of 7 Hz. The mass-spectrometer was calibrated by multielemental solutions. The trace element contents were calculated in the Iolite program using glass standards NIST SRM-610 and USGS GSD-1G and standardization of basic elements to 100% (86–87% excluding OH–) elements in oxide form [54].
The EBSD analysis of Ti minerals was conducted on a Hitachi S-3400N SEM (Hitachi-Science & Technology, Berkshire, UK) equipped with an Oxford Instruments X-Max 20 EDS (Media System Lab, Macherio, Italy). The operation conditions were as follows: 10 kV accelerating voltage and 1 nA beam current to increase the locality of the ED analysis. The EDS was calibrated against a set of standard natural and synthetic samples (Micro-Analysis Consultants Ltd. standards). The samples were treated with Ar plasma using an Oxford Instruments Ionfab300 etcher (Oxford Instruments Plasma Technology, Bristol, UK), an exposition of 10 min, an angle of 45°, an accelerating voltage 500 V, a current of 200 mA and a beam diameter of 10 cm. The minerals were mapped on a Hitachi S-3400N SEM equipped with an Oxford Instruments HKL NordlysNano EDS (Oxford Instruments Plasma Technology, Bristol, UK), at an accelerating voltage of 30 kV, a beam current of 1.5 nA, an exposition of 0.5 s per pattern, and averaging 5–10 (when mapping) or 60 (for individual patterns) images. The EBSD data fitting and mineral phases indexing of structural data of anatase and chlorite were based on the Inorganic Crystal Structure Database. Phase identification and verification used individual EBSD patterns, whereas mapping was aimed to describe microstructural and microtextural orientations, grains and aggregates description. The orientations of individual crystals of anatase, rutile and chlorite aggregates are shown as Euler color schemes, pole figures and orientation distribution density heatmaps.

4. Results

4.1. Structures and Textures of Silica–Iron-Rich Rocks

Jasperites consist of clastic, contraction, spherical or pseudoglobular hematite–quartz aggregates. The hyaloclast relics are typically chloritized and locally enclose residual albite, quartz phenocrysts and leucoxene. At the bottom of the jasperite beds, the hyaloclasts are first gradually transformed to fine hematite–quartz ± chlorite aggregates (Figure 3a,b) and then their volume is reduced that is confirmed by the formation of syneresis cracks during compaction (Figure 3c). The hyaloclasts are often completely altered by postsedimentary processes leading to the loss of signatures of primary substrate and clastic structure of the precursor hyaloclastite (Figure 3d).
Gossanites are composed of oxidized clastic sulfides mixed with hyaloclastic material in various proportions (Figure 3e,f). Abundant hematite or hematite–quartz pseudomorphs formed after small sulfide fragments (Figure 3g,h). The groundmass of the gossanites consists of finely dispersed hematite–quartz ± chlorite aggregates. Fine hyaloclastic material is converted to hematite–quartz aggregates, while the larger hyaloclastic particles are chloritized or silicified and have only an oxidized rim. Typical gossanites show no globular structures like those found in jasperites. In spite of obvious differences between gossanites and jasperites, they both contain abundant hyaloclastic and/or calcareous materials.

4.2. Mineral Composition of Hyaloclastites and Silica–Iron-Rich Rocks

In hyaloclastites, chlorite is dark green or almost black (isotropic) and forms thin flakes. In a few weakly metamorphosed deposits (e.g., Talgan, XIX Parts’ezda and Priorskoe), the XRD patterns of nonoriented chloritized hyaloclastite samples show an intense reflection of 1.543–1.544 Å, which corresponds to a 060 reflection of a mixed-layered chlorite–smectite phase. Chlorite is characterized by high FeOtotal (30.00–33.24 wt%), lower MgO (9.26–10.61 wt%) and low MnO (0.41–1.13 wt%) contents and locally contains CaO (up to 0.52 wt%), Na2O (up to 0.21 wt%) and TiO2 (up to 0.61 wt%) (Table S1, an. 1–11). The larger homogeneous chlorite flakes contain elevated MgO amount (12.55–15.06 wt%) due to a decreasing FeOtotal amount (28.49 to 23.85 wt%) (Table S1, an. 12–24). Minor illite occurs in assemblage with chlorite.
The transformation of hyaloclasts to finely dispersed silica–iron-rich aggregates begins from their surface, depends on their permeability (Figure 3a), and is accompanied by the accumulation of abundant unidentified small (up to 2–3 µm) TiO2 inclusions. They are chaotically distributed in residual chloritized hyaloclast fragments (Figure 4a,b). Further transformation of hyaloclasts leads to the formation of thin (up to 9 μm) concentric zones enriched in whitish–bluish (in dark field view) TiO2 phases in hematite–quartz (±chlorite) pseudomorphs or around them (Figure 4c). The bluish color locally becomes more intense (up to sapphire) due to the presence of small TiO2 crystallites (Figure 4d,e). The complete replacement of hyaloclasts by silica–iron-rich material is accompanied by gradual disappearance of chlorite and TiO2 phases in hematite–quartz pseudomorphs. The larger (up to 60 μm across) well-formed bluish anatase globules are located closely to the hematite–quartz pseudomorphs (Figure 4f,g). The globules locally exhibit crystallographic boundaries (Figure 4h). The internal heterogeneous texture of globules is related to the uneven distribution of the anatase particles in chlorite matrix (Figure 4i). In the globules, chlorite contains varying TiO2 amount (10.33–13.23 wt%) (Table S1, an. 25–33). The formulas of atypically Ti-rich chlorite based on ten cations deviates from an ideal composition and have a lower cation amount in octahedral sites (Table S1, an. 25–33). Anatase exhibits a gradual removal of typical chlorite constituents and an increasing Ti content (Table S1, an. 34–39). The syneresis cracks in globules are filled with calcite and inclusions of apatite and rare earth element (REE) carbonates (Figure 4i). Authigenic Mn-bearing calcite, abundant apatite (locally, REE-bearing), bastnäsite, REE-bearing xenotime, and zircon occur in assemblages with TiO2 globules in quartz–hematite–chlorite ± calcite matrix between silica–iron-rich pseudomorphs.
The complete transformation of hyaloclastites to jasperites leads to the formation of vague contours of silica–iron-rich pseudomorphs and lamellar hematite crystals in a significantly siliceous ± carbonate groundmass of rocks. Here the TiO2 oxides either occur as single corroded anatase–leucoxene aggregates or are absent. In contrast to jasperites, the gossanites contain rutile aggregates and anhedral titanite. Under an optical microscope, brownish yellow (or grayish in a dark field regime) rutile forms euhedral elongated prismatic crystals (locally, twined) typically located in a bluish anatase substrate (Figure 4j,k). A small amount of FeO (0.15–1.70 wt%), SiO2 (0.01–0.47), Al2O3 (0.01–1.13) and MgO (0–0.20) is detected in the composition of rutile (Table S1, an. 40–44). The crystalline rutile aggregates often merge or form radial bundles (Figure 4l). Apatite is associated with rutile (Figure 4m). In gossanites, titanite is rare and forms anhedral light brown or often colorless aggregates in quartz–hematite groundmass closely associated with chlorite, epidote, apatite and REE-bearing minerals (monazite, xenotime and allanite) (Figure 4n,o). Titanite contains (wt%) Al2O3 (5.24–9.97), FeO (0.25–6.68), TiO2 (23.83–31.15), CaO (23.34–29.65), rare MgO (up to 0.47), MnO (up to 0.22), V2O5 (up to 0.87), and F (up to 1.50) (Table S1, an. 45–53).
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Figure 4. Ti mineralization in silica–iron-rich rocks: (a), chaotic distribution of TiO2 phase in partly altered hyaloclast; (b), detail of the previous figure [41]; (c), concentrically zoned silica–iron-rich pseudomorph with TiO2 phases inside and around; (d), crystalline anatase in silica–iron-rich groundmass; (e), detail of the previous figure; (f,g), anatase globules near the silica–iron-rich pseudomorph; (h), anatase with crystallographic boundaries; (i), anatase particles in finely dispersed Ti-rich chlorite mass; (j,k), crystalline rutile inside anatase [41]; (l), intergrowths of prismatic rutile crystals after anatase; (m), rutile–apatite assemblage in quartz–hematite–chlorite matrix; (n), titanite with monazite inclusions associated with epidote in quartz–hematite matrix [41]; (o), titanite–monazite–epidote assemblage. Reflected light (j–l), dark field (a,c,d,f–h), SEM-photo (b,e,i,m–o). Ant—anatase, Hem + Q—finely dispersed hematite–quartz aggregates, Ca—calcite, Rt—rutile, Ap—apatite, Ttn—titanite, Mnz—monazite, Ep—epidote.
Figure 4. Ti mineralization in silica–iron-rich rocks: (a), chaotic distribution of TiO2 phase in partly altered hyaloclast; (b), detail of the previous figure [41]; (c), concentrically zoned silica–iron-rich pseudomorph with TiO2 phases inside and around; (d), crystalline anatase in silica–iron-rich groundmass; (e), detail of the previous figure; (f,g), anatase globules near the silica–iron-rich pseudomorph; (h), anatase with crystallographic boundaries; (i), anatase particles in finely dispersed Ti-rich chlorite mass; (j,k), crystalline rutile inside anatase [41]; (l), intergrowths of prismatic rutile crystals after anatase; (m), rutile–apatite assemblage in quartz–hematite–chlorite matrix; (n), titanite with monazite inclusions associated with epidote in quartz–hematite matrix [41]; (o), titanite–monazite–epidote assemblage. Reflected light (j–l), dark field (a,c,d,f–h), SEM-photo (b,e,i,m–o). Ant—anatase, Hem + Q—finely dispersed hematite–quartz aggregates, Ca—calcite, Rt—rutile, Ap—apatite, Ttn—titanite, Mnz—monazite, Ep—epidote.
Minerals 14 00939 g004

4.3. Biogenic Textures of Ti Phases

The hematitized filamentous, spherical and microstromatolitic bacterial textures and microfauna are observed in all studied samples of jasperites and gossanites. The Ti minerals are closely associated with biogenic textures. The zoned spherules 120–150 µm in diameter with a core made up of a TiO2 phase and a finely dispersed red–brown silica–iron-rich rim were found in hematitized hyaloclastites (Figure 5a). The mineralized vermiform forms are associated with crystalline TiO2 oxides at the contacts of hyaloclast–silica–iron rich aggregates (Figure 5b). These vermiform aggregates exhibit a ball- or ellipsoid-shape structure (fungal spore-like structure ?, cf., [55]) (Figure 5b).
The numerous thread-like forms ~4 μm across and up to 100 μm long often occur along the margins of the hematite–quartz pseudomorphs (Figure 5c). In these cases, the outer rims of the hyaloclasts locally exfoliate during the replacement by hematite–quartz material with the formation of a dark zone consisting of fine-flake chlorite with quartz and apatite inclusions. This dark zone, in turn, is surrounded by bluish anhedral aggregates enriched in TiO2 (Figure 5d). The edges of the hyaloclasts are unevenly (often asymmetrically) altered. Small tubular hematite–quartz aggregates occur in assemblage with anatase (Figure 5d). Rare anatase crystals are scattered in silica–iron-rich rocks of bacteriomorphic texture (Figure 5e).
In gossanites, the clusters of TiO2 minerals consisting of anatase globules and very small crystalline aggregates associated with apatite and REE minerals (bastnäsite, monazite and xenotime) occur in an axial channel of the hematitized tubeworms (Figure 5f–h). The large white to bluish or grayish blue TiO2 mineral aggregates (up to 200 µm) are localized in the hematite–quartz groundmass of gossanites in association with fossilized tubeworms (Figure 5i).
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Figure 5. Microbial alteration textures of silica–iron-rich rocks: (a), spherulitic zoned TiO2 phases in hematitized hyaloclastite; (b), mineralized vermiform structures at the contact with hyaloclast–silica–iron-rich aggregates; (c), mineralized thread-like forms at the margin of the silica–iron-rich pseudomorphosis [41], (d), anatase rim around tubular silica–iron-rich pseuudomorphosis; (e), fine anatase crystals in silica–iron-rich rock with filamentous texture; (f–h), axial channels of hematitized tube forms: (f), anatase globules, transverse section; (g), anatase segregations associated with apatite, oblique cut; (h), apatite and anatase–rutile clusters, longitudinal section; (i), anatase (blue) and anatase–rutile aggregates (white) in matrix between fossilized tubeworms. Reflect light, dark field.
Figure 5. Microbial alteration textures of silica–iron-rich rocks: (a), spherulitic zoned TiO2 phases in hematitized hyaloclastite; (b), mineralized vermiform structures at the contact with hyaloclast–silica–iron-rich aggregates; (c), mineralized thread-like forms at the margin of the silica–iron-rich pseudomorphosis [41], (d), anatase rim around tubular silica–iron-rich pseuudomorphosis; (e), fine anatase crystals in silica–iron-rich rock with filamentous texture; (f–h), axial channels of hematitized tube forms: (f), anatase globules, transverse section; (g), anatase segregations associated with apatite, oblique cut; (h), apatite and anatase–rutile clusters, longitudinal section; (i), anatase (blue) and anatase–rutile aggregates (white) in matrix between fossilized tubeworms. Reflect light, dark field.
Minerals 14 00939 g005

4.4. EBSD-Based Structure of Anatase and Chlorite

The EBSD analysis was applied for an unidentified mineral phase of hyaloclasts and TiO2 phases in relic fragments of hyaloclasts in silica–iron-rich rocks (Figure 6a). The EBSD phase image shows numerous variously oriented vermiform anatase particles up to 5 µm in size in a non-diffracting hyaloclast matrix. Inverse pole figures of anatase exhibit a disordered structure. The main mass of hyaloclasts with anatase inclusions yields no Euler and phase images and the structure of chlorite cannot be identified in contrast to crystalline quartz, apatite and clinochlore in a syneresis crack of the hyaloclast.
Previous EBSD studies of silica–iron-rich rocks of the Talgan deposit showed the presence of TiO2 minerals up to 30–40 µm in form of anatase single crystals in the fine-grained matrix of rocks (Figure 6b) [41]. The band contrast, Euler color scheme, phase images and inverse pole figures showed that an anatase crystal is associated with clinochlore (former hyaloclasts) without crystallographic orientation. The anatase crystals with corroded edges and chlorite microinclusions in quartz–calcite–hematite (minor hyaloclastic relic) masses were found in silica–iron-rich rocks from other massive sulfide deposits of the South Urals (Figure 6c). In addition, the inverse pole figures show more crystallized forms for calcite and quartz. The Kikuchi patterns corresponding to anatase and rutile (see Figure 4j) are obtained for gossanites (Figure 6d).

4.5. Chemical Composition of Hyaloclastites and Silica–Iron-Rich Rocks

Hyaloclastites of the Urals VHMS areas are characterized by a high TiO2 content (0.32–0.55 wt%) (Table 1). The Fe2O3, SiO2, Al2O3, FeO, and MgO contents of hyaloclastites are comparable with those of ferromagnesian chlorite. The TiO2 content of silica–iron-rich rocks varies depending on the amount of hyaloclastic material. The jasperites with relics of hyaloclasts, which occur in transition zones from hyaloclastites to hematite–quartz layers or among hyaloclastites, contain the highest TiO2 (up to 0.71 wt%), lower Al2O3 and MgO, and higher SiO2 and FeO contents in comparison with hyaloclastites (Table 1). The Fe2O3 content is comparable with that of hyaloclastites. The Al2O3, MgO, and FeO contents decrease with an increasing SiO2 value.
The jasperites with fully replaced hyaloclasts demonstrate the lowest (wt%) TiO2 (<0.05), Al2O3 (0.10–3.51), FeO (0.02–3.93), and MgO (0.04–0.45), high SiO2 (60.92–86.87 wt%), and widely varying Fe2O3 (10.15–29.61 wt%) and CaO (0.20–9.70 wt%) contents (Table 1). Gossanites contain high Fe2O3 (up to 88.09 wt%) and SiO2 (up to 57.76 wt%) amount, as well as up to 25.12 wt% Ssulfide due to a large amount of relic sulfides (Table 1). The TiO2 content of gossanites is comparable with that of jasperites. The high TiO2 content (0.11–0.45 wt%) of gossanites is characteristic for oxidized sulfide–hyaloclast beds comparable with silica–iron-rich rocks with hyaloclast relics (Table 1).

4.6. LA-ICP-MS Composition of Hyaloclasts and Hematite–Quartz (±Chlorite) Pseudomorphs

4.6.1. Hyaloclasts in Jasperites

In the foot of the jasperite bodies, the hyaloclasts have moderate Ti content (68–83 ppm) (Table S2, Figure 7a). The hyaloclasts have high content of (ppm) V (83.2–89.4), Mn (5041–5224), Zn (2391–2533), and Ga (43.7–45.7) and varying amounts of Ca (10–500), P (171–308) and Sc (8.9–12.7). The contents of other trace elements are low (<5–10 ppm). The hyaloclast relics in silica–iron-rich pseudomorphs with TiO2 mineral inclusions are characterized by highly variable contents of (ppm) Ti (547–2950), Ca (680–27900), P (160–590), and K (202–1450). The contents (ppm) of Mn (455–1770), Zn (33–134) and Ga (5.5–20.2) contents are lower and the contents of Sc, Co, Ni, Cu, Y, Zr, Nb, Ba, REEs, W, Pb, Th and U are 1–2 orders of magnitude higher compared to unaltered hyaloclasts (Table S2, Figure 7a). The LA-ICP-MS profiles are characterized by Ca, Ti, Y, REE and Zr peaks indicating the presence of mineral inclusions related to these elements.
The Ti (33–1290 ppm), Ca (580–223,000 ppm), P (131–1660 ppm), K (21–1150 ppm), and Mn (1688–4830 ppm) contents of finely dispersed hematite–quartz (±chlorite) pseudomorphs (partly replaced hyaloclasts) are comparable with those of hyaloclast relics. The contents (ppm) of Sc (14.2–29.7), V (91–178), Co (14.5–29.8), Ni (14.5–31.9) slightly increase and the Y, Zr, Nb, Ba, REE, Nb, W, Pb, Th and U contents are significantly lower (Table S2, Figure 7a). The contents of all trace elements in chlorite-free hematite–quartz aggregates of jasperites (Table S2, Figure 7a) are extremely low except for Ca (310–800 ppm), P (108–180), and Mn (21.7–114). Their Ti content is 6.4–11.5 ppm.

4.6.2. Hyaloclasts in Gossanites

In gossanites, the hyaloclast fragments free from inclusions of TiO2 mineral phases are characterized by low Ti (6–20 ppm) and high V (342–427) contents. The hyaloclasts have strongly varying K (431–2294 ppm), Ca (431–7000 ppm), P (70–2390 ppm), and Zn (2729–3059 ppm) content. In comparison with hyaloclasts in jasperites (Table S2, Figure 7b), the hyaloclasts in gossanites have the lowest Mn and Ga contents (1212–1343 ppm and 14.7–18.8 ppm, respectively). The content of other trace elements is very low (<5–10 ppm). The content (in ppm) of Ti (183–599), V (565–757), Y (1.46–28.60), Zr (5.0–35.1), W (22.0–58.8), REEs (5.98–10.64), and U (6.96–19.40) in hematite–quartz–chlorite pseudomorphs are higher, whereas those of Mn (1044–1113), Zn (2342–2829) and Ga (13.0–19.8) remain almost the same (Table S2, Figure 7b). The disappearance of chlorite in the pseudomorphs leads to a decrease in trace element content.

4.6.3. LA-ICP-MS Zonation of Anatase Globules

The LA-ICP-MS element mapping showed that the anatase globules exhibit a striking Ti zonation with a Ti-enriched core and a Ti-depleted rim divided by a narrow zone with an intermediate Ti content (Figure 8). The main trace elements of anatase include V, U, Cr, W, Nb, Pb, and Th. Their distribution (except for Th) in globules is similar to Ti: higher contents in the core and lower in the rim. Similar V, Cr and Pb contents are typical of the rim and the rock groundmass, whereas U, W, Nb, and Th are contained only in the anatase globule. Silica, Al, Fe, and Mg (the main components of chlorite) are mutually correlated and are present in both the anatase globules and the matrix. The outer rim of the anatase globule locally accumulates P and Zr. Carbonates located in syneresis cracks of the anatase globules are characterized by high Mn, Sr, La, Ce and Y contents (Figure 8).

5. Discussion

5.1. Halmyrolytic–Diagenetic Alteration of Hyaloclasts

It is evident from mineralogical studies that the silica–iron-rich aggregates in jasperites and gossanites of the Urals VHMS deposits directly replace the hyaloclast fragments. Enveloping character of primary TiO2 phases and their further transformation to anatase globules near silica–iron-rich pseudomorphs after hyaloclasts indicate the precipitation of the precursor TiO2 phases in unconsolidated sediments.

5.1.1. Pathways of Hyaloclast Alteration

The EBSD studies of hyaloclasts in silica–iron-rich rocks showed a randomly oriented fine-flake disordered chlorite in hyaloclasts. It can be considered a result of spontaneous recrystallization of a precursor amorphous material [54]. It is known that the alteration of hyaloclastites is related to the decomposition of volcanic glass, the replacement of primary glassy or crystalline groundmass, and the formation of authigenic minerals with different density and the degree of compaction of primary pore spaces [56,57]. In these processes, a hydrated hyaloclastite layer is transformed to alumosilicate gel (“gelpalagonite”), which further recrystallizes to smectites, Fe-hydroxides, etc. [58,59,60]. This is consistent with the composition of chlorite of hyaloclasts, which is similar to that of the “gelpalagonite”/clay relics in chlorite, and further diagenetic chlorite after volcanic glass [61,62]. The enlargement of chlorite flakes in hyaloclasts can be related to an increasing temperature of diagenesis [63].
The silica–iron-rich pseudomorphs are a result of Fe2+ oxidation of “gelpalagonite” and synchronous formation of silica–iron-rich gel-like material (e.g., hisingerite) affected by seawater. These pseudomorphs can be transformed to finely dispersed hematite–quartz aggregates during lithification processes, probably, simultaneously with chloritization of hyaloclasts.

5.1.2. Mobilization of Ti and Precipitation of Ti Minerals

The silica–iron-rich pseudomorphs form as some elements of “gelpalagonite” are released to seawater/pore fluid and, vice versa, as some elements are extracted from it. Titanium as TiO2 × nH2O gel is probably released from “gelpalagonite” during the formation of silica–iron-rich gel and could further crystallize as anatase nanoparticles. A disordered anatase structure in hyaloclasts confirms the presence of precursor amorphous Ti phases. The evolution of Ti phases includes the following stages: (i) the precipitation of soft heterogenous colloidal gel with TiO2 × nH2O-enriched “pockets”; (ii) the spontaneous emergence of anatase seeds (they are reported to have the lower formation energy compared to rutile [64]) in areas with local TiO2 oversaturation and (iii) the growth of an anatase crystal until the termination of TiO2 × nH2O supply. Due to low Ti mobility, the anatase crystals are micrometer in size. The complete oxidation of “gelpalagonite” locally leads to higher Ti concentrations and the formation of agglomerates of anatase globules near the silica–iron-rich pseudomorphs after hyaloclasts. The fine-dispersed chlorite–anatase aggregates in globules could indicate that Ti during halmyrolisis is probably retained by clay phases (after the decomposition of hyaloclasts) due to absorption, which is followed by diagenetic precipitation of anatase nano- or microparticles. The growth of the anatase monocrystals is explained by that the thermodynamically unstable TiO2 phases undergo a sequence of irreversible reactions over time to form progressively more stable phases.
Some authors suggest the release of Ti from pore fluids after hydrolysis reactions of Ti-bearing minerals in form of Ti(OH)4 colloids and its precipitation as crystalline TiO2 (pH >5) when dehydrated [65,66]. As a consequence of this process, Ti-bearing gels often precipitate as an “occlusional” phase during authigenesis of clay minerals, limiting the range of Ti mobility to short distances (i.e., from the nanometre scale to the Ti-bearing minerals [4]). It is known that Ti, which results from diagenetic alteration of glassy material, can subsequently be recrystallized to anatase [67]. Kossovskaya et al. [68] believed that Ti can enter the octahedral layers of smectites during glass alteration and can be released and precipitated as Ti phases during the subsequent transformations of smectites. Anatase can occur as pillars in smectite interlayers [69], which were accumulated as a mixture of smectite and amorphous material [70]. The dacitic glasses containing the euhedral anatase crystals of uniform size at the present-day PACMANUS hydrothermal sulfide field (Pacific Ocean) indicate that the precipitation of anatase is a result of the short-range mobilization of Ti liberated during low-temperature glass alteration [71]. Dekov et al. [72] suggested that the Ti enrichment of Fe-hydroxide sediments during volcanic glass alteration is accompanied by the precipitation of anatase, which is a product of alteration of Ti-bearing magnetite from volcanic glasses, and this process is passive (precipitation or/and leaching of other phases) rather than active (due to their mobility).
The stability and growth of TiO2 phases in form of anatase indicate the higher (>5) pH values [73] during the formation of silica–iron-rich sediments, while moderately acidic pH values (3–6) of pore fluids yield the formation of brookite [74]. Further depletion of jasperites in Ti related to the disappearance of anatase can be explained by an increasing pH value of pore water. This is consistent with the completion of Fe2+ oxidation of hyaloclasts, precipitation of authigenic carbonates and microbial mineralization. The increase in pH values up to 8 and more during the oxidation of hyaloclasts and the formation of jasperites [24] probably destabilized the anatase particles. It is known that the high pH value in alkaline environments may facilitate the dissolution of Ti-bearing minerals such as anatase [75]. Experimental data show that the highest solubility of Ti hydroxide (IV) is achieved at pH values of 9.0–9.5 and Ti is mobilized as an anion [Ti2CO3)5(OH)4]5– complex [76,77].
The formation of rutile in gossanites can be related to the oxidation of sulfides, which is favored in acidic conditions (pH < 5) necessary for the conversion of anatase to rutile [74,78]. The nanosized anatase can irreversibly transform to rutile as it coarsens [14]. It is shown that the precipitation, growth and agglomeration of anatase in sediments occur in pore water from low to high temperatures and anatase can directly transform to rutile without intermediate metastable phases [64,79,80]. The fine rutile crystals are stable in acidic conditions, but the stability of the larger rutile crystals is independent from pH values [74]. The formation of rutile crystals associated with hematite and Cu sulfides due to diagenetic recrystallization of a Si–Ti–Fe metacolloidal material similar to anatase leucoxene was registered in ironstones of the Udokan Cu deposit in Siberia, Russia [10].
The assemblage of titanite with epidote, chlorite, muscovite and apatite in gossanites can reflect the direct transformation of glass and clays to chlorite and epidote at an intermediate stage between diagenesis and metamorphism [81]. A combination of rapid sedimentation and abundance of volcanic glass with sulfides during the formation of gossanites could lead to rapid evolution of the composition of pore fluid [81] and titanite could form as a result of modification of the precursor Ti phases due to Ti redistribution upon diagenesis and low-grade metamorphism. Titanite is a typical authigenic mineral of metabasalts of greenschist facies of metamorphism, which forms during the transformation of basaltic glass [82]. Precipitation of authigenic titanite is documented in buried volcaniclastic sandstones and mudstones and its formation is linked to a volcaniclastic source, because Ca, Si and Ti could concurrently be released with the alteration of volcanic rock fragments [83]. The formation of titanite in sedimentary rocks is attributed to zeolite, prehnite–pumpellyite or chlorite facies of metamorphism [84].

5.2. Element Distribution during the Formation of Ti Phase

The LA-ICP-MS studies show that the elevated Ti contents of hyaloclast fragments and highly variable Ti contents of hyaloclasts with inclusions of TiO2 phases and silica–iron-rich pseudomorphs (quartz–hematite–chlorite) are due to the anatase microinclusions. The decrease in Si, Al, Mg, Ca, P, and Na contents and active enrichment in H2O and passive enrichment in Ti and Fe during glass alteration processes have long been known [58,68,85]. The Na loss during the alteration of volcanic glass is consistent with the removal of alkalis to seawater in low-temperature marine environments [86]. Iron is virtually immobile in this process being retained in “gelpalagonite” and the Ti content of an amorphous substance after volcanic glass is 2–3 times higher than its initial concentration in the original glass [59,68,72,86]. In addition, the FeO content increases by a factor of 5, of which only a small part could be attributed to passive accumulation assuming behavior similar to Ti [59].
The formation of Fe-oxyhydroxides accelerates the alteration of glass providing a less-protective gel layer [87]. The alteration processes of basaltic and rhyolitic glasses are generally similar, with leaching of alkalis and accumulation of Fe and Ti on a glass surface [88]. The intermediate products of the formation of silica–iron-rich pseudomorphs in dacitic glass accumulate K, U, Nb, Th, Zr, Sc, Cr, Pb, Zn, Co, Ni, Ba, Th, Pb, Y, and REEs in association with Ti similarly to those of basaltic glass [58,86]. The trace elements likely derive from pore waters in unconsolidated layers rather than from seawater because it has too low concentrations of these elements. The components released from volcanic glass during its palagonitization may not be reprecipitated immediately in adjacent pores, but can migrate to shallower sites and overlying rocks, where smectite forms in other hyaloclastite layers. This, however, does not exclude that U and V can additionally derive from seawater during the transformations of clay material [89,90]. Some data suggest that the incorporation of V4+ in TiO2 phases leads to the preferential formation of anatase [91].
The even Si, Al, Fe, and Mg distribution in the anatase globules (see Figure 8) suggests the amalgamation of abundant TiO2 × nH2O gel phases in the “gelpalagonite” groundmass and their removal during its replacement by silica–iron-rich pseudomorphs. The maturation of gel and the lithification of sediments were accompanied by the desorption of elements accumulated in ferruginous sediments. Some of the elements formed authigenic phases and others are isomorphically incorporated in these authigenic minerals. Vanadium, U, W, Nb, Pb and Th are localized in globules indicating that anatase can serve the area for their release and precipitation and similar behavior with Ti during ferruginization of “gelpalagonite”. The removal of Ca and Mn from hyaloclasts has contributed to the precipitation of Mn-bearing calcite in pore space or syneresis cracks during the lithification processes of the anatase globules. The local Zr, Y, and REE clusters on the periphery of the anatase globules, the presence of calcite veinlets in the globules, and the finding of zircon, bastnäsite and xenotime probably mean that they can form by the precipitation and subsequent crystallization of colloidal particles, but not earlier than anatase.

5.3. Microbial Effect on the Formation of Ti Phases

The occurrence of corrosive biogenic textures in the outer zones of some hyaloclasts (see Figure 5b,c) suggests the microbial mediation during the degradation of hyaloclastic material with Fe accumulation. Key petrographic arguments for the biogenic origin of these corrosion structures include their size similarity with microbes and biotic morphology integrated into Fe-oxyhydroxides. This replacement style of hyaloclasts can be attributed to the release of organic acids by microbes. The abundance of easily oxidized Fe released through glass degradation likely provided a favorable substrate for microbial etching. The calcite matrix and the abundance of apatite in silica–iron-rich sedimentary rocks can be regarded a result of bacterial metabolic activity during a syngenetic period. Phosphorus can be extracted by living organisms and released to form the authigenic apatite (e.g., [92]). The P gain upon the alteration of hyaloclasts can be explained by its scavenging during the decomposition of organic matter by co-precipitation with Fe oxides.
The presence of small anhedral TiO2 particles within bacterial textures could reflect passive Ti accumulation as the surrounding glass is dissolved. The larger accumulations of anatase aggregates in channels of fossil tubes are related to local microenvironment rich in organic carbon and favorable for the complexation of Ti by organic ligands and focused precipitation of TiO2 phases [15]. The maintenance of relatively high pore water Ti concentrations of sediments may be facilitated by the complexation of dissolved Ti by organic matter [93]. It is believed that the biophilic properties of Ti are weak, but there is evidence that some organisms are able to use its soluble forms [94]. It is known that Ti can passively accumulate during etching of volcanic glasses by microbes [95]. In these cases, an origin of microbial tunneling of seafloor volcanic glass that is subsequently chloritized and infilled by titanite has been argued to record the activities of subseafloor microbes. The aggregates with similar textures and morphology but filled almost entirely with fine-grained titanite have been documented in some metabasalts [82].

6. Conclusions

It is shown that halmyrolysis of calcareous hyaloclastites triggered the decomposition of hyaloclastic material with a subsequent release of Ti during the formation of silica–iron-rich sediments and its further aggregation in the form of anatase under slightly alkaline conditions. Anatase in hyaloclastites, which are completely converted to silica–iron-rich sediments (jasperites), disappears at an increasing pH of pore fluids. The oxidation of sulfides in sulfide–hyaloclastite sediments (gossanites) promotes the conversion of anatase to rutile under acidic conditions of mineral formation. Titanite is a result of Ti redistribution at an increasing temperature of diagenesis. The biogenic factor played a major role in Ti accumulation. Our results indicate that the authigenic anatase can be a sensitive indicator of iron depositional environment and is useful in studies of Fe-rich depositional facies related to volcaniclastic rocks.

Supplementary Materials

The following are available online at https://www.mdpi.com/article/10.3390/min14090939/s1, Table S1: Chemical composition (wt%) of minerals in silica–iron-rich sedimentary rocks and Table S2: The contents of elements in hyaloclasts and products of their sequential oxidation (silica–iron-rich pseudomorphs after hyaloclasts. Results of LA-ICP-MS analyses.

Author Contributions

Conceptualization, study of silica-iron-rich sedimentary rocks from the Urals deposits, original draft preparation, N.R.A.; supervision, funding acquisition, V.V.M.; review and editing, I.Y.M., LA ICP MS analyses and mapping, D.A.A.; identification of mineral phases under SEM, E.V.B. All authors have read and agreed to the published version of the manuscript.

Funding

The mineralogical study was supported by the Russian Science Foundation (project no. 22-17-00215). Field trips were conducted in frame of state contract no. 122031600292-6.

Data Availability Statement

All data supporting reported results as Supplementary Materials.

Acknowledgments

The authors are grateful to the Resource Center “Geomodel” of St. Petersburg State University for cooperation (project NO 116234388). The authors thank three anonymous reviewers for their valuable suggestions, which allowed the improvement of the manuscript and Assistant Editor of the journal for the useful recommendations.

Conflicts of Interest

The authors certify that they have no affiliations with or involvement in any organization or entity with any financial or non-financial interest in the subject matter or materials discussed in this manuscript.

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Figure 3. Microtextures of silica–iron-rich rocks. (a–d), Jasperites: (a), replacement of hyaloclasts by finely dispersed quartz–hematite aggregates; (b), relics of hyaloclasts in silica–iron-rich pseudomorphs; (c), silica–iron pseudomorphs with syneresis cracks and small silica–iron particles in siliceous matrix; (d), decomposition of silica–iron pseudomorphs; (e–h), gossanites: (e), chloritized and silicified hyaloclasts in ferruginous matrix; (f), fragments of pyrite–chalcopyrite ore (Py + Chp) clasts and large chloritized hyaloclasts with TiO2 phases and finely dispersed quartz–hematite rim in ferruginous matrix; (g), colloform and zoned hematite pseudomorphs after sulfide particles, (h), replacement of pyrite clasts by finely dispersed quartz–hematite aggregates. Reflected light (g,h), dark field (a–f). Hereinafter, Chl—chlorite (Table S1), Q—quartz, Hem—hematite, Ti—TiO2 phases, Py—pyrite.
Figure 3. Microtextures of silica–iron-rich rocks. (a–d), Jasperites: (a), replacement of hyaloclasts by finely dispersed quartz–hematite aggregates; (b), relics of hyaloclasts in silica–iron-rich pseudomorphs; (c), silica–iron pseudomorphs with syneresis cracks and small silica–iron particles in siliceous matrix; (d), decomposition of silica–iron pseudomorphs; (e–h), gossanites: (e), chloritized and silicified hyaloclasts in ferruginous matrix; (f), fragments of pyrite–chalcopyrite ore (Py + Chp) clasts and large chloritized hyaloclasts with TiO2 phases and finely dispersed quartz–hematite rim in ferruginous matrix; (g), colloform and zoned hematite pseudomorphs after sulfide particles, (h), replacement of pyrite clasts by finely dispersed quartz–hematite aggregates. Reflected light (g,h), dark field (a–f). Hereinafter, Chl—chlorite (Table S1), Q—quartz, Hem—hematite, Ti—TiO2 phases, Py—pyrite.
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Figure 6. EBSD images of anatase and hyaloclasts in silica–iron-rich rocks: (a), variously oriented vermiform anatase particles in non-diffracting hyaloclast matrix; (b,c), single anatase crystals in the fine-grained chlorite–calcite–quartz mass; (d), Kikuchi patterns for anatase and rutile.
Figure 6. EBSD images of anatase and hyaloclasts in silica–iron-rich rocks: (a), variously oriented vermiform anatase particles in non-diffracting hyaloclast matrix; (b,c), single anatase crystals in the fine-grained chlorite–calcite–quartz mass; (d), Kikuchi patterns for anatase and rutile.
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Figure 7. Distribution of trace elements (ppm) in hyaloclasts and their silica–iron-rich pseudomorphs.
Figure 7. Distribution of trace elements (ppm) in hyaloclasts and their silica–iron-rich pseudomorphs.
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Figure 8. LA-ICP-MS elemental maps of anatase globules. Scale bar is 15 µm.
Figure 8. LA-ICP-MS elemental maps of anatase globules. Scale bar is 15 µm.
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Table 1. Major oxide content of hyaloclastites and silica–iron-rich rocks (wt%).
Table 1. Major oxide content of hyaloclastites and silica–iron-rich rocks (wt%).
TiO2Fe2O3FeOAl2O3SiO2MgOCaOLOI *
Chloritized hyaloclastites (n = 11)
av0.3311.5213.5116.5928.2512.252.3912.96
min0.166.345.5613.4322.523.930.387.14
max0.5516.9526.1819.8335.5119.258.6916.25
Silica–iron-rich rocks with relic of hyaloclasts (n = 7)
av0.4612.7117.2213.5638.074.832.197.38
min0.239.6216.3811.7231.904.501.445.48
max0.7115.4917.9615.5141.175.303.3510.86
Silica–iron-rich rocks after hyaloclastites (jasperites) (n = 20)
av<0.0517.103.720.8474.310.362.382.84
min<0.0510.150.020.1060.920.040.200.10
max<0.0529.613.933.5186.870.009.718.58
Oxidized sulfide layers with significant hyaloclasts and sulfides (gossanites) (n = 20)
av0.2538.708.097.4329.852.872.346.13
min0.1120.171.073.118.190.320.203.49
max0.4564.2115.8012.2253.374.576.9320.24
Oxidized sulfide ore with minor hyaloclastic material (gossanites) (n = 15)
av<0.0546.972.181.7837.501.112.964.02
min<0.0531.120.020.054.360.430.360.80
max<0.0588.093.593.5257.762.837.757.69
* High LOI values are related to the presence of chlorite, calcite and sulfides.
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Ayupova, N.R.; Maslennikov, V.V.; Melekestseva, I.Y.; Artemyev, D.A.; Belogub, E.V. The Fate of “Immobile” Ti in Hyaloclastites: An Evidence from Silica–Iron-Rich Sedimentary Rocks of the Urals Paleozoic Massive Sulfide Deposits. Minerals 2024, 14, 939. https://doi.org/10.3390/min14090939

AMA Style

Ayupova NR, Maslennikov VV, Melekestseva IY, Artemyev DA, Belogub EV. The Fate of “Immobile” Ti in Hyaloclastites: An Evidence from Silica–Iron-Rich Sedimentary Rocks of the Urals Paleozoic Massive Sulfide Deposits. Minerals. 2024; 14(9):939. https://doi.org/10.3390/min14090939

Chicago/Turabian Style

Ayupova, Nuriya R., Valery V. Maslennikov, Irina Yu. Melekestseva, Dmitry A. Artemyev, and Elena V. Belogub. 2024. "The Fate of “Immobile” Ti in Hyaloclastites: An Evidence from Silica–Iron-Rich Sedimentary Rocks of the Urals Paleozoic Massive Sulfide Deposits" Minerals 14, no. 9: 939. https://doi.org/10.3390/min14090939

APA Style

Ayupova, N. R., Maslennikov, V. V., Melekestseva, I. Y., Artemyev, D. A., & Belogub, E. V. (2024). The Fate of “Immobile” Ti in Hyaloclastites: An Evidence from Silica–Iron-Rich Sedimentary Rocks of the Urals Paleozoic Massive Sulfide Deposits. Minerals, 14(9), 939. https://doi.org/10.3390/min14090939

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