Sb–Au-bearing chalcedonies in hot geothermal systems: insights from the jasperoids of Poggio Peloso (southern Tuscany, Italy)

Abstract

Detailed characterization was performed on the chalcedonies from the jasperoids of the Pietratonda–Poggio Peloso Sb–Au deposit (southern Tuscany, Italy). The main purpose was to retrieve information on the geothermal fluids that formed the chalcedonies and the source of antimony concentrations. Investigations were performed using optical microscopy, laser ablation inductively coupled plasma mass spectroscopy and X-ray diffraction on both the chalcedonies and the lithotypes cropping out in the area. The results obtained allow the chalcedonies of Pietratonda–Poggio Peloso to be described as a unicum, based on the very high contents of Sb that do not find a comparison in the literature. The textures showed multiple generations of silica that agree well with an environment characterized by multiple injections of mineralizing solutions, bearing variable physicochemical characteristics. The transport likely took place in an alkaline environment, while the acidification of the water may have favoured the precipitation at varying temperatures but not higher than 225 °C. The rocks from which the constituents may have been leached are the hosting carbonates and the surrounding metamorphic rocks. Among the examined rocks, the metamorphic rocks showed the most numerous and significant correspondences with the chalcedonies and were the only ones in which discrete amounts of gold contents were found.

1. Introduction

Jasperoids consist of silica formed by metasomatic carbonate replacement from fluid–rock interaction (Lovering, 1962) in active geothermal systems. In continental settings, many geothermal systems are hosted in carbonate rock volumes that play the role of extensive reservoirs (Barbier, 2002; Faulds et al. 2011; Garland et al. 2012; Brogi et al. 2016). Carbonate rocks, in fact, are prone to forming permeable volumes if interconnected fractures develop during their deformation (e.g. Agosta et al. 2010). It occurs when carbonate rocks are involved in tectonic events at relatively low temperatures (i.e. in the upper crustal level), as in the case of extensional processes triggering detachment zones (Carmignani et al. 2001; Brogi & Cerboneschi, 2007; Matera et al. 2021). In these geological contexts, the carbonate units can offer suitable secondary permeability to host significant volumes of geothermal fluids (Romagnoli et al. 2010; Brogi et al. 2020) and, depending on their chemical features, react with them, depositing neoformational minerals (i.e. banded calcite veins: Hancock et al. 1999; Capezzuoli et al. 2018; Brogi et al. 2021) and jasperoids. Jasperoid formation is often associated with ore development mainly consisting of antimony and gold deposits (Dessau, 1952, 1977; Klemm & Neumann, 1984; Bagby & Berger, 1985; Tanelli et al. 1991; Cline & Hofstra, 2000; Cline et al. 2005; Sillitoe & Brogi, 2021). This is the case for southern Tuscany where jasperoids consist of extensive stratabound volumes (Lattanzi, 1999; Brogi et al. 2011; Brogi & Fulignati, 2012; Morteani et al. 2017) confined between low-permeability units, developed in correspondence with cataclastic horizons formed during the Neogene–Quaternary extensional evolution of the inner Northern Apennines (Carmignani et al. 1994; Brogi et al. 2005). Nevertheless, the metasomatic process triggering the formation of jasperoids and Sb–Au deposition is the result of repeated pulses of fluid injection within the hosting rocks, a consequence of the tectonically controlled dynamic evolution of the geothermal system.

The study of chalcedony and agate generally addresses multiple objectives such as the identification of the sources of the elements, the characterization of the fluids and the reconstruction of the mobilization, transport, precipitation and depositional conditions. The study of chalcedony represents a flourishing research topic, prompting international interest. Several scholars have contributed to the knowledge of these materials, providing essential information on their composition, microstructure and banding (see, e.g. Heaney, 1993; Graetsch, 1994; Heaney et al. 1994; Heaney & Davis, 1995; Merino et al. 1995; Wang & Merino, 1995; Götze et al. 1998; French et al. 2003). Various studies have also been focused on the investigation of their genesis in volcanic (see, e.g. Götze et al. 2001, 2016; Gilg et al. 2003; Moxon & Reed, 2006; Dumańska Słowik et al. 2008; Richter et al. 2015) and sedimentary environments (see, e.g. Moxon & Reed 2006; Götze et al. 2009). Furthermore, the widespread distribution of chalcedony in correspondence with meso- to epithermal ore deposits has also favoured research on polymetallic (Sb, Zn, Pb, Au, Ag) vein-type ore deposits (see, e.g. Dill et al. 2008; Moreira & Fernandez 2015; Radosavljević et al. 2016).

The studies performed on chalcedonies indicate that the deposition temperatures range from 20 to 230 °C (Harris, 1989; Fallick et al. 1985; Saunders, 1990; Heaney, 1993; Götze et al. 2020) and that, in sedimentary and low-temperature hydrothermal environments, near-surface conditions are the most favourable for their formation (see, e.g. White & Corwin, 1961; Cady et al. 1998). In the area under examination, a few studies performed on quartz fluid inclusions estimated homogenization temperatures ranging between 125 ± 3 and 232 °C (unpub. studies performed by Roedder on quartz from Poggio Fuoco, mentioned by Dessau et al. 1972), between 140 and 232 °C (quartz from Pereta) or between 155 and 255 °C (quartz from Poggio Peloso in Brogi & Fulignati, 2012). These temperatures are close to those stated for stibnite precipitation, which has been constrained to between 132 and 245 °C by Lattanzi (1999) and confirmed at 250 °C by Morteani et al. (2017).

Following the work previously carried out in this area, this paper presents results of the study of the chalcedonies from the Pietratonda–Poggio Peloso Sb–Au epithermal mineralization (southern Tuscany, Italy; Fig. 1), which consists of an exhumed hot geothermal system developed during Quaternary time (Brogi & Fulignati, 2012). The main purposes of the study are: (1) to investigate the origin of antimony, gold and other elements present in the examined chalcedonies, by comparing the geochemical features and rare earth element (REE) patterns of the chalcedonies and selected lithotypes which may have reacted with the fluids that deposited the chalcedonies and, (2) to obtain information on the depositional environment of the chalcedonies. To these aims, optical and X-ray diffraction analyses were carried out and major, minor and trace elements were determined on six samples of chalcedonies and 11 samples from selected lithotypes, by laser ablation inductively coupled plasma mass spectroscopy (LA-ICP-MS) and X-ray fluorescence (XRF).

Fig. 1. The geographic distribution of Sb-mineralization in southern Tuscany and northern Lazio. The sampling sites are indicated by the name of the sample except for the chalcedonies (all taken from Poggio Peloso).

2. Geological outline

2.a. Southern Tuscany in the inner Northern Apennines

Southern Tuscany is part of the inner zone of the Northern Apennines orogen (Fig. 1), a Tertiary fold-thrust belt formed during collision between the Adria microplate (part of the African plate) and the Corsica–Sardinia massif (related to the European plate). Late Cretaceous through early Oligocene convergence of the African and European plates by subduction of the intervening western Tethys Ocean led to late Oligocene – early Miocene collision and eastward transport and emplacement of a thrust stack (see, e.g. Molli, 2008). The Northern Apennines stack is formed, from the top down by (1) Jurassic–Cretaceous ophiolite and pelagic sedimentary cover of the Tethyan oceanic realm (Ligurian units, LU) and (2) Cretaceous–Oligocene turbidites representing the continent–ocean transition (Sub-Ligurian units, SLU), both of which were thrust in late Oligocene – early Miocene times over (3) the Upper Triassic to lower Miocene evaporite, carbonate, terrigenous and marine siliciclastic succession of the Adria passive margin (Tuscan Nappe, TN2). The evaporite level of the Tuscan Nappe is made up of partly de-dolomitized, anhydrite-dolostone breccia (known as the Calcare Cavernoso and Burano formations, TN1); its highly permeable and vacuolar texture is due to the interplay among different processes, such as cataclasis, anhydrite dissolution and karstification (see, e.g. Gandin et al. 2000). During early Miocene time, the Tuscan Nappe was thrust over (4) the polydeformed blueschist- to greenschist-facies (Giorgetti et al. 1998; Brogi & Giorgetti, 2012) tectonic wedge complex (known as the Tuscan metamorphic complex; Pandeli et al. 1991) made up of phyllite and metaquartzite of Early to Middle Triassic age (Verrucano Group, MRU3), which overlies (5) phyllite and metacarbonate derived from Carboniferous–Permian protoliths (MRU2). MRU2 overlies (6) Precambrian to early Palaeozoic micaschist (MRU1) and (7) gneissic basement with Precambrian to early Palaeozoic protoliths. Locally, the thrusting caused interleaving of the Calcare Cavernoso and underlying metamorphic units, especially in the upper part of the tectonic wedge complex.

Beginning in early Miocene time, the previously stacked units were affected by extension (Carmignani et al. 2001). Extension produced: (a) lateral segmentation of the previously stacked units and formation of bowl-shaped basins (Brogi & Liotta, 2008; Brogi, 2011); and (b) development of NW-striking Pliocene–Quaternary normal faults, cross-cutting all the previous structures and inducing development of NW-trending tectonic depressions, filled by continental and marine sediments (Martini & Sagri, 1993; Brogi et al. 2013; Martini et al. 2021). The primary evidence of extension is the opening of the Tyrrhenian Basin (Bartole, 1995) and the present crustal and lithospheric thicknesses of ∼17 and 40 km, respectively (Calcagnile & Panza, 1981; Locardi & Nicolich, 1982; Di Stefano et al. 2011; Möller et al. 2013).

At least since Langhian time, extension was accompanied by eastward migrating magmatism (Serri et al. 1993). It affected the area corresponding to the Tuscan archipelago and the inland inner Northern Apennines (Serri et al. 2001; Dini et al. 2005, 2008). Magmas were emplaced at shallow crustal levels (6–8 km depth, mainly: Serri et al. 2001) mostly along NE-striking brittle shear zones playing the role of transfer zones (Dini et al. 2008). Hot geothermal systems (T >200 °C) developed near the intrusive bodies and are now exposed in the whole Tuscan archipelago and southern Tuscany (e.g. Liotta et al. 2010; Vezzoni et al. 2016; Zucchi et al. 2017; Zucchi, 2020). Active systems are presently exploited at depth for geothermal energy production (cf. Liotta et al. 2021) in the Larderello and Monte Amiata areas (Batini et al. 2003 and references therein).

2.b. The Pietratonda–Poggio Peloso area

The Pietratonda–Poggio Peloso area (Fig. 1) is mainly characterized by exposures of jasperoid that formed by the silica replacement of the basal Upper Triassic evaporitic horizon of the Tuscan Nappe (Bernoulli, 2001). This level tectonically overlies the Triassic continental metasiliciclastic succession of the Verrucano Group mainly consisting of quartzite and phyllite with metacarbonate and phyllite at the top (Meccheri et al. 1987; Moretti, 1991; Aldinucci et al. 2005), which recorded a polyphase tectonometamorphic evolution (Aldinucci et al. 2005; Brogi, 2006). At the top of the jasperoid level, discontinuous bodies made up of upper Oligocene – lower Miocene quartz–feldspar sandstone (Macigno Fm) occur. The contact juxtaposing the Macigno Fm on the jasperoid is tectonic and corresponds to a low-angle normal fault (Brogi, 2008). On top, Ligurian units and Miocene sediments formed the cap of the palaeo-geothermal system that was controlled by WSW–ENE strike-/oblique-slip and NNE–SSW normal faults, forming linkage zones between the main strike-slip fault segments (Brogi & Fulignati, 2012). The geothermal fluids consisted of meteoric water circulated to depth and heated owing to the geothermal anomaly. According to Brogi & Fulignati (2012), these fluids reached temperatures of ∼200–225 °C and a maximum salinity of ∼3 wt % NaCl equiv.

2.c. Sb-mineralization and disseminated gold

Southern Tuscany has represented one of the most important mining districts of the Mediterranean basin from the Etruscan period to the twentieth century, with a limited time break during the Roman Empire. Deposits of pyrite, iron, lead, zinc and minor silver and gold are widespread and typically related to the activity of geothermal fluids associated with the Neogene–Quaternary magmatism (Tanelli, 1983; Tanelli et al. 1991; Serri et al. 1993; Lattanzi, 1999; Peccerillo et al. 2001; Dini, 2003; Peccerillo, 2003; Dini et al. 2005; Boschi et al. 2009). The most significant Sb-ore sources are typically located along the NE–SW strip connecting Mount Amiata with Capalbio; however, further deposits are located north of the area of concentration (Fig. 1). The distribution of Sb-deposits shows common features throughout the area: (1) the distribution includes the peripheral zones of present and fossil geothermal areas, coinciding with a Sb mineralized belt; (2) they are associated with fluid flow channelled by extensional faults (Dessau et al. 1972; Tanelli, 1983; Tanelli & Scarsella, 1990; Tanelli et al. 1991; Brogi et al. 2011; Brogi & Fulignati, 2012); (3) they are typically found at the contact between the Upper Triassic evaporite (‘Calcare Cavernoso’ Fm) and the overlying, impermeable, Ligurian and Sub-Ligurian units, which have played a fundamental role in fluid containment (Dessau et al. 1972; Tanelli, 1983; Tanelli & Scarsella, 1990; Tanelli et al. 1991; Lattanzi, 1999); (4) they are characterized by silicic (jasperoid) and, subordinately, argillic alteration (Tanelli & Scarsella, 1990; Tanelli et al. 1991).

The Sb-ore bodies were intensively exploited during the first half of the nineteenth century when at least 17 mines were active (Campiglia, Capita, Casal di Pari, Cetine, Macchia-Casella, Manciano, Montauto, Monticchio, Niccioleta, Pereta, Pietrarotonda–Salamagna, Poggio Fuoco, Rosia, San Martino sul Fiora, Selvena, Tafone, Zolfiere; list provided by Morteani et al. 2017). The deposition is still ongoing as testified to by the stibnite and metastibnite precipitation observed some years ago at the Pereta and Tafone mines (Dessau, 1952; Klemm & Neumann, 1984), as well as by the Sb- and As-rich geothermal fluids (0.2–1.3 mg/L Sb and 22.5–65.4 mg/L As) and the Sb-rich (75 %) scales in the wells (Morteani et al. 2017) and piping (Cappetti et al. 1995; Möller et al. 2009; Morteani et al. 2011) of the Piancastagnaio geothermal field and power plant, respectively.

The Sb-mineralization of Poggio Peloso hill was first documented by Rimbotti (1884) but never exploited due to its low industrial potential (Brizzi & Sabelli, 1985). Reported minerals are: quartz, as the main gangue mineral of Sb-bearing minerals, followed by calcite; stibnite (Sb₂S₃), in radial nodules of thin crystals (2–3 cm); stibiconite (Sb₃O₆(OH)), often pseudomorphic on stibnite; valentinite (Sb₂O₃), generally in spherules with a radius of a few millimetres; rare senarmontite (Sb₂O₃), showing small colourless octahedra; very rare onoratoite (Sb₈O₁₁C_l2); pyrolusite (MnO₂), in the form of concretions of blackish powder; alunite (KAl₃(SO₄)₂(OH)₆), generally present as a white powder often associated with stibiconite; frequent barite (BaSO₄), in tabular crystals typically below centimetre size; and Fe-hydroxides (Braga, 1980; Brizzi & Sabelli, 1985).

Brogi & Fulignati (2012) described the Pietratonda–Poggio Peloso Sb-mineralization as ‘a fossil hydrothermal system that was probably active during the volcanic activity that produced rhyolitic lava flows’ of the neighbouring Roccastrada volcanic complex (∼2.3 Ma). What was the source of the Sb is still a matter of debate. Dessau et al. (1972) suggested that the metals had either a magmatic origin (direct) or that they were leached from the rocks crossed by the fluids (indirect). Lattanzi (1999) suggested the existence of a distal ‘magmatic connection’ with the igneous rocks of the ‘Tuscan Magmatic Province’, even in the absence of a physical contact between the mineralization and the igneous bodies. On this basis, Brogi & Fulignati (2012) speculated on the existence of a similar ‘magmatic connection’ between the Sb-deposits of Poggio Peloso and the Roccastrada volcanic complex.

Gold occurrences have long been associated with Sb- and Ba-deposits (Tanelli et al. 1991), and ‘disseminated’ gold has been found in this ore deposit (Pipino, 1988; Tanelli & Scarsella, 1990; Tanelli et al. 1991; Lattanzi, 1999). Lattanzi (1999) hypothesized that the igneous rocks of the Tuscan Magmatic Province may have been likely gold contributors, as well as the ophiolites of the Ligurian units as suggested by Montini et al. (1995). The latter found a positive Au–Cr correlation in mineralized samples at Frassine (Monterotondo Marittimo, Grosseto) and observed that the only likely source of Cr in the area was the ophiolites occurring in the Ligurian units.

3. Materials

The list of investigated samples is provided in Table 1 and sampling sites are shown in Figure 1.

Table 1. Rock samples list

Abbreviations: Qz – quartz; wm – white mica; Chl – chlorite group; Cal – calcite; Dol – dolomite; Fsp – feldspars; Gp – gypsum; Opq – opaque minerals such as Fe- and Ti-oxides; sul – sulfates of the alunite-jarosite group. Further information on the mineralogical assemblage is provided in online Supplementary Material Note 1.

Six samples of chalcedonies were collected at Poggio Peloso (Fig. 2). All the chalcedonies showed the typical light blue colour, alternating to lighter (up to whitish) or darker (up to greyish) bands. Only sample PP5 showed light orange to light grey/blue bands. The hand specimens showed primary crustiform and colloform textures, while zonal marking was only present in sample PP5. Banding was generally parallel to the rim. Overall, these agates can be described as wall-lining and vein agates. The host rock was the Calcare Cavernoso (Triassic evaporitic Burano Fm) for all samples.

Fig. 2. The chalcedonies from Poggio Peloso investigated in this study.

Considering the various hypotheses formulated on the origin of antimony and gold, it was immediately clear that a sampling exclusively focused on chalcedonies would not have provided convincing answers. Therefore, samples of the carbonate base of the Tuscan Nappe (sample V9), the underlying Verrucano Group (samples V1–V3), the Upper Triassic evaporites (samples V4, V5) and the Palaeozoic phyllites (samples V6–V8) were also collected with the aim of determining their trace compositions. In addition, two samples of Palaeozoic metamorphic rocks drilled in the Monte Amiata geothermal area, consisting of graphitic quartz-metasandstone and metasiltite and graphitic phyllite, coming from the geothermal wells BG3bis (sample A2) and BG25 (sample A1), at depths of 3111 and 3663 m below the ground level, respectively, were also analysed for comparison. Sample A2 was collected in correspondence with a fractured volume (Ruggieri et al. 2004) hosting the deeper (2500–3500 m below ground level) exploited geothermal reservoir (Batini et al. 2003; Bertini et al. 2005).

4. Methods

4.a. LA-ICP-MS

LA-ICP-MS measurements were performed on flat-polished samples. The instrument combines an ablation microbeam based on a Nd:YAG laser source (Brilliant, Quantel) operating at 266 nm (for details see Tiepolo et al. 2003) and a quadrupole ICP-MS (PerkinElmer Sciex-Elan DRC-e). Thirty-four masses from ⁷Li to ²³⁸U were acquired; the laser was operated at a 10 Hz repetition rate, the power on the sample was 1.5 mW and the spot size was set at 50 µm. The optimization of the LA-ICP-MS to minimize elemental fractionation was performed by ablating NIST 610 glass and adjusting the nebulizer Ar and the carrier laser cell He gas flows to obtain the ratio of ²³²Th and ²³⁸U signals close to 1 by minimizing the ThO+/Th+ ratio (<1 %) in order to reduce the formation of polyatomic oxides. Accuracy was assessed on the USGS BCR-2 reference glass (analysed as an unknown in each analytical run) and was better than 20 % at the sub parts per minute (ppm) level. Data reduction was carried out with the software package GLITTER (van Achterbergh et al. 2001) and using NIST SRM 610 and ²⁹Si as external and internal standards, respectively.

4.b. X-ray fluorescence

Major and minor element compositions of the bulk rocks were determined by XRF. Samples were mechanically crushed in a planetary mill and manually ground into a powder in an agate mortar. Quantitative analyses were performed on powder discs obtained by pressing 0.5 g of sample on a support of boric acid. The XRF instrument was a Philips MagiX-Pro. Background and mass absorption intensities were calculated using calibrations based on 24 international geological reference materials. Loss on ignition was determined by heating samples to 1050 °C for 2 hours.

4.c. Optical microscopy

All chalcedonies were cut orthogonally to the banding and polished. Half of the polished sample was prepared in thin-section for optical microscopy (OM) observations, while the other opposite half was investigated by LA-ICP-MS. Rock sample nos V1–V9 were prepared in thin-section to evaluate the possible presence of microfeatures that could lead to their exclusion and to decide the points on which to conduct further analyses.

4.d. Powder X-ray diffraction

An aliquot of ∼20 g was crushed and pulverized in an agate mortar. About 1 g was used for powder X-ray diffraction (PXRD), carried out on an automated PANalytical X’Pert PRO diffractometer, using Cu–Ka radiation (Ni-filter, 2θ range = 5–70°, step size = 0.02°, time per step = 1 s, room temperature = 25 °C). Measurements were performed on powders gently pressed into a back-loading cavity mount. The phase analysis was carried out using the software X’Pert HighScore Plus that uses the Inorganic Crystal Structure Database (ICSD). The starting coordinates of the atoms, cell parameters and space groups were taken from Le Page & Donnay (1976) for quartz, Brigatti et al. (1998) for phengitic muscovite, Walker & Bish (1992) and Guggenheim & Nelson (1993) for clinochlore, Kawano et al. (1998) for sericite, Aleksandrova et al. (1973) for sudoite, Neder et al. (1999) for kaolinite, Cole et al. (1949) for sanidine, Chao et al. (1940) for orthoclase and Ferguson et al. (1958) for albite.

5. Results

5.a. The chalcedonies

5.a.1. LA-ICP-MS

The average composition of all samples is characterized by SiO₂ contents ranging from 98.0 to 99.8 wt % (av. 99.1 wt % ± 0.3) and a rather low minor- and trace-element content, i.e. the typical composition of chalcedonies and agates (see, e.g. Blankenburg, 1988; Tanaka & Kamioka, 1994; Merino et al. 1995; Götze et al. 2001; Möckel & Götze, 2007; Götze et al. 2009, 2016). Some spots with low SiO₂ and high Ca, Mg, Al, K or Fe contents were also identified as carbonates, feldspars and metal oxide inclusions. The presence of inclusions made it necessary to exclude some measurements (fully provided in the online Supplementary Material Tables S1–S6 and Figs S1–S6) from the calculation of the average composition of the chalcedonies provided in Table 2. Table 2 shows that SiO₂ contents of the chalcedony range between 99.2 and 99.4 wt % and that none of the other minor and trace components is above 1000 ppm, except for Sb. The latter show contents (1543–1846 ppm; average 1697 ± 119 ppm) that greatly exceed those reported for the upper continental crust (UCC) (0.4 ppm; Rudnick & Gao, 2003) or for surficial waters (<0.001 ppm) (Filella et al. 2007). To a lesser extent, B and Ge are also enriched, ranging on average between 81 and 113 ppm for the former and between 7 and 9 ppm for the latter. As for Ba contents, it is worth underlining that the amounts reported in Table 2 are underestimated. Indeed, numerous Ba-rich inclusions were found but omitted from the calculation of the average chalcedony composition.

Table 2. LA-ICP-MS data. The chemical composition of the chalcedony. Average values provided as ppm if not differently specified. Measurements of quartz and inclusions were omitted from calculation

The correlations between elements and element pairs are shown in online Supplementary Material Figures S1–S6 and Tables S7–S12, respectively. Here the attention is focused on the correlation coefficients of Sb with major, minor and trace elements (Table 3) and on correlated pairs of elements (Table 4). Sb can be strongly correlated (r > 0.9) with B, Mn, As and Cs or well correlated (0.60 < r < 0.89) to poorly correlated (0.40 < r < 0.59) with all the remaining elements. The correlation with B and, to a lesser extent, Mn, As and Cs appears consistent among the different samples (Table 3; Fig. 3). On the contrary, the correlation with the other elements is only observed in a few samples. The well to strongly correlated pairs Be–Mn, Be–Ga, B–Rb, K–Rb, K–Cs, Rb–Cs and Ba–Sr, and the poorly correlated pair SiO₂–Ge are worth noting (Table 4).

Table 3. Correlation coefficients (r) of Sb with major, minor and trace elements in the chalcedonies (LA-ICP-MS data)

Highest values are marked in bold.

Table 4. Correlated pairs of elements in the chalcedonies (LA-ICP-MS data)

Highest values are marked in bold.

Fig. 3. LA-ICP-MS measurements. The Sb, B, Mn, As and Cs patterns of the chalcedonies from Poggio Peloso.

As for gold, despite that the values are to be considered qualitative, this element has been frequently detected, both in chalcedony and carbonate inclusions. The significance of its correlation is lowered by the small number of available measurements but appears consistent with Ag in sample PP7 (r = 0.865). As for the correlation between chromophores and agate colours, the techniques used are not suitable for proposing conclusive information. For example, (a) the variability of Fe and Mn amounts cannot be correlated to light/dark bands in the geochemical profiles of samples PP1, PP2, PP5, PP7 and PP8, and (b) the whitish and blue bands of sample PP4 show comparable contents of both Fe and Mn. Conversely, the brown band of the latter sample (online Supplementary Material Table S3, measurement no. 11) shows the highest Fe and Mn values.

As for REE contents, the chalcedonies show a depleted chondrite-normalized pattern (Fig. 4a), which is typical of these materials. Except for PP7 showing a La_n/Lu_n ratio below 1, light REEs (LREEs) prevail over heavy REEs (HREEs) in the other samples (La_n/Lu_n ratio 1.3–4.2). The LREE normalized patterns show a decreasing slope in all samples (La_n/Sm_n 1.3–5.2) except PP5 (La_n/Sm_n 0.6). A slight negative Ce anomaly is present in all samples (Ce/Ce* 0.57–0.94), except PP5 and PP7 for which it cannot be determined. Conversely, a weak to strong positive Eu anomaly characterizes all samples (Eu/Eu* 1.67–6.32). The HREE normalized patterns show an increasing slope (Gd_n/Lu_n 0.4–0.6), except in PP1 and PP5 (Gd_n/Lu_n 1.2 and 2.3, respectively). The UCC-normalized REE pattern (Fig. 4b) is particularly depleted and shows an enrichment of HREEs with respect to LREEs. Both La_n/Lu_n and Gd_n/Lu_n ratios are constantly below 1 (range 0.06–0.76), as shown by the overall increasing slope of the patterns. The Ce anomaly is negative in all samples (Ce/Ce* 0.20–0.95) except PP4 (Ce/Ce* 2.05), while the Eu anomaly is always positive (Eu/Eu* 1.03–7.28).

Fig. 4. (a) The chondrite-normalized REE concentration pattern and (b) the upper continental crust (UCC)-normalized trace-element pattern of the chalcedonies (average values from Table 2). (c) The REE patterns of the metamorphic rocks. Black and grey lines indicate the samples investigated in this study. Red lines correspond to the Radicondoli (5.1960 m) drill core chip analysed by Möller et al. (2003) – indicated with ‘R’ – and the phyllites from the well Ciabattino 2 (305 m depth) and the Boccheggiano outcrops (sample nos 10179-B and 10179-D) analysed by Möller et al. (2009). For comparison, the pattern of the chalcedonies is visible in the background. (d) REE pattern of selected Tuscanian volcanites compared to the chalcedonies. Reference data from Giraud et al. (1986) and Pinarelli et al. (1989). The REE concentrations of the chondrite are from McDonough & Sun (1995); those of the UCC are from Rudnick & Gao (2003).

5.a.2 . PXRD

Quartz was identified in all samples. Calcite has been clearly detected in sample PP2 while barite and haematite were rarely detected. The results of the XRD analyses performed on the rocks analysed for comparison are provided in the online Supplementary Material (Note 1) and confirmed the macroscopic identification of the sampled rocks.

5.a.3. Optical microscopy

The samples are mainly constituted by fibrous chalcedony (length-fast) and micro-/macro-crystalline quartz. In the following description, the correspondence with the spot-analyses (i.e. ‘S(no.)-no./nos’) refers to LA-ICP-MS measurements (online Supplementary Material Tables S1–S6). In sample PP1, long fibres of fan-shaped chalcedony with flamboyant extinction (S1- nos 14, 15) are richer in Mg, V, Mn, B, Na, K, Cs, As and Rb, and lower in Be, Al and Cu with respect to the thin parallel-bladed fibres (S1-nos 10, 11). Sb, Sr, Ba and Ga contents are similar in both types of textures. In sample PP2, the sequence includes several textures: fibrous flamboyant (S2-nos 1–4), comb, parallel-bladed (S2-nos 9, 10 and 13–15) and pseudo-granular (S2-nos 1, 11, 12, 16, 17). The last band of granular quartz (S2-nos 22, 23) contains the highest SiO₂ and the lowest levels of almost all other elements, except for Mg, Ca and As. A thick and heterogeneous layer mainly constituted by calcite (S2-nos 5–8) is also present. Sample PP4 shows several generations of hatched bands (‘Runzelbänderung’), constituted by long and thin parallel fibres, alternating with thin layers of comb quartz. In sample PP5, the pseudo-granular texture (S4-nos 3, 4, 7–17) is generally poorer in most elements with respect to the flamboyant spherulites and fan-shaped fibres (S4-nos 1, 2, 5, 6). In sample PP7, the chemical composition varies without an apparent consistency among the different depositional layers. The layers show flamboyant (S5-nos 1, 2, 15–17, 22–27), parallel-bladed (S5-nos 3–8, 20–21) and feathery chalcedony (S5-nos 13, 18), along with comb (S5-nos. 14, 19) and granular quartz (S5-nos 9–12, 28, 29). Similarly, to the previous sample, PP8 also shows parallel-bladed (S6-nos 1, 2, 5, 6), flamboyant (S6-nos 3, 4, 7) fibres along with comb micro- and macro-quartz (S6-nos 8–10). The last layer is constituted by granular quartz and abundant residues of carbonaceous materials (S6-nos. 11, 12).

Regarding the distribution of the elements in the various layers, it may be interesting to note that while the antimony remains at almost comparable levels in all types of textures and generally decreases drastically in correspondence with carbonate inclusions, gold is present in chalcedony and carbonates while absent in comb quartz. A series of photos of the main textures (with a scale of detail greater than those given in the online Supplementary Material) is provided in Figure 5.

Fig. 5. The textures of chalcedony. OM images (crossed polarized light). (a) Parallel-bladed and comb textures. The arrows indicate two distinct layers of comb quartz. Along the lower one, a concentration of iron oxides approximately corresponds to the dark line indicated by the arrow. (b) Parallel-bladed texture. (c) Flamboyant texture. (d) Flamboyant and comb (indicated by the arrow). (e) Feathery quartz (jigsaw-puzzle texture). (f) Pseudo-granular. The scale bar equals 1 mm.

5.b. The rocks

The rocks were analysed only to confirm their identification and to estimate their Sb and Au contents; however, a complete analysis is provided for most of them in the online Supplementary Material (Note 1; Table S13).

The rocks are characterized by low to very low amounts of Sb: below 1 ppm in samples V1, V4, V5, V6, V7, V9 and R; between 1 and 2 ppm in samples V3, V8 and A1; and between 3 and 4 ppm in samples V2 and A2. Among the other elements, while minor and trace elements typically show higher values in the rocks than in the chalcedonies, the contents of Be and Ge appear systematically higher in the chalcedonies. The relatively high contents of B in samples V9 and A2, Mn in samples V9, A1 and A2, as well as the amounts of Au measured in samples V1, V7–V9, A2 and, especially, A1 may be informative for comparison.

Figure 4c shows the REE patterns of the examined rock samples and those of the Palaeozoic phyllites from the Larderello–Travale geothermal area analysed by Möller et al. (2009). The first observation to make is that none of the rock samples show such pronounced depletion as that observed in the chalcedonies except V1. The Ciabattino and Boccheggiano samples appear particularly enriched compared to those investigated in this study, while the Radicondoli sample is comparable with both V6 and V7. The general trend of the phyllites is decreasing except in samples A1 and A2, which also show an increase in HREEs compared to LREEs. The different pattern of these two samples compared to the other rocks is probably due both to the partially different mineralogy (including the relative abundance of the accessory phases) and to the different migration speed of the HREEs compared to the LREEs in different conditions of alteration (in particular pH variability). As they fall beyond the scope of this research, we will not dwell on these aspects, while it is worth noting another important feature concerning the Eu anomalies. While the samples A1, A2 and V3 show a weak positive Eu anomaly (similar to the chalcedonies), the phyllites analysed by Möller et al. (2003, 2009) and the other metamorphic rocks investigated here are characterized by a negative Eu anomaly (Eu/Eu* 0.5–0.9). Furthermore, the samples A1 and A2 show a La_n/Lu_n ratio (0.8–1.2) closest to that of the chalcedonies.

In Figure 4d, the REE patterns of the chalcedonies are further compared with those of the Tuscan volcanites. The data obtained by Giraud et al. (1986) and Pinarelli et al. (1989) on the volcanites from both Roccastrada and Monte Amiata evidence a clear difference not simply related to the absolute REE contents but also regarding the overall behaviour of the REEs.

6. Discussion

6.a. Poggio Peloso chalcedony: an outlier based on Sb contents

The samples from Poggio Peloso contain the highest Sb amounts found so far in chalcedonies. Table 5 shows a selection of reference data including chalcedony, carnelian and agate, as well as quartz, amethyst, chert/flint, jasper and quartzites from several locations. Low amounts of Sb are typically found in all silica-based minerals and rocks; however, the contents measured by the authors in agates rarely exceed 50 ppm.

Table 5. Sb amounts in silica-based minerals and rocks

Abbreviations: N – number of analysed samples; Min. – minimum value; Max. – maximum value; Av. – average value; sd – standard deviation value

* From Götze et al. (2001).

Therefore, the comparison describes the chalcedonies of Poggio Peloso as ‘outliers’ with respect to Sb amounts, precisely as the high contents of Zr (between 295 and 687 ppm) found in vein agates from ‘Borówno’ quarry (Lower Silesia) distinguished the Polish samples (Powolny et al. 2019).

6.b. Positive Eu anomaly: hydrothermal environment

Comparing the REE patterns provided by Morteani et al. (2017) for calcite with those obtained for the Poggio Peloso chalcedony, a striking difference emerges regarding Eu anomalies. While calcite patterns are characterized by negative Eu anomalies, the chalcedonies show positive anomalies. This result represents an anomalous feature even for chalcedonies, and it has been rarely observed (e.g. in vein and moss agates from the ‘Borówno’ quarry in Poland; Powolny et al. 2019).

As for calcite, Morteani et al. (2017) explained the negative anomaly due to the poor Eu²⁺ uptake capacity of Ca minerals, based on the thermochemical reduction of Eu³⁺ to Eu²⁺ by high-temperature deep-seated aquifers. As regards the chalcedonies, it should be remembered that Eu²⁺ is present in a variety of magmatic and hydrothermal environments while it is rare in sedimentary environments because it is linked to particularly reducing conditions (Brookins, 1989). In aqueous solutions, the redox potential of Eu/Eu* is affected by temperature and, to a lesser extent, by pH, pressure and REE speciation (Bau, 1991). As a consequence, positive anomalies such as those observed in the chalcedonies are found in acidic, reducing hydrothermal fluids and sediments in an active ridge system (Kamber & Webb, 2001; Michard et al. 1983; Michard, 1989; German et al. 1990). Furthermore, Eu enrichment may also reflect plagioclase weathering since Eu substitutes into the calcium site of plagioclases (see, e.g. Cabioch et al. 2006). However, Eu²⁺ can also substitute Ba²⁺ (Shannon, 1976; Guichard et al. 1979; Mazumdar et al. 1999), thus making barite a possible carrier phase. In chalcedonies, barite inclusions are indicative of the activity of SO₄ ²⁻ in the mineralizing medium, and Ba is one of the elements promoting silica precipitation in alkaline solutions (Dumańska-Słowik et al. 2018).

Möller et al. (2009) reported the C1-chondrite-normalized patterns of REY in liquid and vapour produced by some wells of the Monte Amiata geothermal field. In particular, the liquid phase is characterized by strong Eu and moderate Y positive anomalies, which were interpreted to be the result of REY fractionation during water–rock interaction in the fluid source region. In particular, the deep fluid productive horizons of the Monte Amiata geothermal field are hosted in the Palaeozoic metamorphic rocks represented by the A1 and A2 samples. It is worth reiterating that such samples, contrary to the other Palaeozoic rocks, are characterized by a weak positive Eu anomaly.

6.c. Silica source and transport and Ge enrichment

The conditio sine qua non for the formation of chalcedonies is the availability of silica. Silica commonly derives from the weathering or hydrothermal alteration of silicate minerals (effective process) or from the dissolution of amorphous silica or of quartz (less effective process). As a consequence, chalcedony tends to form mainly in areas affected by thermal springs or in formations rich in weatherable silicates. Landmesser (1984) has shown that diffusion of monomeric silicic acid (H₄SiO₄) is the main mechanism in silica transport, based on numerous considerations including the large size/low velocity of colloidal silica and its clear prevalence under a wide range of pH conditions (1–9). This mechanism excludes the transport of silica via free liquid solution, which would also be difficult to explain compared to the low solubility of silica (Krauskopf, 1956).

Considering that the host rocks of the Poggio Peloso chalcedonies are carbonates, the high amounts of silica needed for the formation of chalcedony could have been derived from the volcanic and metamorphic rocks present in the territory or from residual magmatic fluids. Comparing REE contents in the surrounding rocks and chalcedonies, it is evident that the volcanic rocks have such different patterns (Fig. 4d) that it becomes difficult to invoke them as the primary silica source. Most of the metamorphic rocks are overall enriched with respect to the chalcedonies, but their patterns show different LREE and HREE trends (Fig. 4c). As shown above, an interesting analogy with chalcedonies is represented by the weak Eu anomaly shown by the A1–A2 samples from the Monte Amiata geothermal field. It is worth noting that the fluids produced from this geothermal field are characterized by significant concentrations of dissolved silica (up to 1160 ppm; Minissale et al. 1997) so that the brines resulting from flashing at the separator are highly supersaturated in silica, causing the formation of silica scales in geothermal pipelines (Vitolo & Cialdella, 1995). Moreover, in one geothermal well, the high silica content is accompanied by significant Sb concentration (up to 50 mg/L), which caused the deposition of SiO₂-rich, stibnite and/or metastibnite-bearing scales in the pipes (Morteani et al. 2011). However, the lack of a convincing direct correlation with any of the surrounding rocks suggests that the silica does not derive from the leaching of a unique Si-rich lithology.

The behaviour of Ge in chalcedonies has been frequently related to that of SiO₂, given the similar geochemical behaviour of these two elements and the preferred structural incorporation of Ge in the quartz crystal structure (Götze et al. 2009). Ge contents measured in the chalcedonies from Poggio Peloso (7–9 ppm) are higher than those reported for the continental crust (1.4 ppm), lower than those shown by most volcanic agates, while slightly higher than those reported for sedimentary agates (Götze et al. 2020; Table 1). As noted in other case studies (e.g. Gliozzo et al. 2019), Ge contents are poorly (PP1, 2, 5, 7, 8) to well (PP4) correlated with those of SiO₂; however, the absolute Ge amounts suggest the possible interaction of volatile fluids (HF) and the transport of SiF₄, GeF₄ and BF₃ compounds (for this chemical transport reaction, see Götze et al. 2012).

6.d. B, Mn, As and Cs and their correlations with Sb: the fluids

6.d.1. Boron

Contents ranging between 81 and 113 ppm are relatively high, both compared with the crustal B abundance (17 ppm) and the typical range of B content in quartz (from 0 to 25 ppm; Stavrov & Khitrov, 1962), where B³⁺ may replace Si⁴⁺ in the SiO₄ tetrahedra (see Müller et al. 2012 for details on the configuration of trace elements in the quartz lattice).

In magmatic systems, the incompatible behaviour of this lithophile element leads to its strong enrichment in magmatic fluids and vapours as well as in hydrous liquids. Hence, boron enrichment in hydrothermal fluids can derive from magmatic fluids, heating and feeding hydrothermal aquifers with volatile components. However, it can also be influenced by the geochemistry of the local rocks (water–rock interactions; see Bernard-Romero et al. 2010) and therefore, being strictly conditioned by the hydrothermal circulation system itself. Among the surrounding rocks, the B-rich Triassic evaporites (Dessau et al. 1972) may have represented the most abundant source of B. In this regard, however, it is worth noting that the A2 sample of the Monte Amiata geothermal field also contains remarkable B amounts (191 ppm) and that also the fluids produced by this field are characterized by significant B concentrations (up 3932 ppm; Minissale et al. 1997).

6.d.2. Manganese

The correlation of Sb with Mn was unexpected despite the testified presence of concretions of pyrolusite (blackish powder) in the area. Pyrolusite typically occurs in oxidation zones, mainly in sedimentary beds, and in Tuscany, its occurrence is typically related to supergene alteration. In the chalcedonies, manganese contents (324–655 ppm) are lower than those reported for the UCC (0.1 wt %; Rudnick & Gao, 2003). In addition, they are often correlated with Be, which also shows high contents (22–48 ppm) with respect to both the continental crust and the chondrite composition. Be enrichment is frequent in jasperoids (e.g. Johnson, 1977; Lovering & Heyl, 1980; McLemore, 2002, 2010), and beryllium deposits often contain concentrations of manganese oxides together with fluorite, lithium, zinc, uranium and several trace elements (Lindsey, 1977). Consequently, while the observed correlation between Mn and Be is not surprising, it appears more difficult to trace the origin of these two elements. The rocks investigated in this study bear low levels of Mn, except for the carbonates and the A1 and A2 metapelites. Therefore, it seems plausible to hypothesize that the high Mn contents of the chalcedonies mainly derive from the leaching of underlying carbonates. As regards Be enrichment, none of the surrounding rocks shows high levels of this element. However, no data are available on Be amounts contained by the Roccastrada rhyolites, which, in turn, could represent likely candidates. The conduits provided by faults should have then allowed beryllium to circulate, and a wall-rock reaction may have been essential for mineralization processes.

A not compelling but useful example to remember in this regard is the one offered by Lindsey (1977) for Spor Mountains and Honeycomb Hills (Utah). That is to say, carbonate-hosted beryllium (replacement) deposits genetically related to rhyolites (the only ones to return Mn-rich red beryl) and associated with hydrothermal silica, carbonates and other minerals. Moreover, it can be worth remembering that Be-bearing minerals such as bertrandite mainly occur in carbonate replacement deposits such as the jasperoids.

6.d.3. Arsenic

Arsenic enrichment has been frequently used to trace either epithermal gold mineralization (Henley, 1985; Berger & Silberman, 1985; Silberman & Berger, 1985) or geothermal activity (Bingqiu et. al. 1986). Geothermal waters and sinters frequently bear high amounts of As, besides Sb and Hg (Weissberg, 1969; Boyle & Jonasson, 1973). Furthermore, this element can be concentrated in hydrothermally altered metasediments (see, e.g. Codeço et al. 2021) and in surficial conditions. In the latter environment, As may be released by Fe-hydroxides under oxidizing acid conditions which, in turn, can be obtained owing to sulfide decomposition (see e.g. Craw et al. 2000). In the investigated rocks, the highest concentrations of As (17–43 ppm) were measured in the metamorphic rocks (samples V2, V6 and V8). Consequently, the As in the chalcedonies may have been derived from magmatic/hydrothermal fluids and As-enriched meteoric waters.

6.d.4. Caesium

Cs, Rb and K (LFS elements) show a consistent correlation in several samples. In hydrothermal fluids, the alkali metals such as Rb and Cs remain in solution and are later incorporated into other phases such as micas (see, e.g. Codeço et al. 2021). Accordingly, the metamorphic rocks, especially the metapelite A2, turned out to be the most Cs-rich rocks among those examined. The derivation of Cs from water–rock interaction with these rocks thus appears as plausible.

6.e. Gold and iron

As specified above, while the absolute quantities have to be considered with caution, the presence of Au in chalcedonies is evident and not surprising. Au leaching and transport require suitable redox conditions as well as the availability of suitable complexing agents (Mann, 1984 ; Webster & Mann, 1984; Stoffregen, 1986; Bowell et al. 1993; Freyssinet et al. 2005). It is commonly believed that, at low temperatures (<400 °C), Au covalently bonds with bisulfide ligand (Wood et al. 1987) and is then sequestered by sulfides (Cygan & Candela, 1995; Jugo et al. 1999; Simon et al. 2000) while, at high temperatures (>350–400 °C), gold tends to form chloride complexes (Candela, 2003). In the first case, the precipitation of gold can occur from the ore-fluids, for example, owing to fluid dilution or fluid–rock interactions, or from the achievement of oxidizing conditions that cause the breakdown of the sulfide complexes (Ilchik & Barton, 1997). In the second case, however, it is cooling that favours the precipitation of gold from gold-chloride solutions (Gammons & Williams-Jones, 1997).

In this regard, it is also worth mentioning the investigation by Williams-Jones & Normand (1997), according to whom Sb and Au jointly reach maximum concentrations when antimony is transported as HSb₂S₄ ⁻. The same authors highlight how evident Sb–Au associations are reached when the transport takes place in an alkaline environment (such as limestones) and precipitation occurs in an environment suitable for the acidification of the waters (such as phyllosilicate-rich rocks). In this way, hydrogen ions are produced and can promote the deposition of gold and stibnite. Undoubtedly other processes are possible, but the similarity of the conditions described by Williams-Jones & Normand (1997) to those provided by the area under investigation make this reconstruction one of the most likely.

On the other hand, the absence of correlation of Au with iron seems to diminish (but not exclude) the possible importance of an association with Fe-oxyhydroxide, as often reported in the literature (Karasyova et al. 1998; Ran et al. 2002; Cohen & Waite, 2004; Corral et al. 2018). Moreover, in the territory under examination, the lack of a correlation between stibnite and pyrite occurrences is observed. As explained by Dessau (1952), ‘pyrite and stibnite were deposited at the footwall and at the roof of the limestone, respectively, due to the chemical action, together with the drop in temperature and pressure, or due to a chronological shift for which pyrite should have been deposited in a “slightly earlier, higher temperature period”’. In a nutshell, the precipitation of pyrite and Sb–Au should be in succession. While the former lowers the amount of reduced sulfur in the ore fluid, the latter is favoured by a lowering of temperatures able to decrease Sb solubility and lead to Sb–Au precipitation. Finally, it may be interesting to recall the study performed by Nelson (1990), which reports the range of Au concentration in jasperoids (up to 94 ppm) in US Carlin-type orebodies and traces the origin of gold to marine black-shales. Metamorphic rocks and chalcedonies are those bearing discrete Au amounts.

6.f. Textures

The various growth and textures observed in chalcedony testify to the physicochemical variability of the mineralizing solutions, chiefly in terms of SiO₂ concentration and pH (Richter et al. 2015). On the other hand, the feathery quartz (jigsaw-puzzle texture) observed in PP8 indicates recrystallization. Dong et al. (1995) established that this texture forms at quartz stability temperatures (>180 °C) from the transformation of pre-existing silica phases such as amorphous silica or chalcedony (Lovering, 1972).

The presence of calcite layers in PP4 may indicate a process involving boiling, with consequent loss of CO₂, generation of CO₃ ²⁻ (from dissociation of HCO₃ ⁻) and precipitation of calcite (Henley & Brown, 1985; Reed & Spycher, 1985). However, the analysis of fluid inclusions performed on other agates from Poggio Peloso by Brogi & Fulignati (2012) did not retrieve evidence of boiling. Otherwise, calcite may have been generated by contact of cold fluids with hot rocks (retrograde solubility; see Simmons & Christenson, 1994).

7. Conclusions

The chalcedonies of Poggio Peloso are the first Sb-rich chalcedonies reported in the literature. Not only is there no comparison among chalcedony and agate analysed to date, but not even the rocks present in the surrounding area have comparably high Sb levels. Antimony is strongly correlated with B, Mn, As and Cs in these chalcedonies. While the high presence of B and Mn may indicate the role of the fluids that, leaching the hosting carbonates, led to their replacement with chalcedony, As and Cs can be better associated with fluids interacting with metamorphic rocks.

The fluid transport in an alkaline environment agrees with the presence of carbonates, whereas silica precipitation in such an environment is possibly confirmed by features such as the abundance of barite inclusions. The latter may be further correlated with and explain the positive Eu anomaly observed in these chalcedonies (as well as in the Palaeozoic metamorphic rocks A1 and A2 from the Monte Amiata geothermal field).

The alkaline environment may also be indicated as responsible for the Sb–Au association, while the acidification of the waters promoted by metamorphic rocks or volatile fluids (HF) and SiF₄, GeF₄ and BF₃ compounds (related to the abundance of Ge) should have favoured their precipitation.

Among the acidic components carried by the fluids, silicic acid must be included. For neither the latter nor any other constituent of these chalcedonies was it possible to draw a convincing correlation with the surrounding volcanic rocks. The comparison established on the basis of the REEs drew significant analogies only with the carbonate (B and Mn seen above) and metamorphic rocks. Among the latter, the Palaeozoic rocks (graphitic quartz-metasandstone and metasiltite and a graphitic phyllite) from two geothermal wells of the Monte Amiata geothermal field provided the most convincing comparison, also regarding the content of gold. In addition, some features of the geothermal fluids of this field (i.e. significant concentration of Sb, B and SiO₂; positive Eu anomaly) seem to be similar to those of the fluid that deposited the chalcedonies. Thus, the latter fluid likely circulated in and reacted with metamorphic rocks analogous to those hosting the deep reservoir of the Monte Amiata geothermal field.

As for the deposition temperatures, it is possible to assume that they did not exceed the threshold of 225 °C, which was set as the upper term for the deposition of the Poggio Peloso quartz. In any case, the textures of the chalcedonies testify to the physicochemical variability of the mineralizing solutions and different precipitation. Therefore, they are representative of a diachronic process in which fluids with different characteristics have formed different generations of the silica phases.