Mafic VMS Mineralization in the Mesozoic Metavolcanic Rocks of the Evros Ophiolite, Xylagani Area, Greece

Abstract

The sulfide mineralization at Xylagani is hosted in metamorphosed mafic massive and pillow lava. It has an Early–Middle Jurassic age and belongs to the Makri unit, which represents the upper crustal section of the Evros ophiolite in the Circum Rhodope Belt, Northern Greece. The protolith of the host rock is basalt that has a boninitic-to-low-Ti tholeiitic composition and was formed in an intra-oceanic supra-subduction zone within a juvenile forearc-to-volcanic arc setting. The volcanic rocks were subjected to ocean-floor metamorphism at very low-grade prehnite–pumpellyite facies and low-grade greenschist facies at temperatures of up to 360 °C and pressures between 1 and 4 kbar. The mineralization shows typical features of a stratabound–stratiform deposit and occurs as silicified lenses and layers with disseminated and massive sulfides and gold. Based on host rock composition, geotectonic setting, and base metal content, the mineralization at Xylagani is classified as a Cu-rich mafic volcanic-associated deposit, i.e., Cyprus-type VMS (volcanogenic massive sulfide). The mineralization consists of pyrite, chalcopyrite, gold, pyrrhotite, sphalerite, galena, and tennantite-(Zn). It was formed at a subseafloor setting where hydrothermal fluids circulated through the host volcanic rocks, resulting in a pervasive alteration (silicification and chloritization) and the development of a replacement VMS deposit. The very low-to-low-grade orogenic metamorphism and related deformation during the Alpine collision in the Middle Jurassic to Early Cretaceous periods remobilized the mineralization and formed milky quartz veins with rare sulfides, crosscutting the metavolcanic rocks.

1. Introduction

The Hellenides are part of the Alpine orogeny and contain numerous ore deposits that belong to the Tethyan metallogenic belt. This belt extends from the Iberian Peninsula and the Pyrenees along the Alps, the Dinarides, and the Carpathian–Balkan region, reaching through Anatolia and the Caucasus to Iran, the Tibetan Plateau, and Indonesia. Metallogenesis in this belt resulted from the formation and destruction of the Tethys oceanic lithosphere between the southern margin of Eurasia and the continental plates of Africa, Arabia, and India [1,2,3].

This geodynamic model is reflected in the geotectonic zones of Greece and includes successive events of oceanic rifting, extension, subduction, and collision between continental and oceanic lithospheric plates, parts of which constitute the geology of the Aegean region today. As a consequence, these dynamic events have influenced metallogenesis in the Hellenides from the Triassic period to the present, spanning the last 250 million years and giving rise to significant metallogenic provinces. These metallogenic provinces in Greece have a defined geographical distribution and similar ages and fall into two major categories: (i) Mesozoic age chromite, volcanogenic massive sulfide, and lateritic nickel mineralization associated with ophiolitic sequences; (ii) Cenozoic age mineralization (e.g., intrusion-related, skarn, carbonate replacement, porphyry, epithermal, and vein-type deposits) with precious, base, and rare metals, related to the geodynamic regime of oceanic (and continental) crust subduction beneath the continental lithosphere [4,5,6,7,8].

The mineralization examined here is located near the village of Nea Petra, 5 km northeast of the Xylagani town in the Thrace region of Greece and has not been explored in detail in the past (Figure 1 and Figure 2). This contribution focuses on the host rock lithology and geochemistry, ore textures, sulfide mineralogy, mineral chemistry, geochemical characterization, and sulfur isotopic analysis of the mineralization, aiming to assess the ore deposit type and provide data for metallogenic interpretation. The purpose of this work is to document the paragenetic evolution of sulfide minerals to determine the geotectonic environment and the sulfur sources and to suggest a model for the evolution of mineralization for a better understanding of the ore formation mechanism.

2. Geological Setting

The Triassic–Jurassic ophiolites of the Hellenides in the northern Aegean are considered remnants of the Neotethys Vardar–Axios ocean and represent tectonic sutures formed during the progressive closure of the oceanic basin [9,10,11]. The eastern Vardar suture zone of Hellenides, known as the Circum Rhodope belt (CRB), consists of major units that tectonically overlie the crystalline basement of the Serbo-Macedonian and the Rhodope massifs (inset of Figure 1). In Thrace of Greece, CRB is characterized by low-grade metamorphic rocks of Mesozoic age, which form the upper tectono-stratigraphic unit of the Rhodope massif and crop out along the hanging wall of the extensional detachment faults (Figure 1) that were active during the Paleogene period. The rocks of the CRB in Thrace are subdivided into the Makri unit and the Melia formation [9,12,13,14].

The Makri unit has a Middle Triassic-to-Middle Jurassic age and overlies tectonically the crystalline basement of the Rhodope massif along low-angle faults, which represent reworked thrust contacts that were later activated as extensional detachments (Figure 1; [14,15]). It is composed of a lower metasedimentary series and an upper metavolcanic series [9,12,13,14,15,16]. The metasedimentary series includes chlorite–sericite schists, calc-schists, and phyllites, interbedded with minor clastic rocks (e.g., meta-greywackes and meta-quartzites) and abundant crystalline limestone–dolomite or marble beds of variable thickness. The sedimentary protoliths were deposited in a shallow marine platform-type deposit. The metamorphic grade ranges from very low to low grade, up to greenschist facies.

Figure 1. Simplified geological map of the eastern Rhodope in the Greek Thrace region showing the Makri unit, the Evros ophiolite, and the Melia formation from the Circum Rhodope belt (modified from [9,12,13,16,17,18]). Inset map: map showing the main geotectonic zones of Greece, Western Turkey, and the area of the geological map (modified from [14]).

Figure 1. Simplified geological map of the eastern Rhodope in the Greek Thrace region showing the Makri unit, the Evros ophiolite, and the Melia formation from the Circum Rhodope belt (modified from [9,12,13,16,17,18]). Inset map: map showing the main geotectonic zones of Greece, Western Turkey, and the area of the geological map (modified from [14]).

Figure 2. Geological sketch map of the Xylagani area, showing the mineralization and the inactive mines along Mylorema Creek, near the village of Nea Petra (after [12]). Inset map: map of NE Greece showing the studied area.

Figure 2. Geological sketch map of the Xylagani area, showing the mineralization and the inactive mines along Mylorema Creek, near the village of Nea Petra (after [12]). Inset map: map of NE Greece showing the studied area.

The metavolcanic series of the Makri unit conformably overlies metasedimentary rocks. It consists of greenschists (chlorite-, sericite-, tremolite-, actinolite-, epidote-schists) that are mainly derived from mafic pillow, massive or brecciated mafic-to-acid lava, and pyroclastics, with occasional small serpentinite bodies [15,16]. The metavolcanic and meta-pyroclastic rocks were metamorphosed from prehnite–pumpellyite to greenschist facies at temperatures of 220 to 400 °C and pressures of 1 to 4 kbar and belong to the upper crustal section of the Evros ophiolite [11,16,18].

The Evros ophiolite includes mafic volcanic and intrusive magmatic rocks that are associated with an intra-oceanic supra-subduction zone with a forearc-to-arc signature. They are characterized by boninitic and tholeiitic to calc-alkaline affinities and date to the Early–Middle Jurassic [9,11,14,15,16,18]. However, an input of a depleted MORB-type mantle source, with a contribution of crustal contamination and/or recycled sediments, has been proposed by [11]. The rocks of the Evros ophiolite comprise unmetamorphosed or weakly metamorphosed plagiogranite, gabbro, metadiabases, pillow basalt lava, boninitic–tholeiitic massive basalt, basaltic andesite to andesite lava, and boninitic–tholeiitic dikes [11,16,17,18,19].

Aliki limestone, with an Early Cretaceous age, lies unconformably at the top of the greenschists of the Makri unit [9,15]. The Melia formation includes metamorphosed volcanic and sedimentary rocks of the Middle Jurassic to Early Cretaceous age and is considered to have been overthrusted on the Makri unit [14,20,21]. The protoliths of these rocks include pillow lava, massive mafic lava, and sheeted dykes placed in a Jurassic-to-Cretaceous sedimentary series [15,17]. The sediments were deposited during a rapid uplift of the Rhodope metamorphic complex [14].

3. Sampling and Analytical Methods

Sixty-eight samples of fresh-to-slightly altered host rocks and the mineralization outcrops were collected along Mylorema Creek and from underground galleries. Twenty polished-thin sections and thirty-three polished sections from selected samples were studied under a ZEISS Axioskop 40 dual transmitted–reflected polarizing light microscope (ZEISS, Jena, Germany) and a JEOL JSM-6390LV scanning electron microscope (SEM) (JEOL, Tokyo, Japan) equipped with an OXFORD INCA 300 energy-dispersive system (EDS) (Oxford Instruments Ltd., Abingdon, UK) at the School of Geology, Aristotle University of Thessaloniki. For the back-scattered electron (BSE) mode, the operating conditions were a 20 kV accelerating voltage and 0.4 mA probe current, 80 s analysis time, and a beam diameter of 1 μm.

Twenty-eight ore samples were analyzed by atomic absorption spectrometry (AAS) at the analytical laboratory of the Institute for Geology and Mineral Exploration, Xanthi, Greece. Six samples of the host rocks were analyzed by X-ray fluorescence (XRF) for major and trace elements and by inductively coupled plasma atomic emission spectroscopy (ICP-AES) for rare earth elements (REEs) at the Institute of Mineralogy and Petrology, University of Cologne, Germany. In order to determine trace metal concentrations, pure pyrite separates without any inclusions of other sulfides was extracted and was analyzed by ICP-AES at Svensk Grundämnes-analys AB in Luleå, Sweden. Neutron activation analyses (NAAs) for gold on mineralized samples were carried out at the Max-Planck Institute, Heidelberg, Germany.

The chemical composition of the ore minerals was determined by a Cameca Camebax Microbeam electron microprobe at the Institute of Mineralogy and Petrology, University of Cologne, Germany. Operating conditions were 20 kV and 20 nA, with a beam diameter <1 μm. The following analytical wavelengths and standards were employed: SKα and FeKα: FeS₂; CuKa: CuFeS₂; PbMa and SeLa: PbSe; ZnKa: ZnS; AgLa: Ag₂S; BiLa and TeLa: BiTe; MnKa: MnTi, CdLa: CdSe, HgMa: HgTe; SbLa, AsLa, CoKa, NiKa, AuMa and PtMa: pure metals. Count times depended on the expected element concentration and ranged from 20 to 60 s. The detection limits were as follows: Ag: 0.01 wt%, As: 0.01 wt%, Au: 0.01 wt%, Bi: 0.01 wt%, Cd: 0.01 wt%, Co: 0.01 wt%, Cu: 0.01 wt%, Fe: 0.02 wt%, Mn: 0.01 wt%, Ni: 0.01 wt%, Pb: 0.01 wt%, S: 0.02 wt%, Sb: 0.01 wt%, Se: 0.01 wt%, Te: 0.01 wt%, and Zn: 0.01 wt%.

Sulfur isotope measurements were carried out on five pure pyrite separates, in which pyrite contained very small amounts of impurities (mainly chalcopyrite), less than 5% in all concentrates. Isotope data from the other mineralization types were not obtained due to the high contamination (>10%) of the other sulfides intergrown with pyrite. The isotopic analyses were carried out at the Institute of Geochemistry, University of Göttingen, Germany.

4. Results

4.1. Petrography and Geochemistry of Metavolcanic Rocks

The sulfide mineralization in the Xylagani area is hosted in metamorphosed mafic lava and pyroclastics of Jurassic age, which consist of chlorite-, sericite-, tremolite-, actinolite-, and epidote-greenschists (Figure 3a–c). They belong to the Makri unit, which represents the upper crustal section of the Evros ophiolite in the CRB at Thrace.

These rocks are foliated on a macro-to-microscopic scale and are sometimes extensively altered, especially when they are associated with mineralization. They often show pillow structures (Figure 3a), while primary textures, such as porphyritic, subophitic, or amygdaloidal textures (Figure 3b,c), are well preserved. Based on textural and mineralogical features, the metavolcanic rocks can be subdivided into pyroxene-phyric lava, aphyric-oligophyric lava, and albite-rich lava.

The secondary minerals observed in metavolcanic rocks, such as quartz, chlorite, sericite, prehnite, pumpellyite, epidote, zoisite, titanite, actinolite, and calcite, are characteristic of very low-to-low-grade metamorphism, from prehnite–pumpellyite to greenschist facies. Frequently, volcanic rocks are rich in vesicles filled with quartz, prehnite, pumpellyite, chlorite, and calcite, forming characteristic amygdaloidal textures (Figure 3d–f). Silicification is the main alteration feature resulting from the interaction of hydrothermal solutions with lava during the ore-forming process. In addition to silicification, the devitrification and chloritization of volcanic glass also occur. Sporadically, massive milky quartz veins were observed, crosscutting the metavolcanic rocks and their foliation. These veins range in width from 0.5 to 5 cm and are believed to have formed during the orogenic metamorphism that accompanied the Alpine Orogeny.

Major and trace elements were determined for six representative samples of the metavolcanic rocks that are related with the mineralization, and the results are presented in

Table 1. Chemical analyses using XRF and ICP-AES of the metavolcanic rocks in the Xylagani area. dl: detection limit; bdl: below the detection limit, Fe₂O₃: Fe total; LOI: lost on ignition material at 1100 °C.

	dl	NP2	NP11a	NP11b	NP19	NP36a	NP36b
wt%
SiO2	0.01	57.78	48.60	49.80	54.60	51.01	59.00
TiO2	0.01	0.32	0.32	0.53	0.51	0.33	0.28
Al2O3	0.01	13.23	12.20	17.10	14.00	13.55	13.50
Fe2O3	0.01	7.21	9.45	9.22	8.98	8.60	6.51
MgO	0.01	7.64	13.89	9.05	7.93	10.76	6.46
CaO	0.01	5.46	7.78	5.94	5.99	8.86	4.09
Na2O	0.01	4.59	1.76	1.07	0.25	2.23	4.65
K2O	0.01	0.64	0.56	1.47	1.56	0.42	0.02
MnO	0.01	0.09	0.19	0.09	0.14	0.15	0.08
P2O5	0.01	0.01	0.03	0.02	0.18	0.02	0.03
LOI		3.15	4.99	4.89	4.77	4.67	5.91
Tot		100.12	99.77	99.18	98.91	100.60	100.53
ppm
As	100	bdl	bdl	bdl	bdl	bdl	bdl
Ba	2	824	2032	2164	3175	bdl	bdl
Co	2	30	36	51	32	36	20
Cr	2	240	36	66	104	32	29
Cu	1	57	68	69	1090	15	39
Ga	1	13	18	19	16	9	5
Nb	2	bdl	bdl	bdl	bdl	bdl	bdl
Ni	2	43	56	66	105	37	41
Pb	5	6	8	16	10	8	10
Rb	1	bdl	19	20	18	bdl	bdl
Sc	1	33	31	32	31	32	31
Sr	1	110	230	235	200	45	34
Th	1	bdl	bdl	bdl	bdl	bdl	bdl
U	1	bdl	bdl	9	8	bdl	bdl
V	1	201	265	272	180	231	177
Y	1	13.0	13.5	13.8	11.8	13.0	9.7
Zn	1	50	120	130	4145	40	43
Zr	2	17	18	20	19	20	18
REE (ppm)
La	0.01	-	-	1.46	4.74	-	1.04
Ce	0.01	-	-	3.15	6.19	-	2.26
Nd	0.01	-	-	3.06	7.94	-	1.69
Sm	0.01	-	-	0.99	2.30	-	0.50
Eu	0.01	-	-	0.51	0.90	-	0.18
Gd	0.01	-	-	1.37	4.32	-	0.79
Dy	0.01	-	-	2.04	5.37	-	1.37
Er	0.01	-	-	1.59	3.87	-	1.18
Yb	0.01	-	-	1.48	3.23	-	1.20
Lu	0.01	-	-	0.21	0.48	-	0.17
ΣREE		-	-	15.86	39.34	-	10.38
Al2O3/TiO2		41.34	38.13	32.26	27.45	41.06	48.21
Zr/Y		1.31	1.33	1.45	0.45	1.54	1.03
Zr/TiO2		0.01	0.01	0.00	0.00	0.01	0.00

. Concentrations of SiO₂ (48.60–59.00 wt%), Na₂O+K₂O (1.81–5.23 wt%), CaO (4.09–8.86 wt%), MgO (6.46–13.89 wt%), and Fe₂O₃ (6.51–9.45 wt%) show a broad distribution, while TiO₂ (0.28–0.53 wt%), Zr (17–20 ppm) and Y (9.7–13.8 ppm) are low.

The metavolcanic rocks of Xylagani are characterized by relatively high contents of Cu (15–1090 ppm), Zn (40–4145 ppm), Cr (29–240 ppm), V (177–272 ppm), and Ni (37–105 ppm). The abundance of rare earth elements (REEs) (Table 1) and chondrite-normalized patterns of the metavolcanic rocks (Figure 4) have also been used to interpret their eruptive setting. The REE concentrations are low (2 to 11 times chondritic values), with their low total content (ΣREE) ranging from 10.38 to 39.34 ppm.

4.2. Types of Mineralization

Field observations show that the mineralization at Xylagani, which is concentrated mainly around Mylorema Creek near the Nea Petra village (Figure 2), occurs in the form of silicified bodies hosted in the metavolcanic rocks. Over 20 ore outcrops occur in the area, and local residents report that this mineralization was exploited by a French–Italian company between 1900 and 1910. The intense mining activity in the area is witnessed by 23 inactive underground galleries and surface workings, along with stockpiles of mined material (Figure 5a,b). These works were probably exploiting pyrite and chalcopyrite ore. The galleries, which have a length of up to 120 m, are flooded, and entry to most of them is rather difficult. Remnants of mining installations or ore concentration facilities have not been found in the region, indicating that the ore was likely transported out of the mined area.

The main ore bodies are up to 100 m long and more than 10 m thick and cover an area of ~2 km². They form lenses or layers, which are conformable with the foliation of the associated metavolcanic rocks at different stratigraphic levels and exhibit a stratabound or stratiform structure. The mineralization is disseminated at the edge and is massive compared to the core of the ore bodies. Disseminated mineralization also occurs in the host rocks close to the ore bodies. Stockworks of stringer-type ore were not observed in the area.

Based on textural features, five types of mineralization are recognized: (1) Thin layers, consisting mainly of pyrite, aligned parallel to the bedding and occasionally folded (Figure 5c) with thicknesses of 0.5 to 3 mm. (2) All the studied silicified ore bodies contain disseminated sulfides (<10–20%), mainly pyrite. Significant disseminated sulfide enrichment also occurs in altered metavolcanics. This is the most widespread type of mineralization. (3) Semi-massive mineralization occurs in the central parts of lenses and contains about 50 to 80% sulfides, mainly pyrite and chalcopyrite (Figure 5d). (4) Massive mineralization is characterized by a very high sulfide content (80%) and is subdivided into two types: massive pyrite and massive chalcopyrite–pyrite mineralization. Massive pyrite occurs as lenses, up to 30 cm long and 10 cm thick, forming stratabound mineralization (Figure 5e) or as layers, extending up to 10 m in length and 5 cm in thickness. Massive chalcopyrite–pyrite mineralization occurs as layers (Figure 5f) with a length of up to 10 m and a thickness of up to 10 cm, forming stratiform mineralization. (5) Milky quartz veins have a low sulfide content, crosscutting metavolcanic rocks and their foliation.

Chemical analyses from the different mineralization types in the Xylagani area (

Table 2. Average AAS chemical analyses of the ore types from the Xylagani area. (1) Thin layers, (2) disseminated, (3) semi-massive, (4) massive pyrite, and (5) massive chalcopyrite–pyrite mineralization. dl: detection limit; bdl: below the detection limit; n: number of analyses.

	dl	1 (n = 3)	2 (n = 13)	3 (n = 6)	4 (n = 3)	5 (n = 3)
wt%
Fe	1	8.40	12.59	25.70	43.65	30.95
ppm
Au	0.1	2.0	3.6	5.0	3.6	0.1
Ag	1	8.3	1.0	3.0	bdl	54
Cu	1	144	3548	16,717	154	189,500
Zn	1	377	99	848	bdl	1010
Pb	1	243	28	73	80	5115
As	10	93	bdl	216	248	29
Co	1	333	68	36	73	2
Ni	1	11	13	8	11	9

; Supplementary Table S1) showed that the studied ore bodies demonstrate a Cu concentration ranging from 132 ppm to 19.00 wt%. In contrast, Zn and Pb are lower, with contents reaching 4290 ppm and 5190 ppm, respectively. The contents of As, Co, and Ni in the mineralized bodies are relatively low, up to 248 ppm, 333 ppm, and 13 ppm, respectively. The ore bodies contain values of Au between 0.1 and 6.4 ppm, with an average of 3.5 ppm (n = 28). The highest average Au concentration (5 ppm) is found in semi-massive mineralization. In addition, the concentration of Ag reaches 54 ppm, with the highest content in massive chalcopyrite–pyrite mineralization.

4.3. Ore Mineralogy and Alteration Minerals

The main ore minerals in the mineralization at Xylagani are pyrite and chalcopyrite, with sphalerite as the minor phase. Pyrrhotite, galena, tennantite-(Zn), and gold occur in traces (Figure 6). Secondary iron and copper minerals, such as goethite, malachite, and covellite, are rare. Quartz, sericite, and chlorite are the most common alteration minerals. The ore mineral assemblages, characterizing the different mineralization types, are presented in

Table 3. Ore mineral assemblages characterizing the different types of mineralization in the Xylagani area (au: gold; ccp: chalcopyrite; gn: galena; po: pyrrhotite; py: pyrite; sp: sphalerite; tn: tennantite-(Zn)).

Ore types	Main Ore Minerals	Assemblages of ore Minerals
Thin layers	pyrite–chalcopyrite	py I, py II, ccp I, sp I
Disseminated	pyrite–chalcopyrite	py I, ccp I, py II, sp I, ccp II, sp II, po, tn I
Semi-massive	pyrite–chalcopyrite	pyI, au, ccp I, py II, sp I, ccp II, sp II, tn I
Massive pyrite	pyrite	py II
Massive chalcopyrite–pyrite	chalcopyrite–pyrite	py II, ccp II, sp II, gn I
Milky quartz veins	pyrite–chalcopyrite	py III, ccp III, sp III, gn II, tn II

, and the microprobe analyses of the ore minerals are represented in

Table 4. Average microprobe analyses in wt% of framboidal fined-grained pyrite (py I), coarse-grained pyrite (py II), pyrite in milky quartz veins (py III), the second generation of chalcopyrite (ccp II), chalcopyrite in milky quartz veins (ccp III), pyrrhotite (po), sphalerite (sp), galena from massive chalcopyrite–pyrite mineralization (gn I), galena from milky quartz veins (gn II) and tennantite-(Zn) (tn I and tn II)) (n: number of analyses; bdl: below the detection limit; -: not analyzed).

	py I	py II	py III	ccp II	ccp III	po	sp I,II,II	gn I	gn II	tn I	tn II
	(n = 6)	(n = 21)	(n = 6)	(n = 14)	(n = 14)	(n = 3)	(n = 16)	(n = 3)	(n = 8)	(n = 9)	(n = 9)
Bi	-	-	-	-	-	-	0.04	bdl	bdl	bdl	bdl
Pb	-	-	-	-	-	-	0.02	85.10	86.63	bdl	bdl
Au	bdl	0.06	0.02	0.06	bdl	bdl	-	-	-	-	-
Te	bdl	0.02	0.01	0.03	0.02	bdl	0.01	0.07	0.03	bdl	bdl
Sb	-	-	-	-	-	-	0.01	bdl	0.01	4.21	4.55
Cd	-	-	-	-	-	-	0.14	0.04	0.02	0.07	0.13
Ag	bdl	0.02	0.01	0.05	0.02	0.04	0.03	0.17	0.08	0.15	0.40
Se	bdl	0.05	0.04	0.05	0.16	0.02	0.01	1.47	0.18	0.09	0.09
As	0.33	0.14	0.06	0.07	0.02	0.08	bdl	0.01	0.02	17.71	17.64
Zn	-	-	-	-	-	-	65.19	0.06	0.03	6.13	7.16
Cu	0.01	0.13	0.10	34.33	34.68	0.02	0.03	0.07	0.05	41.61	40.61
Ni	0.03	0.03	0.03	0.02	0.01	0.06	0.01	0.02	0.01	0.02	0.01
Co	0.02	0.08	0.04	0.03	0.02	0.03	-	-	-	-	-
Fe	46.45	46.37	46.56	30.47	30.40	60.76	1.38	0.05	0.07	2.23	1.91
V	bdl	0.01	bdl	-	-	-	-	-	-	-	-
Mn	-	-	-	-	-	-	0.03	bdl	0.01	0.01	0.01
S	52.91	53.67	53.90	35.00	35.22	38.79	32.94	12.87	13.41	27.63	27.63
Total	99.75	100.58	100.76	100.11	100.61	99.80	99.84	99.93	100.55	99.84	100.15
	Formula units calculated on the basis of total atoms.
Pb	-	-	-	-	-	-	-	0.99	1.00	-	-
Sb	-	-	-	-	-	-	-	-	-	0.53	0.56
Cd	-	-	-	-	-	-	-	-	-	0.01	0.02
Ag	-	-	-	-	-	-	-	-	-	0.02	0.06
Se	-	-	-	-	-	-	-	0.05	0.01	0.01	0.02
As	0.01	-	-	-	-	-	-	-	-	3.56	3.55
Zn	-	-	-	-	-	-	0.97	-	-	1.41	1.65
Cu	-	-	-	1.00	1.00			-	-	9.87	9.64
Fe	1.00	1.00	1.00	1.00	0.99	0.90	0.03	-	-	0.60	0.51
S	1.99	2.00	2.00	2.00	2.01	1.00	1.00	0.96	0.99	13.00	12.99
Atoms	3.00	3.00	3.00	4.00	4.00	1.90	2.00	2.00	2.00	29.00	29.00

and Supplementary Table S2.

Pyrite is the predominant sulfide of all ore types during mineralization and occurs in three generations (py I, py II, py III). Framboidal and fine-grained pyrite was found in the thin layers and in the disseminated and semi-massive ore types and represents the early generation (py I); it does not contain inclusions of other sulfides. The framboids (Figure 7a) are found either isolated or in clusters within fine-grained quartz. Thin layers consisting of framboidal pyrite, parallel to the host rock bedding with a thickness of about 100 μm, have also been observed in a matrix of hydrothermal quartz and chlorite (Figure 7b,c). The framboids are spherical and, more rarely, oval-shaped, ranging from 3 to 70 μm across. In some cases, they display a concentric internal arrangement of microcrystals. Microprobe analyses of py I (Table 4; Supplementary Table S2) showed As contents ranging from 0.07 to 0.91 wt% (0.33 wt% average), corresponding to a chemical formula of Fe_1.00(S,As)_2.00.

Following the deposition of framboidal and fine-grained pyrite, the second generation of pyrite (py II) displays coarse-grained euhedral and subhedral textures with various sizes of up to 5 mm. Frequently, a growth of py II in the form of radiating, subhedral pyrite and thin pyritic zones occurs around framboids (Figure 7d). This second pyrite generation also forms veinlets crosscutting the framboidal layers. Within some pyrite grains, well-preserved outlines of framboidal pyrite (Figure 7d) suggest that coarse pyrite (py II) postdates the formation of py I. The second pyrite generation is characterized by inclusions of other sulfides, such as chalcopyrite and sphalerite, and less abundant pyrrhotite, tennantite-(Zn), and gold (Figure 7e). These inclusions range between 5 and 200 μm across and are rounded, oval-shaped, or irregular. The cataclastic deformation of pyrite is common, with chalcopyrite and quartz filling the cracks and fractures of pyrite (Figure 7f). Quartz fibers are usually developed in pressure fringes around euhedral pyrite (py II) (Figure 7g).

The second generation of pyrite (py II) shows lower As contents than py I and up to 0.37 wt% (0.09 wt% on average). Gold reaches 0.17 wt% (0.06 wt% on average). The average chemical formula of py II corresponds to Fe_1.00S_2.00. The Co and Ni concentrations of pure pyrite (py II) from the massive pyrite ore type (

Table 5. ICP-AES analyses of pure pyrite from massive pyrite ore-type mineralization. dl: detection limit; bdl: below the detection limit.

	dl	NP 48-2a	NP 48-2b	NP 48-2c	Average
wt%
Fe	0.01	48.1	48.0	48.3	48.1
Al	0.01	0.01	0.01	0.01	0.01
Ca	0.01	0.01	0.01	0.01	0.01
K	0.01	bdl	bdl	bdl	bdl
Mg	0.01	0.02	0.02	0.02	0.02
Na	0.01	0.01	0.01	0.01	0.01
P	0.01	0.02	0.02	0.02	0.02
Ti	0.01	bdl	bdl	bdl	bdl
ppm
Ag	2	<2	<2	<2	<2
As	20	293	291	299	293
Au	1	1.8	1.7	1.9	1.8
Ba	2	5	5	5	5
Be	2	<2	<2	<2	<2
Co	2	59	63	70	63
Cr	2	10	11	14	11
Cu	2	338	297	262	297
La	10	<10	<10	<10	<10
Mn	2	31	30	31	31
Mo	10	<10	<10	<10	<10
Nb	10	13	11	10	11
Ni	2	19	16	18	18
Pb	10	53	51	53	53
Sc	2	<2	<2	<2	<2
Sn	10	<10	<10	<10	<10
Sr	2	<2	<2	<2	<2
Zn	2	<2	<2	<2	<2
Zr	2	<2	<2	<2	<2
V	2	24	22	24	24
W	10	<10	<10	<10	<10
Y	2	<2	<2	<2	<2
Co/Ni		3.11	3.94	3.89	3.56

), ranged from 59 to 70 ppm and from 16 to 19 ppm, respectively. The Co/Ni ratios range from 3.11 to 3.94. Arsenic varies from 291 to 299 ppm, Cu varies from 262 to 338 ppm, Pb varies from 51 to 53 ppm, and Au varies from 1.7 to 1.9 ppm. The third generation of pyrite (py III) contains even less As than the older generations, up to 0.16 wt% (0.06 wt% on average). Gold is also lower, reaching 0.06 wt%. The average chemical formula of py III is Fe_1.00S_2.00.

The most abundant sulfide after pyrite is chalcopyrite, which occurs mainly as disseminated anhedral grains ranging between 5 μm and 1 mm across. It may be massive and locally very abundant, mainly in the semi-massive and the massive chalcopyrite–pyrite ore types. It is intergrown mainly with pyrite and sphalerite and less frequently with pyrrhotite and galena. Chalcopyrite occurs in three generations. The first early generation (ccp I) appears as interstitial fillings among pyrite framboids and as inclusions within pyrite (Figure 7d,e). Figure 7d demonstrates a good example of early-formed chalcopyrite that was crystallized after framboidal pyrite (py I) and before coarse subhedral pyrite (py II). The abundance of chalcopyrite blebs (<1 μm) in sphalerite indicates “chalcopyrite disease”, and they belong to the first chalcopyrite generation (ccp I). Often, these blebs are orientated along dislocations in sphalerite.

The second generation of chalcopyrite (ccp II), which is more widespread in mineralization, appears as disseminated grains, void fillings in quartz, veinlets in pyrite, and massive layers parallel to the foliation of the host rocks. The chalcopyrite veinlets crosscutting pyrite grains (Figure 7f) indicate that part of chalcopyrite was formed later than the main pyrite deposition (py II). The third generation of chalcopyrite (ccp III) occurs in the milky quartz veins and is intergrown with pyrite, sphalerite, galena, and tennantite-(Zn). Chalcopyrite is quite homogeneous in composition (Table 4, Supplementary Table S2). In ccp II, Co and Ni concentrations are low (up to 0.05 wt%), whereas As and Au reach 0.16 and 0.13 wt%, respectively. Variable Se values were detected in chalcopyrite at up to 0.16 wt%. In ccp III, Ni reaches up to 0.03 wt%, As reaches up to 0.11 wt%, Au reaches up to 0.02 wt%, and Se reaches up to 0.41 wt%. Chalcopyrite has stoichiometric average chemical formulae: Cu_1.00Fe_1.00S_2.00 for ccp II and Cu_1.00Fe_0.99S_2.01 for ccp III.

Pyrrhotite occurs exclusively as discrete minute inclusions (5 to 200 μm) in pyrite, and in some cases, it is intergrown with chalcopyrite. Microprobe analyses (Table 4, Supplementary Table S2) revealed that the composition of pyrrhotite varies within a narrow range from 60.46 to 61.05 wt% Fe and is homogeneous within individual grains. The calculated chemical formula is Fe_0.90S_1.00.

Sphalerite is found in anhedral grains up to 1 mm and occurs in three generations (sp I, sp II, and sp III). Sphalerite I contains fine blebs of chalcopyrite (chalcopyrite disease). Sphalerite II is intergrown with pyrite and chalcopyrite (Figure 7h), while the third generation of sphalerite (sp III) occurs in milky quartz veins, and is intergrown with pyrite, chalcopyrite, galena, and tennantite-(Zn). Sphalerite (Table 4, Supplementary Table S2) of all generations has a similar chemical composition and is characterized by low contents in Cd (up to 0.25 wt%) and Mn (up to 0.09 wt%) and variable Fe concentrations (0.16–3.88 wt%; 0 to 6.45 mole% FeS). The average chemical composition of the analyzed sphalerite is (Zn, Fe)_1.00S_1.00.

Galena has been identified in traces in massive chalcopyrite–pyrite mineralization (gn I) and in milky quartz veins (gn II). It occurs as rounded or elongated inclusions within chalcopyrite and pyrite. Silver contents reach 0.19 wt%, and the highest concentrations (0.17 wt% on average) are confined to the massive chalcopyrite–pyrite ore type. With respect to Se contents, two chemically distinct types of galena exist (Table 4, Supplementary Table S2): one with a higher Se concentration (1.10 to 1.68 wt%) in massive chalcopyrite–pyrite mineralization (gn I), and the other with lower Se contents (up to 0.42 wt%) hosted in quartz veins (gn II). These two variations in galena correspond to the chemical formulae Pb_0.99(S,Se)_1.01 and Pb_1.00(S,Se)_1.00, respectively.

Tennantite-(Zn), the only identified sulfosalt phase at Xylagani, is rare and occurs in disseminated and semi-massive mineralization (tn I) and in quartz veins (tn II). It is associated with pyrite, sphalerite, galena, and chalcopyrite, and the size of the grains range between 10 and 300 μm. Microprobe analyses (Table 4, Supplementary Table S2) revealed that Ag is low, reaching 0.35 wt% in tn I and 0.77 wt% in tn II. The average chemical formula is (Cu,Ag,Cd)_9.91(Fe,Zn)_2.01(As,Sb)_4.09(S,Se)_13.00 for tn I, and (Cu, Ag,Cd)_9.72(Fe,Zn)_2.16(As,Sb)_4.11(S,Se)_13.01 for tn II. Both tennantite types are characterized as tennantite-(Zn) due to the elevated Zn content, which is 6.13 wt% (1.41 apfu) for tn I and 7.16 wt% (1.65 apfu) for tn II.

Native gold occurs in pyrite from the semi-massive ore type and forms elongated, angular, or rounded grains of up to 20 μm across (Figure 7e). Data obtained from microprobe analyses (

Table 6. Microprobe analyses in wt% of native gold from semi-massive mineralization. bdl: below the detection limit.

wt%	16 D4	16 D5	16-7 A1	Average
Bi	0.24	bdl	0.08	0.08
Au	97.41	97.18	93.67	97.18
Ag	1.94	3.05	5.94	3.05
Cu	0.25	0.26	0.13	0.25
Total	99.84	100.49	99.82	100.56
Atomic composition
Bi	0.22	0.00	0.07	0.15
Au	95.55	93.84	89.22	93.84
Ag	3.47	5.38	10.33	5.38
Cu	0.76	0.78	0.38	0.76

) show that Ag contents range between 1.94 and 5.94 wt% and for Cu from 0.13 to 0.26 wt%, while Bi rises in one case up to 0.24 wt%. The average chemical formula is (Au,Ag,Cu,Bi)_1.00.

Mineralization is associated with extensive hydrothermal alterations, e.g., silicification and chloritization, which often envelopes the mineralized lenses and replace volcanic rocks. Alteration intensity increases from the margins toward the center of the mineralized zones, progressing from weak to pervasive, with complete replacement of the host rocks by quartz, chlorite, sericite, and sulfides.

Hydrothermal quartz, the main host of the ore minerals, shows a variety of textures. Most commonly, it is polycrystalline with adulatory extinction and suturing of the contacts between crystals, implying extended recrystallization. In many cases, brittle deformation has produced a cataclastic texture. One of the most common textures is the development of quartz fibers around euhedral pyrite crystals, forming pressure fringes (Figure 7g). Hydrothermal white mica (sericite) and chlorite were deposited with quartz (Figure 7c), and all were interpreted as syngenetic with sulfide mineralization.

Thirty microanalyses of hydrothermal chlorite from the mineralization at Xylagani (Supplementary Table S3) are plotted on the classification diagram of [23] and fall in the compositional field of ripidolite and pycnochlore (Figure 8). The contents of the calculated Al^IV range between 1.926 and 2.714 atoms per formula unit (apfu), and the Fe/(Fe + Mg) ratio varies from 0.322 to 0.420 (Figure 8).

Chlorite composition is related to formation conditions, e.g., temperature, pressure, redox conditions, and fluid and bulk-rock composition [24]. One of the most significant factors is temperature, and various methods have been proposed to determine it based on the chemical composition of chlorite. The most common methods use the empirical geothermometer that relates the temperature of the formation to the tetrahedral aluminum (Al^IV) content of chlorite [25,26] or the semi-empirical geothermometer that incorporates the influence of the Fe/(Fe + Mg) ratio in addition to the Al^IV content [27]. The application of both geothermometers at the hydrothermal chlorites of the mineralization at Xylagani yields similar temperatures to chlorite formation. The empirical thermometer of Cathelineau and Nieva [25] yielded temperatures between 222 and 301 °C, and the semi-empirical thermometer of Zang and Fyfe [27] yielded temperatures from 222 to 306 °C.

4.4. Sulfur Isotopes

The δ³⁴S values of pyrite range from +2.02 to +2.24‰ (average +2.09 ‰) and are homogeneous in both mineralization types (

Table 7. Sulfur isotopes of pyrite from semi-massive and massive mineralization.

Sample	Mineralization type	δ34S (‰)
NP 16-12	Semi-massive	2.12
NP 43	Semi-massive	2.08
NP 48-2a	Massive pyrite	2.02
NP 48-2a	Massive pyrite	2.09
NP 48-2a	Massive pyrite	2.24

).

5. Discussion

5.1. Geotectonic Environment of Formation and Metamorphism of Host Rocks

The mineralization of Xylagani occurs within metamorphosed mafic massive, pillow, or brecciated lava and in meta-pyroclastic rocks; based on the host rock composition and metal content, it is classified as a mafic-type (Cyprus-type) volcanogenic massive sulfide deposit (VMS). Major element concentrations were affected by metasomatic processes, e.g., hydrothermal alteration and/or low-grade metamorphism. This is reflected by the broad distribution of major oxides and the relatively high values of LOI (>2.5 wt%). The determination of the protolith type of the host metavolcanic rocks could not be obtained from the diagram of Nb/Y vs. Zr/TiO₂ by Pearce [28] because the Nb concentration is below the detection limit (<2 ppm).

The diagram of SiO₂ vs. Na₂O + K₂O (Figure 9a) by Le Bas et al. [29] was used with caution because the addition of SiO₂ during the hydrothermal alteration influenced the chemical composition of metavolcanic rocks. In this diagram, out of six samples, three are plotted in the field of basalt, one in the field of basaltic andesite, and two in the field of andesite. We assume that the samples with lower SiO₂ are the most representative, and so the protoliths must have a basaltic composition.

Therefore, the definition of the geotectonic environment of the deposition of the volcanic rocks was based on major and trace elements, e.g., Al, V, Y, and the high field strength elements Zr and Ti. These elements remain immobile during the geological evolution of the magmatic rocks, even if they are hydrothermally altered or metamorphosed, allowing us to identify their primary petrogenetic features [30].

The metavolcanic rocks of Xylagani are characterized by high contents of Cr, Ni, and low TiO₂, Zr, Y, and ΣREE. The ratio of Al₂O₃/TiO₂ is high and ranges from 27.45 to 48.21. These geochemical features suggest boninitic affinities. According to Piercey [31] and Pearce and Reagan [32], boninitic rocks that are associated with VMS deposits are characterized by high Cr and Ni concentrations, high Al₂O₃/TiO₂, low TiO₂ (<0.60 wt%), Zr (<58 ppm), and Y (<19 ppm). When the samples are plotted in the Zr/Y vs. Al₂O₃/TiO₂ diagram by Piercey [31], it is indicated that the protoliths of metavolcanics have a boninitic and low-Ti tholeiitic (LOTI) characteristic (Figure 9b).

Plots of the volcanic rocks in the discrimination diagrams of Ti vs. V (after [33]) and Ti vs. Zr (after [34,35]) show that these rocks lie in the fields of boninites and arc tholeiites (Figure 9c,d). Based on Piercey [31], this characteristic of metavolcanic rocks indicates the association of VMS mineralization with a juvenile environment at a forearc-to-volcanic arc spreading setting.

Figure 9. Discrimination of geochemical diagrams for the metavolcanic rocks of Xylagani. (a) SiO₂ vs. Na₂O + K₂O diagram (after [29]); (b) Zr/Y vs. Al₂O₃/TiO₂ diagram (after [31]) for rocks associated with VMS deposits from mafic-dominated environments (LOTI = Low-Ti tholeiites, MORB = mid-ocean ridge basalts); (c) V vs. Ti diagram (after [33]) for boninites, juvenile island arc tholeiites, mid-ocean ridge basalts (MORBs), ocean island basalts (OIBs) and alkaline island basalts OIBs; and (d) Ti vs. Zr diagram (after [34,35]) for boninites, juvenile island arc tholeiites, and mid-ocean ridge basalts (MORBs).

Figure 9. Discrimination of geochemical diagrams for the metavolcanic rocks of Xylagani. (a) SiO₂ vs. Na₂O + K₂O diagram (after [29]); (b) Zr/Y vs. Al₂O₃/TiO₂ diagram (after [31]) for rocks associated with VMS deposits from mafic-dominated environments (LOTI = Low-Ti tholeiites, MORB = mid-ocean ridge basalts); (c) V vs. Ti diagram (after [33]) for boninites, juvenile island arc tholeiites, mid-ocean ridge basalts (MORBs), ocean island basalts (OIBs) and alkaline island basalts OIBs; and (d) Ti vs. Zr diagram (after [34,35]) for boninites, juvenile island arc tholeiites, and mid-ocean ridge basalts (MORBs).

Chondrite-normalized patterns of metavolcanic rocks in Xylagani are subparallel, with a slight depletion in light REE (LREE) and flattening in the heavy REE (HREE) region. These characteristics are consistent with the interpretation that the studied volcanic rocks belong to the tholeiite association and attest to a supra-subduction zone setting [36,37,38]. A slightly U-shaped REE profile is observed in sample NP 36b, indicating a boninitic composition of the protolith. A clear but weak negative Eu anomaly appeared in the analyzed samples, except for one case (sample NP 11) where Eu demonstrated a positive anomaly, which may be attributed to local plagioclase removal or accumulation, respectively, in the magma chamber. A small negative anomaly was also observed in Ce, apart from sample NP 19, which demonstrates a strong negative Ce anomaly. The negative Ce anomalies indicate the involvement of the subduction of sediments in arc systems in a strong oxidizing environment [17,39]. Recently, it has been suggested that negative Ce anomalies may also be caused by the incorporation of subcontinental lithospheric material at shallow depths during plume ascent [40] or by magmatic processes such as partial melting [41].

The boninitic-to-low-Ti tholeiitic composition of the metavolcanic rocks in Xylagani indicates that mineralization is associated with a juvenile environment where mantle-derived volcanic rocks were formed. In many ophiolite-associated mafic boninitic-to-low-Ti tholeiitic rocks that incorporate VMS deposits, like in Cyprus and Oman, the geotectonic setting is associated with forearc or arc extension, and with the initiation of a back-arc basin, formed in a supra-subduction and roll-back settings [42]. The intra-oceanic subduction of the Vardar oceanic slab was active from around 180 Ma until 163.5 Ma (in the Early–Middle Jurassic period) and produced forearc or volcanic arc boninitic and tholeiitic magmas [18].

In the case of Xylagani VMS mineralization, the metavolcanic rocks indicate a juvenile environment at a forearc-to-volcanic arc spreading setting during intra-oceanic subduction initiation at the Vardar–Axios ocean in the Early–Middle Jurassic period. A similar geotectonic environment has been previously documented for the Evros ophiolite [11,16,17,18].

The Pb isotopic compositions of pyrite from Xylagani, analyzed by Frei [43], range from 18.37 to 18.384 for ²⁰⁶Pb/²⁰⁴Pb, from 14.646 to 15.672 for ²⁰⁷Pb/²⁰⁴Pb, and from 38.507 to 38.586 for ²⁰⁸Pb/²⁰⁴Pb. These values align with those of the Evros ophiolite reported by Bonev et al. [44], which exhibit narrow ranges and indicate a contribution from a depleted MORB-type lithospheric mantle source (Figure 10). Additionally, Pb isotope compositions from the Evros ophiolite, along with Nd-Sr isotope systematics, suggest a mantle–crust interaction driven by mantle wedge magmatic processes in the CRB arc/back-arc system [45].

The studied volcanic rocks at Xylagani were subjected to two metamorphic events: (i) ocean-floor metamorphism, probably via water–rock interactions during the circulation of seawater in the Early–Middle Jurassic period; (ii) orogenic metamorphism during the northward thrusting of the Makri Unit together with the Evros ophiolite over the Rhodope, during the Late Jurassic–Early Cretaceous period [14,15,18].

The incorporation of seawater heated by thermal transfer from the underlying magma during the ocean-floor metamorphism of the volcanic rocks in the Xylagani area in an oxidizing environment is documented by REE patterns, especially by the negative Ce and Eu anomalies. It was proposed [46,47] that the metamorphism grade due to seawater convection systems may reach up to greenschist facies temperatures (350–400 °C).

The ocean-floor metamorphism of the host metavolcanic rocks at Xylagani is evidenced by the presence of quartz, chlorite, sericite, prehnite, pumpellyite, epidote, zoisite, titanite, actinolite, and calcite. The metamorphic mineral parageneses contain quartz–albite–pumpellyite–chlorite, quartz–prehnite–chlorite, prehnite–actinolite and pumpellyite–actinolite facies rocks. The stability field of prehnite–pumpellyite phases, according to Starr et al. [48], is between ~200 °C and 280 °C. Moreover, the parageneses of quartz–albite–pumpellyite–chlorite and/or quartz–prehnite–chlorite observed in the study area are stable at pressures of 1 to 4 kbar. The presence of actinolite in the metavolcanic rocks indicates that the metamorphic temperature for pressures between 1 and 4 kbar does not exceed 360 °C in the prehnite–actinolite and pumpellyite–actinolite phases [49,50]. Similar conclusions were drawn by Magganas [21], who determined metamorphic temperatures ranging from 220 °C to 400 °C, and pressures from 1 to 4 kbar. Consequently, the metavolcanic rocks at Xylagani were metamorphosed at very low-grade prehnite–pumpellyite facies to low-grade greenschist facies conditions at temperatures up to 360 °C and pressures between 1 and 4 kbar during seawater circulation.

The subsequent orogenic metamorphism was produced during the northward tectonic emplacement of Makri Unit together with the Evros ophiolite onto the Rhodope continental margin in the Late Jurassic–Early Cretaceous period. This collision event is associated with Alpine Orogeny, which affected the Carpathians and the Balkan units [15]. The conditions of orogenic metamorphism were similar to those of ocean-floor metamorphism, reaching up to greenschist facies [19,21], making it quite questioning to distinguish the characteristic minerals of each event. In this collisional setting, the metavolcanic rocks were folded, and nearly barren milky quartz veins were developed, crosscutting the foliation of the metavolcanic host rocks.

5.2. Formation Conditions and Origin of the Mineralization

The sulfide mineralization at Xylagani shows typical features of a stratabound–stratiform deposit and occurs as silicified lenses and layers at different stratigraphic levels, with disseminated-to-massive sulfides and gold. On the ternary Pb-Cu-Zn diagram, which includes the VMS classification according to their base-metal content and the nature of the host rocks [51,52], the majority of the analyzed samples are plotted around Cu (Figure 11), apart from a few exceptions, which show higher local Pb and Zn contents. This reflects the type of VMS mineralization, which is mafic (Cu), i.e., Cyprus-type or Cu-pyrite with Au.

Due to the absence of typical structures that indicate a seafloor deposition, e.g., sulfide chimneys or fossils within the host rocks, it is suggested that the mineralization of Xylagani can be classified as a subseafloor deposit. Several textural features of the mineralization support the syn-volcanic-to-early diagenetic subseafloor origin of the ore bodies. These textures are still well-preserved despite the subsequent deformation and the very low-to-low-grade ocean floor and orogenic metamorphism. They include mineralized lenses with semi-massive mineralization, thin layers, lenses of massive pyrite, and layers of massive chalcopyrite–pyrite that exhibit stratabound or stratiform structures.

Within the mineralized lenses, alteration intensity increases from weak at the margins to pervasive at the core, where the original rocks are completely replaced by quartz, chlorite, and sericite, with sulfides. The mineralization and associated hydrothermal alteration, which have been overprinted by deformation related to subsequent metamorphism, along with the strong stratigraphic control on mineralization, indicate that ore formation is synchronous with volcanic activity. The Xylagani deposit exhibits alteration that parallels stratigraphy and extends into the metavolcanic rocks, which is a feature commonly observed in subseafloor replacement-style VMS [53].

Framboidal and fine-grained pyrite is the early generation pyrite (py I), and forms disseminated grains or thin layers, implying subseafloor replacement processes, which are similar to those that have been previously described by [53]. Disequilibrium conditions caused by the rapid crystallization of fine-grained pyrite were the main mechanism for the formation of the framboidal texture as a result of mixing between hydrothermal fluid and circulating seawater in the host rocks [53,54]. The presence of framboidal pyrite in magmatic–hydrothermal environments indicates low temperatures, slightly higher than 200 °C [55]. This upper-temperature limit is comparable to the formation temperature of hydrothermal chlorite at Xylagani, which ranges between 222 and 306 °C. The close intergrowth of chlorite with framboidal pyrite (Figure 7c) suggests similar formation conditions in the subseafloor.

Moreover, framboidal pyrite occurs as cores or isolated clusters within larger euhedral pyrite crystals (Figure 7d), suggesting that framboids act as nuclei for subsequent crystal growth. According to Piercey et al. [53], this is evidence of a replacement texture in the subseafloor. Sphalerite I contains fine blebs of chalcopyrite (chalcopyrite disease). These blebs form after sphalerite deposition [56,57], with chalcopyrite replacing sphalerite with Cu-rich fluid impulses.

The δ³⁴S values of pyrite (+2.02 to +2.24‰) in semi-massive and massive mineralization types (Table 7) are homogeneous, indicating that the physicochemical conditions did not change during ore deposition. These values can result from mixing between Early Triassic seawater (δ³⁴S = 16–17‰) [58] and igneous rocks originating from the mantle, e.g., basalts (δ³⁴S = −2 to −1‰) [59]. The more positive values recorded in pyrite from Xylagani cannot be attributed to sulfur fractionation during partial melting, but they are most likely a result of hydrothermal seawater alteration through the convection of hydrothermal fluids and fluid–rock interactions [60,61]. It should be noted that pre-metamorphic sulfur isotope compositions are preserved regardless of the metamorphic grade or whether the sulfides have undergone recrystallization [62]. Similarly, in Xylagani, the S-isotopic composition of pyrite was not affected by the subsequent ocean floor and orogenic metamorphism.

In this setting, the circulation of hydrothermal fluids through the host volcanic rocks led to pervasive hydrothermal alteration, e.g., intense silicification and chloritization and the development of a subseafloor replacement VMS deposit. Continued hydrothermal activity caused the leaching of metals, such as Au, Ag, Cu, Fe, Zn, Pb, As, Co, and Ni, from the volcanic rocks of the Evros ophiolite. These metals were subsequently precipitated as sulfide and gold mineralization, along with quartz, chlorite, and minor sericite, forming replacement-style ore bodies within the pore spaces and zones in the volcanic material produced by pervasive hydrothermal alteration. This process reflects the increasing intensity of fluid–rock interactions during deposit evolution. A similar mechanism for VMS deposits has been proposed in ref. [53,63]. However, discordant stockwork veins, which typically underlie massive sulfide lenses, were not identified in the study area and may be preserved below the current level of exposure.

5.3. Post-Ore Metamorphism of Mineralization

The very low-to-low-grade ocean-floor metamorphism and the subsequent orogenic Alpine metamorphism with related deformation have affected mineralization in the Xylagani area. Characteristic textures resulting from this post-depositional process are exhibited in the sulfides, especially in pyrite and quartz of the mineralized bodies. The thin mineralized layers, which display macroscopic ductility with the development of folds, show that they have been deformed simultaneously with volcanic rocks. In deposits that have been subjected to orogenic Alpine metamorphism and deformation, the development of folds in pyrite-rich units is a common process [64].

Hydrothermal quartz shows extended recrystallization. This is the reason that hydrothermal quartz at Xylagani is free from fluid inclusions due to dynamothermal processes, as has been previously suggested [62]. In many cases, brittle deformation has produced a cataclastic texture.

The effects of metamorphism on mineralization are also shown by the compaction and recrystallization of framboidal pyrite. This process has obliterated the internal structure of the framboids in many cases and resulted in their homogenization. The total destruction of the framboidal textures was avoided due to the presence of fine-grained quartz, which preserved them from compaction [65,66]. The pressure fringes of quartz developed around euhedral pyrite crystals are also an indication of the effects during low-grade metamorphism [67]. In most cases, this texture forms due to pyrite’s resistance to deformation because of its high hardness. It is suggested that the quartz coexisting with sulfides at Xylagani was deposited as fine-grained quartz and is synchronous with the formation of mineralization. Post-depositional recrystallization due to orogenic metamorphism resulted in the local coarsening of fine-grained quartz to sutured grain boundaries and the development of pressure fringes.

In addition to textural changes, the chemical composition of pyrite was slightly affected due to metamorphism in the investigated mineralization. Previous studies [68,69,70,71] have shown that the Co/Ni ratios in pyrite are useful in distinguishing ore-forming environments. Pyrites from Xylagani mineralization have Co/Ni ratios > 2 (3.11 to 3.94; Table 5), indicating a volcanogenic–hydrothermal genesis. In the Ni vs. Co diagram (after Bajwah et al. [70] and Brill [71]), pyrites from massive mineralization are plotted close to but not within the volcanogenic field (Figure 12), indicating that they were chemically affected by recrystallization during orogenic metamorphism.

Sulfides are rarely found within milky quartz veins that crosscut both the metavolcanic rocks and their foliation, indicating that they were formed during the metamorphic stage and most probably during collisional orogenic metamorphism. During this event, mineralization experienced limited remobilization into milky quartz veins.

6. Genetic Implications and Conclusions

The deposit at the Mylorema Creek of Xylagani belongs to the type of deposits that are characterized as a mafic Cyprus-type VMS deposit. It is genetically associated with mafic magmatic rocks with a basaltic composition and boninitic-to-low-Ti tholeiitic affinities. These rocks formed in a submarine volcanic environment during the Early–Middle Jurassic period, within a primitive forearc or arc with intra-oceanic supra-subduction spreading and roll-back setting at the Vardar–Axios ocean, are part of the ophiolitic sequence of Evros (Figure 13a).

Heated seawater played a significant role in the pervasive hydrothermal alteration of volcanic rocks and the genesis of mineralization. The hydrothermal circulation of the subseafloor through fractures and faults of the volcanic rocks leached metals, such as Au, Ag, Cu, Fe, Zn, Pb, As, Co, and Ni, from the underlying volcanic rocks of the Evros ophiolite, through fluid–rock interactions, and transported them to the subseafloor environment (Figure 13b). These metals were subsequently precipitated as sulfide and gold mineralization, along with quartz, chlorite, and minor sericite, forming replacement-style ore bodies within the pore spaces and zones in the volcanic material produced by pervasive hydrothermal alteration. The hydrothermal alteration occurred at temperatures reaching up to 300 °C, while mineralization began slightly above 200 °C, marked by the formation of early framboidal pyrite in a subseafloor setting.

The general model for the mineralization genesis of mafic Cyprus-type VMS suggests the presence of stockwork feeder zones enriched in Cu and other metals. However, such stockworks leading to the formation of the stratabound mineralization at Xylagani have not been identified in the area, making it impossible to determine the exact location of the fluid discharge.

The very low-to-low-grade ocean-floor metamorphism in the Early–Middle Jurassic period and the subsequent orogenic metamorphism and related deformation during the Alpine collision in the Late Jurassic–Early Mesozoic period affected the mineralization and formed milky quartz veins with rare sulfides, crosscutting the metavolcanic rocks. The ore minerals were possibly formed due to the remobilization of primary mineralization during metamorphism.