Genesis and mineralization style of gold occurrences of the Lower Lom Belt, Bétaré Oya district, eastern Cameroon (original) (raw)

Genesis and mineralization style of gold occurrences of the Lower Lom Belt, Bétaré Oya district, eastern Cameroon

Kevin Igor Azeuda Ndonfack a,b{ }^{\mathrm{a}, \mathrm{b}}, Yuling Xie a,∗{ }^{\mathrm{a}, *}, Richard Goldfarb c{ }^{\mathrm{c}}, Richen Zhong a{ }^{\mathrm{a}}, Yunwei Qu a{ }^{\mathrm{a}}
a{ }^{a} School of Civil and Environmental Engineering, University of Science and Technology Beijing, Beijing 100083, PR China
b{ }^{\mathrm{b}} Department of Earth Sciences, University of Yaoundé I, Yaoundé, P. O. Box 812, Cameroon
c{ }^{c} State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Beijing 100083, PR China

A R T I C L E I N F O

Keywords:
Cameroon
Bétaré Oya district
Lower Lom Belt
Fluid inclusions
Stable isotopes
Orogenic gold

A B STR ACT

The Bétaré Oya district has a substrate composed of the Neoproterozoic metavolcanic-metasedimentary rocks of the Lom Belt in eastern Cameroon. The district is well-known for alluvial gold mining activities, however, the primary gold mineralization has received little attention. In the current study, we newly report geological, ore mineralogy, fluid inclusion microthermometric, laser Raman spectroscopy, and stable isotope data from auriferous veins of the Lom Belt to characterize the style of gold mineralization and to constrain the origin of the oreforming fluids. The auriferous quartz veins are laminated, fractured, N - to NE-trending, and spatially associated with the Bétaré Oya Shear Zone. Fieldwork coupled with microscopic examination and the textural relationships of ore minerals revealed two stages of mineralization. The first stage is characterized by the presence of pyrite, sphalerite, galena, chalcopyrite, pyrrhotite, hematite, petzite, hessite, wolframite, electrum, and gold, while the second stage is characterized by the presence of a later deposition of galena and pyrite, as well as minor greenockite. The gangue minerals are quartz, sericite, muscovite, chlorite, calcite, ankerite, and barite, whereas the supergene assemblage (stage 3) includes goethite, hematite, covellite, and enargite. Two fluid inclusion assemblages containing three types of inclusions are identified in mineralized quartz veins. Assemblage 1 consists of gold-related types 1 and 2 with daughter minerals consisting of nahcolite, magnesite, and arsenolamprite (black native arsenic), whereas assemblage 2 consists of post-gold type 3 fluid inclusions. Type 1 has H2O−NaCl−\mathrm{H}_{2} \mathrm{O}-\mathrm{NaCl}- CO2±N2±CH4\mathrm{CO}_{2} \pm \mathrm{N}_{2} \pm \mathrm{CH}_{4} primary fluid inclusions (FI), with a TmCO2\mathrm{Tm}_{\mathrm{CO} 2} ranging from -59.8 to −56.6∘C-56.6{ }^{\circ} \mathrm{C}, salinities from 0.5 to 10.8wt%NaCl10.8 \mathrm{wt} \% \mathrm{NaCl} eq., densities from 0.87 to 1.00 g.cm−31.00 \mathrm{~g} . \mathrm{cm}^{-3}, and total homogenization temperatures between 280 and 360∘C360^{\circ} \mathrm{C}. Type 2 contains CO2(±H2O−NaCl)±N2±CH4FI\mathrm{CO}_{2}\left( \pm \mathrm{H}_{2} \mathrm{O}-\mathrm{NaCl}\right) \pm \mathrm{N}_{2} \pm \mathrm{CH}_{4} \mathrm{FI} that exhibit TmCO2\mathrm{Tm}_{\mathrm{CO} 2} ranging between -60.0 and −56.7∘C,ThCO2-56.7^{\circ} \mathrm{C}, \mathrm{Th}_{\mathrm{CO} 2} from 13 to 25∘C25^{\circ} \mathrm{C}, and densities between 0.73 and 0.85 g.cm−30.85 \mathrm{~g} . \mathrm{cm}^{-3}. Type 3 shows H2O−NaClFI\mathrm{H}_{2} \mathrm{O}-\mathrm{NaCl} \mathrm{FI} with salinities between 0.2 and 10.1wt%NaCl10.1 \mathrm{wt} \% \mathrm{NaCl} eq., densities between 0.82 and 0.98 g.cm−30.98 \mathrm{~g} . \mathrm{cm}^{-3}, and total homogenization temperatures from 160 to 235∘C235^{\circ} \mathrm{C}. Measured δ18O\delta^{18} \mathrm{O} for gold-bearing quartz ( +11.5 to +16.0%+16.0 \% ), δD\delta \mathrm{D} from FI(−50.6\mathrm{FI}(-50.6 to −21.8%-21.8 \% ), δ13C\delta^{13} \mathrm{C} from FI(−5.8\mathrm{FI}(-5.8 to −5.5%-5.5 \% ), and δ34 S\delta^{34} \mathrm{~S} from galena and pyrite grains (+5.3%(+5.3 \% and +8.2%+8.2 \%, respectively) suggest a metamorphic source as most likely for the ore-forming fluids and sulfur, although a mantle CO2\mathrm{CO}_{2} contribution cannot be ruled out.

The gold deposition probably took place by fluid-rock interaction and fluid unmixing at −310∘C-310{ }^{\circ} \mathrm{C} and at a depth of about 6−9 km6-9 \mathrm{~km}. The ore-forming fluid was a low salinity (−6.2wt%\left(-6.2 \mathrm{wt} \%\right. NaCl eq. )H2O−NaCl−CO2±N2±CH4) \mathrm{H}_{2} \mathrm{O}-\mathrm{NaCl}-\mathrm{CO}_{2} \pm \mathrm{N}_{2} \pm \mathrm{CH}_{4} solution. These data revealed that the Lom Belt gold mineralization is best classified as a mesozonal orogenic gold deposit.

1. Introduction

Orogenic gold deposits are hosted in rocks typically metamorphosed at lower to upper greenschist facies (McCuaig and Kerrich, 1998). These deposits were formed during compressional to transpressional tectonic
processes at convergent plate margins in accretionary and collisional orogens (Groves et al., 1998; Goldfarb et al., 2005). Generally, orerelated fluids are aqueous-carbonic with a variable content of CH4, N2\mathrm{CH}_{4}, \mathrm{~N}_{2}, and H2 S\mathrm{H}_{2} \mathrm{~S}, and with low to moderate salinities (Groves et al., 1998). However, the source of fluids and metals, particularly whether involving

[1]

1. a{ }^{a} Corresponding author.
   E-mail address: yulingxie63@hotmail.com (Y. Xie). ↩︎

mantle, magmatic, and/or metamorphic reservoirs, is still debated (Goldfarb and Groves, 2015).

Lode gold occurrences hosted by Neoproterozoic metavolcanicmetasedimentary rocks are located in the Bétaré Oya district of the Lom Belt, eastern Cameroon. Initial exploration for primary gold mineralization in the area started in the 1960s after the release of the 1:500,000 geological map (Gazel and Gerard, 1954). In 1983, the Department of Mines and Geology of Cameroon targeted the Lom Belt as a prospective area for primary gold mineralization owing to the high production from alluvial gold deposits. However, until now, very little has been known about the primary gold occurrences in this particular area of Cameroon. Mlési et al. (2006) reported the gold resource of primary occurrences of the Bétaré Oya gold district at 20 t , although no systematic exploration had yet been done in the district. Studies undertaken on the geochemistry of gold-bearing veins revealed an association of Au with enrichments in Ag,Bi,As,Pb,Mo,W\mathrm{Ag}, \mathrm{Bi}, \mathrm{As}, \mathrm{Pb}, \mathrm{Mo}, \mathrm{W}, and Cu (Freyssinet et al., 1989; Vishiti et al., 2017). Gold mineralization in the Bétaré Oya district occurs in quartz veins and as disseminations in altered wallrocks. The veins cut the low- to medium-grade meta-volcanic-metasedimentary rocks that comprise the Lom Belt. The nature and genesis of the ore-forming fluids responsible for gold mineralization in the district are poorly understood.

This paper aims to define the features of ore-forming fluid responsible for gold mineralization hosted by the Lower Lom Belt in the Bétaré Oya district in order to develop an ore genesis model. The work is based on our geology, detailed ore mineralogy, fluid inclusion microthermometry, laser Raman spectroscopy, and C-O-H-S stable isotope studies. The mineralized quartz veins are controlled by the NE-trending Bétaré Oya Shear Zone and contain sulfides, tellurides, and gold. The fluid inclusion study indicates a low salinity aqueous-carbonic oreforming fluid of metamorphic origin that can be related to late Neoproterozoic D3 deformation in the belt.

2. Geological framework of the Pan-African orogeny in Cameroon

The Pan-African orogeny in central Africa comprises the Central African Belt that cuts across a major part of the north equatorial countries of Nigeria, Chad, Central African Republic, and Cameroon, and then extends eastwards to Sudan and Uganda (Toteu et al., 2006). It is interpreted as the deformation zone that was formed from the amalgamation of the Congo-São Francisco, West African, and Metasahara Cratons during the Neoproterozoic era (Kusky et al., 2003; Fig. 1a). In Cameroon, the Central African Belt consists of low- to high-grade Mesoproterozoic to Neoproterozoic metavolcanic-metasedimentary rocks and pre-, syn- and post-tectonic plutonic rocks (Ngako et al., 2008; Njonfang et al., 2008; Penaye et al., 2006; Toteu et al., 2006). These plutonic rocks are the result of the reworking of Archean to Paleoproterozoic continental crust in the Neoproterozoic and many intrusions contain xenoliths of Archean and Paleoproterozoic rocks (Ganwa et al., 2016; Tchakounté et al., 2017; Kamguia Kamani et al., 2021).

The Pan-African plutonic rocks in Cameroon consist of a variety of mafic to felsic bodies (Tchameni et al., 2006; Tchouankoue et al., 2016; Kwekam et al., 2020). The pre- to syn-tectonic plutonic rocks formed an older batholith in the west and northwest, while the syn- to late-tectonic plutonic rocks formed the younger batholith in the east and southeast of Cameroon (Fig. 1b). The emplacement of the older batholith occurred prior to Pan-African transcurrent shearing (Bouyo et al., 2016). The associated pre- to syn-tectonic plutonic rocks are calc-alkaline mafic to intermediate intrusions dated at ca. 740-630 Ma and are interpreted to have formed in a subduction-related continental arc (Toteu et al., 1987; Pinna et al., 1994; Bouyo et al., 2016). The syn- to late-tectonic plutonic rocks were formed during the main collisional period between cratonic blocks and were associated with the development of continental-scale transcurrent shear zones and widespread migmatization. These plutonic rocks are composed of mafic to felsic intrusions and were
emplaced during two episodes at ca. 620-600 Ma and ca. 590-580 Ma (Toteu et al., 2004; Asaah et al., 2015; Tchouankoué et al., 2016; Li et al., 2017).

The Central African Belt in Cameroon is dominated by three intracontinental basins. From the north to the south, these basins include the Poli, the Lom, and the Yaoundé basins, respectively (Toteu et al., 2006). The Poli basin is interpreted as a back-arc basin that opened between 830 Ma and 665 Ma (Toteu et al., 2006; Bouyo et al., 2015). The basin consists of metavolcanic-metasedimentary rocks. The Lom Belt, host to widespread alluvial and lode gold deposits, is a late-tectonic pull-apart basin composed of low- to medium-grade metamorphic rocks (Ngako et al., 2003). Toteu et al. (2006) argued that the sediments in the basin were deposited, deformed, and metamorphosed within the period of 613-600 Ma. However, new geochronological data indicate that the metamorphic event started before ca. 650 Ma (Azeuda Ndonfack, unpublished data). The Yaoundé basin is a pre- to syn-tectonic basin and consists of low- to high-grade metamorphic rocks that were intruded by a series of pre- to syn-tectonic plutonic rocks of alkaline to transitional composition (Fuh et al., 2021; Nkoumbou et al., 2006, 2014). The U-Pb dating of detrital zircons revealed a depositional age of sediments as ca. 625 Ma (Toteu et al., 2006). These sediments were metamorphosed up to granulite facies at ca. 620 Ma and subsequently thrust southward onto the Congo Craton (Penaye et al., 1993).

2.1. Gold mineralization in East Metallogenic Province of Cameroon

The East Metallogenic Province of Cameroon has three main gold districts including the Bétaré Oya, Woumbou-Colomine-Ketté, and Batouri districts. Primary gold mineralization in these districts is controlled by local splays of two N070 ∘{ }^{\circ}-trending, crustal-scale, sinistral strike-slip shear zones: the Central Cameroon Shear Zone and the Sanaga Shear Zone (Fig. 1b). In these splays, mineralized veins strike NNE-SSW and NE-SW along the Bétaré Oya Shear Zone, NE-SW and E-W along the Colomine Shear Zone, and NNE-SSW and NE-SW along the Batouri Shear Zone (Fig. 1b). Some veins cut the Neoproterozoic metavolcanicmetasedimentary rocks, undated older metamorphic rocks, and the Pan-African plutonic rocks, but other auriferous veins occur along the contact between rock layers (Wippern and Seyferls, 1966; Freyssinet et al., 1989; Suh et al., 2006; Fon et al., 2012; Vishiti et al., 2015, 2017; Fils et al., 2020; Azeuda Ndonfack et al., 2021; Fon et al., 2021). Milési et al. (2006) estimated gold resources as ∼20t\sim 20 \mathrm{t} in the Bétaré Oya district, ∼12t\sim 12 \mathrm{t} Au in the Woumbou-Colomine-Ketté district, and ∼15tAu\sim 15 \mathrm{t} \mathrm{Au} in the Batouri district. Most of the gold mineralized zones consist of quartz veins and adjacent altered wallrocks. Veins are most commonly composed of hydrothermal minerals such as quartz, carbonates, chlorite, white mica, sulfides, and/or tellurides. Previous fluid inclusion studies carried out in the Woumbou-Colomine-Ketté and Batouri districts indicated a low salinity, H2O−CO2\mathrm{H}_{2} \mathrm{O}-\mathrm{CO}_{2}-bearing fluid with minor amounts of N2\mathrm{N}_{2} and CH4\mathrm{CH}_{4}, which deposited gold at ∼300∘C\sim 300{ }^{\circ} \mathrm{C} during fluid unmixing (Suh et al., 2006; Azeuda Ndonfack et al., 2021).

3. Geology of the gold occurrences of the Bétaré Oya district

3.1. Metavolcanic-metasedimentary host rocks of the Lower Lom Belt

The southern portion of the Lom Belt, which is ∼126 km\sim 126 \mathrm{~km}-long and ∼10 km\sim 10 \mathrm{~km}-wide, is underlain by NE-trending Neoproterozoic metavolcanic-metasedimentary sequences that formed various schist types including quartzite and black shales or slate (Freyssinet et al., 1989; Ngako et al., 2003; Fig. 2b). The belt was intruded by plutonic rocks and mafic dikes. Schists along the southeastern margin of the belt are highly sheared and mylonitized. The most dominant area of shearing, along the contact between undated high-grade migmatitic gneisses and schists, as well as defining the eastern edge of the Lom Belt (Fig. 2a), is termed the Bétaré Oya Shear Zone (Kankeu et al., 2009). The rocks of the belt were metamorphosed into greenschist facies

Fig. 1. a) Map of western Gondwana showing the amalgamation between the Sahara, West African, and Congo-Sáo Francisco cratons (modified from Kusky et al., 2003), TS: Trans Sahara Belt; DA: Damara Belt; and b) Geological map of Cameroon showing the Neoproterozoic intercontinental basins and the Pan-African plutonic rocks associated with major litho-tectonic structures (from Azeuda Ndonfack et al., 2021). TBSZ: Tcholliré-Banyo Shear Zone; CCSZ: Central Cameroon Shear Zone; SSZ: Sanaga Shear Zone.

Fig. 2. a) Geology map of the Lower Lom Belt (modified from Gazel and Gerard, 1954; Kankeu et al., 2009; Fils et al., 2020); and b) Geological profile of the mineralized zone of the Mborguéné gold occurrence from Freyssinet et al. (1989). The locations of the studied gold occurrences are also shown: MB (Mbal); HH Haya Haya; MG (Mborguéné). BOSZ: Bétaré Oya Shear Zone.
assemblages.
The rocks of the Lower Lom Belt can be grouped into three categories based on their structure and mineralogy (Fig. 2a). Gold occurs in all three litho-units, which from the margins to the interior of the belt include: (1) Chlorite-sericite schist that is spatially associated with the margins of the ductile to brittle shear zones and includes the Bétaré-Oya Shear Zone (Fig. 2a). The chlorite-sericite schist is strongly deformed, with interbedded quartz veins displaying pinch-and-swell and shearing features (Fig. 3a, f), which appear to represent Riedel structures related to the main shear zone. The NE-striking schistosity generally has a high to moderate northwesterly dip. The mineralogical composition of the schist consists of fine-grained quartz, ilmenite, rutile, zircon, pyrite, muscovite, sericite, and staurolite. Garnet, magnetite/hematite, and graphite are also locally abundant in this unit (Gazel and Gerard, 1954). (2) Sericite-muscovite schist occurs father inward within the belt
(Fig. 3b, 3g). It is also NE-striking and has a high NW dip, and is characterized by microstructures including crenulations and garnet fish. White mica (sericite-muscovite) is the main phyllosilicate mineral in the schist and occurs with fine-grained biotite, quartz, plagioclase, and garnet, and minor pyrite, ilmenite, chalcopyrite, sphalerite, zircon, monazite, and apatite. All minerals display a C-S microstructure. (3) The complexly interbedded quartzitic schist (Fig. 3c), quartzite (Fig. 3d, 3h), biotite-muscovite schist, sericite schist, slate, and actinolite-biotitechlorite schist occur in the center of the belt (Fig. 2b). The rocks are weakly to strongly deformed, NE-trending, and show steep to nearvertical dips and most of the gold occurrences are preferentially hosted in these relative low strain rocks (Fig. 2). Based on textures and mineral shapes, suites of minerals were distinguished as detrital (quartz, biotite, feldspar), metamorphic (quartz, albite, actinolite, epidote, clinozoisite, garnet, titanite), hydrothermal hypogene (quartz, sericite,

Fig. 3. Macro- and microphotographs of the Lower Lom Belt rocks. a) NE-trending sericite-chlorite schist expressing sinistral shear movement; b) Sericite-muscovite schist showing schistosity; c) Quartzitic schist showing NE-dipping schistosity; d) Hand sample of quartzite; e) Hand sample of mafic dike (lamprophyre); f) Microphotograph of the rock seen in the Fig. 3a showing C-S microstructures; g) Microphotograph of the rock seen in the Fig. 3b displaying C-S microstructure marked by sericite (Ser), muscovite (Ms), quartz (Qtz), and pyrite (Py); h) Microphotograph of the rock seen in the Fig. 3d showing C-S microstructure; and i) Microphotography of the rock seen in the Fig. 3e showing plagioclase (PI) and biotite (Bt).
muscovite, chlorite, calcite, sulfides), and supergene (barite, jarosite, goethite).

Plutonic rocks in the district are syn-tectonic (see below section 3.2) and show heterogeneous deformation features varying from undeformed to strongly deformed mineral grains within a single pluton (Ngako et al., 2003). The Bétaré Oya peraluminous, calc-alkaline granitic stock was emplaced into the chlorite-sericite schist and has a granodioritic to tonalitic affinity (Ateh et al., 2017). The pluton shows evidence of D2 sinistral deformation along its NW and NE margins when examined by aeromagnetic and radiometric images (Ateh et al., 2017), but the total amount of displacement along the Bétaré Oya Shear Zone is unknown. The intrusion was dated at ca. 635 Ma and formed in a subduction-related magmatic arc setting (Ateh et al., 2017). Mafic dikes occur in the belt (Fig. 3e, 3i) and underwent the same regional greenschist facies metamorphism. They display porphyritic textures with plagioclase phenocrysts in a groundmass of plagioclase and a relic olivine-pyroxene-amphibole assemblage strongly metamorphosed to biotite, chlorite, calcite, epidote, and sericite, with traces of zircon and apatite. Some phenocrysts of plagioclase are fractured and display kink band microstructures, which are commonly filled by calcite and sericite.

3.2. Structural features of the Lower Lom Belt

The metavolcanic-metasedimentary rocks of the Lower Lom Belt underwent three phases of deformation (D1-D3). The D1 event is defined by compressional tectonics that developed an S1 foliation. This was overprinted by a second compressional D2 deformation and a D3 transpressional event (Kankeu et al., 2009, Fils et al., 2020). The D2 and D3 deformation episodes, commonly transitional, are identified by C-S microstructures, drag folds, local dextral and sinistral shearing, and development of the Bétaré Oya Shear Zone. The C-S microstructures are most intense along the contact between schists and surrounding gneiss and are relatively weak in the inner part of the belt. The drag folds that we observed in the field consistently showed sinistral movement, although both dextral and sinistral structures had been previously reported (Kankeu et al., 2009). Gold-bearing quartz veins were formed during the D3 transpressional/transtensional event and are spatially associated with the P and P′\mathrm{P}^{\prime} of the Riedel model (Fils et al., 2020).

3.3. Gold mineralization and quartz vein microstructures

Gold-bearing quartz veins commonly occur along the contact

between lithologies, such as schists and quartzites, or within shears in the low strain schist units parallel to or cut the schistosity. Mineralized quartz veins are sheared, vuggy, brecciated, and lenticular. They typically define a stockwork system and surrounding host rocks are highly silicified (Gazel and Gerard, 1954). Ore minerals consist of pyrite, galena, chalcopyrite, and arsenopyrite along with quartz, sericite, and carbonate as gangue phases (Gazel and Gerard, 1954; Vishiti et al., 2017). Veins are NE-trending and −60∘NW-60^{\circ} \mathrm{NW} dipping, have lengths of more than 100 m , and are as wide as 10 m at the Haya Haya and Mborguéné occurrences (Fon et al., 2012; Freyssinet et al., 1989; Fig. 2). Freyssinet et al. (1989) and Fils et al. (2020) emphasized how gold occurrences could be hosted in various lithological units including the pelitic metasediments, metavolcanics, and/or quartzites (Fig. 2).

The mineralization style at the Mbal (MB), Haya Haya (HH), and Mborguéné (MG) occurrences is characterized by laminated veins displaying transverse fractures (Fig. 4a-c). The veins are whitish with tarnished pinkish areas comprising an abundance of coarse-grained pyrite, sphalerite, galena, and carbonate minerals (Fig. 4d-f). Microscopic features of the quartz veins consist of undulatory extinction, lamellae deformation, subgrain boundaries, and highly sutured grain boundaries (Fig. 4g-i). These microstructure features were recognized as the result of plastic deformation (Jacques et al., 2014). The subgrains developed
along boundaries of coarse quartz grains (Fig. 4g-i) were interpreted to have been formed by bulging recrystallization (Kankeu et al., 2009). Also, large elongated quartz grains with varying sizes of subgrains may have resulted from ductile deformation. The undulatory extinction with a strongly sweeping character is dominant throughout the mineralized veins. The sericite and muscovite observed along quartz grain boundaries show a relative sheared character (Fig. 4h). These white micas are also found as inclusions within quartz. Despite the identified plastic deformation features, the mineralized veins at the Mbal, Haya Haya, and Mborguéné occurrences lack C-S microstructures, thus implying that the veins were formed during the late D3 transpressional/transtensional deformation.

4. Methods

4.1. Sampling

Thirty samples were collected within the Lower Lom Belt for analytical work, including 15 mineralized vein samples from open pits and surface outcrops (seven from the Mbal occurrences, four from Haya Haya, four from Mborguéné), five samples of altered wallrocks, and ten samples of relatively fresh host rocks (chlorite-sericite schist, sericite-

Fig. 4. Outcrops, hand samples, and microphotographs of the mineralized veins of the Lower Lom Belt. a) NE-striking (NW dip) quartz veins outcropping at Mborguéné; b) NE-striking (SE dip) quartz veins showing oxidation zone from Haya Haya occurrence; c) N-trending mineralized veins from Mbal occurrence; d) Sample showing grains of pyrite and galena from Mborguéné; e) Sample from Haya Haya showing disseminated pyrite; f) Samples from Mbal and Mborguéné occurrence showing coarse grains of pyrite sphalerite, and galena as well as calcite; g) Anhedral crystal of quartz showing recrystallized subgrain boundaries (confer arrow) from Mborguéné; h) Sericite preferentially located in the suture boundaries of quartz from Haya Haya occurrence; and i) Elongated quartz crystals expressing laminated microstructure associated with recrystallized subgrain boundaries from Mbal occurrence.

muscovite schist, actinolite-biotite-chlorite schist, quartzitic schist, quartzite, mafic dike). Thin sections were made on all samples for mineralogical investigation. Subsequently, representative samples from mineralized veins were selected for fluid inclusion microthermometry and C-H-O-S isotope study of the quartz and related sulfides.

4.2. Ore mineralogy

The identification of ore minerals was performed using reflected and transmitted light microscopy with an Olympus BX51 microscope and scanning electron microscopy coupled with energy dispersive spectroscopy (SEM/EDS) using a Phenom-World instrument at the Key Laboratory of Mineral Resources at the University of Science and Technology Beijing. Polished or thin sections were coated with gold for SEM/EDS analysis. The detector for the Phenom-Word analysis used a beam voltage of 15 kV and a measuring time of 30 s for each analysis. Based on observed textural relationships between minerals, the paragenetic sequence was established.

4.3. Fluid inclusion microthermometry

Microthermometric measurements of fluid inclusions were performed on the Linkam THMS600 heating-freezing stage ( +600/−196∘C+600 /-196^{\circ} \mathrm{C} ) coupled to an Olympus microscope at the Key Laboratory of Mineral Resources, University of Science and Technology Beijing. The uncertainty associated with microthermometric measurements, based on the reproducibility of measurements conducted on reference standards and the measured precision of the equipment, ranged from 0.1 to 5∘C5{ }^{\circ} \mathrm{C}. The rate of heating-freezing was 25∘C/min25{ }^{\circ} \mathrm{C} / \mathrm{min} but was reduced to 0.1−0.5∘C/min0.1-0.5^{\circ} \mathrm{C} / \mathrm{min} when close to a phase transition temperature, such as the melting point of CO2\mathrm{CO}_{2}, ice, and clathrate or homogenization temperatures. Salinities, densities, and bulk compositions of fluid inclusions were calculated using the hokieflincs and algorithm of Steele-MacInnis et al. (2012) and Steele-MacInnis (2018) for the H2O−NaCl\mathrm{H}_{2} \mathrm{O}-\mathrm{NaCl} system and CO2−H2O−NaCl\mathrm{CO}_{2}-\mathrm{H}_{2} \mathrm{O}-\mathrm{NaCl} system, respectively.

4.4. Laser Raman spectroscopy

The volatile components of fluid inclusions, as well as the characterization of daughter minerals, were identified using a HR800 Laser Raman Microthermometer at the China National Nuclear Cooperation (CNNC), Beijing Research Institute of Uranium Geology (BRIUG). The excitation wavelength of the equipment was 532 nm operating at 20 mW . The spectrums were recorded at a counting time of ∼1 mm/\sim 1 \mathrm{~mm} / spectrum, with spectra ranging from 0 to 4500 cm−14500 \mathrm{~cm}^{-1} for each analysis.

4.5. Stable isotope geochemistry

The stable isotope measurements of carbon and hydrogen from fluid inclusions trapped in quartz, oxygen from quartz, and sulfur from galena and pyrite were obtained using the method of Hanbin et al. (2013) at the CNNC. Quartz grains from the three occurrences and numerous pyrite and galena grains from the Mbal occurrence were handpicked from crushed and sieved samples under a binocular microscope.

The CO2\mathrm{CO}_{2} was extracted during the reaction between the quartz samples and Cu2O\mathrm{Cu}_{2} \mathrm{O} that generated water and CO2\mathrm{CO}_{2} from the collected fluid inclusion solutions. Measurements were made using a MAT-253 mass spectrometer and reported as δ13CG-PDB \delta^{13} \mathrm{C}_{\text {G-PDB }} with analytical precision of ±0.2%\pm 0.2 \%. The water for δD\delta \mathrm{D} analysis was released from fluid inclusions by thermal decrepitation. The hydrogen isotopes were measured by reducing the water to H2\mathrm{H}_{2} over a uranium-metal-bearing tube and then transferring it to the mass spectrometer. The results for the hydrogen measurements are reported in the δ\delta notation with respect to Vienna Standard Mean Ocean Water (VSMOW) and the analytical precision is ±2%\pm 2 \%. Although we tried to collect the carbon and hydrogen isotope data from material that showed relatively low contents of secondary fluid
inclusions, the significance of these data must be viewed with caution because the analyzed solutions, nevertheless, likely represent mixtures from both primary and secondary fluid inclusions and may not be fully representative of the ore-forming fluids.

Oxygen isotope analyses of quartz were conducted on 10−20mg10-20 \mathrm{mg} of this mineral using the bromine pentafluoride method of Clayton et al. (1972) with oxygen converted to CO2\mathrm{CO}_{2} on a platinum-coated carbon rod. The results for the oxygen measurements are reported in the δ\delta notation with respect to Vienna Standard Mean Ocean Water (VSMOW) and the analytical precision is ±0.2%\pm 0.2 \%. The δ18O\delta^{18} \mathrm{O} data were collected using a gas isotope MAT-253 mass spectrometer and the δ18O\delta^{18} \mathrm{O} of the fluid was calculated using the equation of Clayton et al. (1972). The separated pyrite and galena grains were reacted with Cu2O\mathrm{Cu}_{2} \mathrm{O} until transformation to pure SO2\mathrm{SO}_{2}. The results are reported as δ34 S\delta^{34} \mathrm{~S} relative to Vienna Canyon Diablo Troilite (VCDT) and the analytical precision is ±0.2%\pm 0.2 \%.

5. Results

5.1. Wallrock alteration of gold occurrences in the Bétaré Oya district

The hydrothermal alteration halos in the country rocks surrounding the quartz veins are commonly affected by weathering that leads to the formation of supergene minerals (Fig. 5a) and some difficulty in the definition of specific alteration assemblages. Nevertheless, the hypogene alteration mineral assemblage is locally preserved in some outcrops. The alteration halos extend up to 15 m from the widest veins. The alteration halos include a proximal zone of strongly altered rock with gold enrichments for as much as 5 m outward from the veins within altered sericite schist (Fig. 5b) and a gold-free distal alteration zone within actinolite-biotite-chlorite schist (Fig. 5c). The altered sericite schist comprises an assemblage of fine-grained sericite, muscovite, quartz, minor chlorite, and sulfides (Fig. 5d). The greenschist facies actinolite-biotite-chlorite schist displays a metamorphic assemblage of actinolite, biotite, epidote, clinozoisite, garnet, biotite, plagioclase, and titanite. Where this latter lithotype is altered in the outer halo to the veins, quartz, sericite, muscovite, chlorite, calcite, and sulfides define the hydrothermal assemblage (Fig. 5e-g). The sulfide minerals disseminated in both the proximal sericite schist and distal actinolite-biotite-chlorite schist include pyrite, chalcopyrite, sphalerite, arsenopyrite, galena, and pyrrhotite. In addition, cobaltite, gersdorffite, loellingite, and pentlandite are observed in the actinolite-biotite-chlorite-schist. This latter group of minerals probably reflects elemental contributions from the metavolcanics.

5.2. Hydrothermal mineralogy of the Bétaré Oya gold district

5.2.1. Gangue minerals

The gangue minerals in the gold mineralized rocks consist of quartz, sericite, muscovite, calcite, ankerite, and barite. Quartz is the most conspicuous gangue mineral and comprises as much as 90%90 \% of the rock in the most highly mineralized zones. Some micro-grains of quartz occur as inclusions in pyrite and sphalerite. Sericite/muscovite occurs locally in or adjacent to the vein quartz and is widespread within the altered wallrocks. Carbonate minerals are calcite and ankerite. Calcite occurs in the interstices of quartz grains or as inclusions in pyrite. Ankerite, detected using SEM/EDS, is commonly associated with quartz or present as inclusions in pyrite. The ankerite is subhedral to euhedral and typically ∼60×30μ m\sim 60 \times 30 \mu \mathrm{~m} in size. Barite appears as tabular crystals or veinlets (Fig. 7n) crosscutting all other minerals and is considered either a late hypogene phase or supergene in origin.

5.2.2. The paragenetic sequence of ore minerals

Based on field observations and textural relationships, the ore minerals of the Lower Lom belt gold occurrences can be classified into two paragenetic stages, as well as into a superimposed supergene assemblage. Stage 1 is composed of pyrite (Py 1), sphalerite, galena (Gn 1),

Fig. 5. Photographs (a-c) and microphotographs (d-g) showing the mineralized veins and wallrock alteration zones. a) Profile showing quartz vein and oxidized altered wallrock; b) Profile of proximal halos of the hydrothermal altered sericite schist digging by local mining workers; and c) Sample from distal hydrothermal altered actinolite-biotite-chlorite schist; d) Association of sericite-muscovite-quartz-calcite from Fig. 5b; e-f) Hydrothermal chlorite and calcite from Fig. 5c within the quartz-plagioclase matrix; and g) Occurrence of pyrite and arsenopyrite from Fig. 5c.
chalcopyrite, pyrrhotite, arsenopyrite, electrum, wolframite, tellurides (petzite, hessite), hematite, and gold. Stage 2 consists of pyrite (Py 2), galena (Gn 2), and greenockite, with the latter phase possibly stage 3. The more common stage 3 minerals consist of hematite, covellite, goethite, and enargite. The mineral paragenesis is presented in Fig. 6.

5.2.3. Ore mineralogy

Pyrite is the dominant ore mineral and the main host mineral for gold
(Fig. 7a). It is either fragmented, massive, or present as granular aggregates in the quartz veins. The grains of pyrite display anhedral to euhedral shape with diameters ranging from tens of microns up to 1.5 cm . Pyrite, either within quartz veins or in surrounding wallrock, commonly coexists with sphalerite, chalcopyrite, and galena (Fig. 7b). Inclusions of quartz, carbonate, sericite/muscovite, wolframite, petzite, hessite, electrum, and gold have been observed in grains of pyrite (Py 1).

Sphalerite represents the second most common ore mineral within

Fig. 6. Paragenetic sequence of the Lower Lom Belt gold mineralization.
the mineralized veins. It occurs as anhedral to subhedral grains displaying brownish color. As with pyrite, sphalerite may be massive or more finely disseminated, and with individual grains ranging from 10 μm\mu \mathrm{m} to 70 mm in diameter. It is associated with pyrite, pyrrhotite, chalcopyrite, hematite, and gold (Fig. 7a, c), and commonly contains inclusions of quartz, sericite/muscovite, chalcopyrite, galena, pyrite, and gold. The exsolution microtexture of sphalerite in pyrite (Py 1) is common (Fig. 7b), as well as the exsolution of chalcopyrite in sphalerite (Fig. 7d). Some coarse grains of sphalerite are fractured and filled with quartz, galena (Gn 2), euhedral pyrite (Py 2), and greenockite.

Galena, dark-colored, is skeletal, massive, and present as veinlets within the mineralized quartz veins. Skeletal galena appears as subhedral single or aggregate crystals filling the interstices of quartz grains. The early galena (Gn 1) within the mineralized veins is associated with chalcopyrite, pyrite (Py 1), sphalerite, and gold. The coarser grains of galena exhibit perfectly developed cleavages. Post-gold, late galena (Gn 2), locally associated with greenockite, appears as veinlets crosscutting sphalerite (Fig. 7e).

Chalcopyrite occurs in the interstice of quartz grains. The mineral displays anhedral microphenocrysts to euhedral cubic microcrystals about 15μ m15 \mu \mathrm{~m} in diameter. Chalcopyrite in sphalerite is present as an exsolution texture with fine rectangular, triangular, square, and rounded shapes (Fig. 7d), which is commonly referred to as chalcopyrite disease (Barton and Bethke, 1987). Some of the fine chalcopyrite grains themselves contain inclusions of pyrite.

Tellurides (petzite, hessite), wolframite, electrum, and native gold, because of their small size, were mainly identified using SEM/EDS (Fig. 7g-i, k-l). The telluride minerals appear as elongated grains up to 15μ m15 \mu \mathrm{~m} in sphalerite (Fig. 7i, l). Electrum appears in trails of aligned inclusions in pyrite (Fig. 7h). Gold occurs as native gold, as well as in electrum and Ag-Au-Te-bearing phases. The native gold displays
subrectangular, trapezoidal, needle-like, and irregular shapes up to 50 μm\mu \mathrm{m} in maximum dimension. Gold commonly occurs as inclusions in pyrite and sphalerite. Wolframite is also observed as inclusions within pyrite and mainly is anhedral and up to 20μ m20 \mu \mathrm{~m} in diameter (Fig. 7 k ). Greenockite occurs filling fractures within sphalerite (Fig. 7j).

Iron oxide minerals consist of hematite and goethite. Hematite, in association with chalcopyrite, sphalerite, and pyrrhotite, occurs as inclusions in pyrite. The hematite also occurs, along with goethite, as fracture filling in quartz. Covellite and enargite grains, averaging 110 μm\mu \mathrm{m} in length, are formed by the replacement of chalcopyrite and sphalerite (Fig. 7f, m).

5.3. Fluid inclusion study

5.3.1. Fluid inclusions types

Three types of FI have been observed in quartz from gold-bearing veins at the three studied occurrences. They were classified into two assemblages based on petrography, microthermometry, and Laser Raman spectroscopy. Their features are summarized in Table 1. Assemblage 1 consists of type 1 and 2 primary and pseudosecondary FI, and assemblage 2 of type 3 secondary FI. These inclusions are classified as follows:
i. Types 1: aqueous-carbonic inclusions commonly contain a variable amount of CO2\mathrm{CO}_{2}. They have a relatively regular shape including negative, semi-negative, or rectangular crystal shapes. Their size ranges between 4 and 25μ m25 \mu \mathrm{~m} in diameter and they have two-phase or three-phase. The type 1 inclusions are the most representative and are found as clusters, isolated inclusions, or in trails (Fig. 8a-c). These trails occur along growth zones in the host quartz; they are distinct from healed micro-fractures. The twophase inclusions (H2O\left(\mathrm{H}_{2} \mathrm{O}\right. - and CO2\mathrm{CO}_{2}-liquid phases) are small in size commonly between 4 and 10μ m10 \mu \mathrm{~m}. Some inclusions contain a solid phase as daughter minerals (Fig. 8d). Cooling below the room temperature of many of the two-phase inclusions leads to nucleation of a third phase CO2\mathrm{CO}_{2}-vapor phase at a temperature below 20∘C20^{\circ} \mathrm{C} (Fig. 8e).
ii. Type 2: carbonic inclusions are one-phase or two-phase with a few inclusions showing a small amount of visible H2O\mathrm{H}_{2} \mathrm{O} (Fig. 8f). Their size ranges from 4 to 15μ m15 \mu \mathrm{~m}. The type 2 inclusions frequently display typical negative crystal to subrectangular shapes. Similar to type 1, type 2 inclusions often contain a daughter mineral. These inclusions occur in assemblages with type 1 FI (Fig. 8a-b, f). The one-phase carbonic inclusions nucleate a second vapor phase during cooling below room temperature.
iii. Type 3: aqueous inclusions are two-phase and between 3 and 10 μm\mu \mathrm{m} in diameter. These inclusions occur along micro-fractures crosscutting mineral boundaries and display a sub-rectangular shape (Fig. 8g). They are secondary FI and comprise inclusion assemblage 2 .

5.3.2. Fluid inclusion microthermometry

During freezing and heating experiments, temperatures corresponding to different phase changes in the FI were recorded including the melting temperatures of CO2\mathrm{CO}_{2}, ice, clathrate, CO2\mathrm{CO}_{2} homogenization, and total homogenization. These measurements were used to estimate salinities, densities, compositions, and trapping pressures and temperatures of the trapped fluids.

Analyzed representative type 1 aqueous-carbonic inclusions ( n=\mathrm{n}= 146) included 95 from the Mbal, 30 from the Haya Haya, and 21 from the Mborguéné occurrences. The melting temperatures of the carbonic phase were between -59.8 and −56.6∘C-56.6^{\circ} \mathrm{C} (Fig. 9a). The values lower than −56.6∘C(-56.6^{\circ} \mathrm{C}\left(\right. pure CO2)\left.\mathrm{CO}_{2}\right) imply the presence of a minor amount of other volatile components in the inclusions that depressed the melting temperatures of CO2\mathrm{CO}_{2}. The melting temperatures of clathrate ranged from 3.8

Fig. 7. Microphotographs and SEM/EDS spectra showing the relationship between ore minerals. a) Association of gold (Au), pyrite (Py), and sphalerite (stage 1); b) Exsolution microtexture of sphalerite and chalcopyrite in pyrite (stage 1); c) The coexistence of chalcopyrite (Ccp), pyrrhotite (Po), sphalerite (Sp), and Fe-oxide (hematite?) within pyrite (stage 1); d) Microtexture of chalcopyrite (Ccp), pyrite (Py) in sphalerite (stage 1); e) Microvein of galena (Gn 2) cutting sphalerite; f) Replacement texture of chalcopyrite (Ccp) by covellite (Cv) within pyrite (Py); g) Spectrum of native gold (contain minor Ag) as inclusion in pyrite (Py); h) Finegrained electrum included within pyrite (Py); e) Association of Ag-Au-Te telluride (petzite) and chalcopyrite (Ccp) within pyrite (Py); j) Greenockite occurring along fracture crosscutting sphalerite (Sp); k) Inclusion of wolframite within pyrite (Py); l) Inclusion of Ag-Te telluride (hessite) within sphalerite (Sp); m) Supergene enargite of stage 3; and n) Microvein of barite cutting quartz from hydrothermally altered wallrock.

Table 1
Summary of the fluid inclusion data for the three studied gold occurrences.

L410 s\mathrm{L}_{410 \mathrm{~s}} : water liquid phase; V1420\mathrm{V}_{1420} : water vapor phase; L4202\mathrm{L}_{4202} : liquid carbonic phase; V1032\mathrm{V}_{1032} : vapor carbonic phase; N2\mathrm{N}_{2} : nitrogen; CH4\mathrm{CH}_{4} : methane; nd: not detected.
to 9.1∘C9.1^{\circ} \mathrm{C}, corresponding to a range of salinities between 0.5 and 10.8 wt %NaCl\% \mathrm{NaCl} eq., with a mode clustering at 6.2wt%NaCl6.2 \mathrm{wt} \% \mathrm{NaCl} eq. (Fig. 9b). The homogenization temperatures of CO2\mathrm{CO}_{2} through vapor bubble disappearance were generally between 10 and 30∘C30^{\circ} \mathrm{C} with a mode at −24∘C-24^{\circ} \mathrm{C} (Fig. 9c). The total homogenization temperatures of type 1 inclusions commonly occurred between 280 and 360∘C360^{\circ} \mathrm{C}, with most values clustering between 300 and 320∘C320^{\circ} \mathrm{C} (Fig. 9d). The calculated compositions of XNaCl,XO2O\mathrm{X}_{\mathrm{NaCl}}, \mathrm{X}_{\mathrm{O} 2 \mathrm{O}}, and XH2O\mathrm{X}_{\mathrm{H} 2 \mathrm{O}} and bulk densities ranged from 0.01−0.030.01-0.03, 0.03−0.24,0.74−0.960.03-0.24,0.74-0.96, and 0.87−1.00 g.cm−30.87-1.00 \mathrm{~g} . \mathrm{cm}^{-3}, respectively. It is important to note that many of the type 1 inclusions with high CO2\mathrm{CO}_{2} contents decrepitated prior to total homogenization of the FI.

Type 2 carbonic inclusions (n=25)(\mathrm{n}=25) have melting and homogenization temperatures for CO2\mathrm{CO}_{2} ranging from -60.0 to −56.7∘C-56.7^{\circ} \mathrm{C} and 13 to 25∘C25^{\circ} \mathrm{C}, respectively (Fig. 9e-f). A few of these measured FI displayed a narrow meniscus that corresponds to a water phase. The melting temperatures of clathrates ranged between 5.5 and 8.3∘C8.3^{\circ} \mathrm{C}, corresponding to salinities between 3.4 and 9.3wt%NaCl9.3 \mathrm{wt} \% \mathrm{NaCl} eq. (Fig. 9g). The calculated bulk densities of the type 2 inclusions range between 0.73 and 0.85 g . cm−3\mathrm{cm}^{-3}.

Type 3 FI consist of aqueous fluids ( n=39\mathrm{n}=39 ) of assemblage 2(Mbal,n2\left(\mathrm{Mbal}, \mathrm{n}\right. =22=22; Haya Haya, n=7\mathrm{n}=7; Mborguéné, n=11\mathrm{n}=11 ). The melting temperatures of ice are between −6.7-6.7 and −0.2∘C-0.2{ }^{\circ} \mathrm{C}, corresponding to salinities of 0.2−10.1wt%NaCl0.2-10.1 \mathrm{wt} \% \mathrm{NaCl} eq. The total homogenization temperatures of this FI type are between 160 and 235∘C235^{\circ} \mathrm{C} (Fig. 9h) and their calculated densities between 0.82 and 0.98 g.cm−30.98 \mathrm{~g} . \mathrm{cm}^{-3}.

5.3.3. Laser Raman spectroscopy

Thirty representative FI of types 1 and 2 were analyzed using laser Raman spectroscopy. Type 3 inclusions were not analyzed owing to their scarcity and small size. Type 1 and 2 FI consisted of H2O\mathrm{H}_{2} \mathrm{O} and CO2\mathrm{CO}_{2} as the major components, consistent with the microthermometry, but small amounts of N2\mathrm{N}_{2} and CH4\mathrm{CH}_{4} were detected in the CO2\mathrm{CO}_{2}-phase. The solid-phase includes nahcolite (Fig. 10a-b), arsenolamprite (Fig. 10c-d), and magnesite (Fig. 10g) daughter minerals.

5.4. C-O-H-S isotopes

The δ18Oquartz \delta^{18} \mathrm{O}_{\text {quartz }}, δDfluid \delta \mathrm{D}_{\text {fluid }} and δ13Cfluid \delta^{13} \mathrm{C}_{\text {fluid }} values for the studied goldbearing quartz and δ34 S\delta^{34} \mathrm{~S} of galena and pyrite are summarized in Table 2. The δ18Ofluid \delta^{18} \mathrm{O}_{\text {fluid }} was calculated based on the oxygen isotopic compositions for quartz veins and the estimated trapping temperatures of FI of 310∘C310^{\circ} \mathrm{C} (Table 2), using the equation of Clayton et al. (1972). The δ18Ofluid \delta^{18} \mathrm{O}_{\text {fluid }} values are between +5.0%+5.0 \% and +9.5%+9.5 \% and the δDfluid \delta \mathrm{D}_{\text {fluid }} values between −50.6%-50.6 \% and −21.8%-21.8 \%. The δ13C\delta^{13} \mathrm{C} of the ore-forming fluid ranged between −5.8%-5.8 \% and −5.5%-5.5 \%. Sulfur isotope compositions are +5.3%+5.3 \% and +8.2%+8.2 \% for the galena and pyrite, respectively. Data from the Woumbou-Colomine-Ketté granite-hosted gold mineralization are also presented in Table 2 (Azeuda Ndonfack et al., 2021 and unpublished data).

6. Discussion

6.1. Composition and P-T conditions of gold-forming fluids

The first FI assemblage (FIA 1), composed of inclusion types 1 and 2, is interpreted to contain the fluids responsible for the gold mineralization within the Lower Lom Belt. Evidence supporting a genetic relationship includes: (1) The analyzed FI that appear as primary in the the gold-bearing veins are aqueous-carbonic, thus explaining the abundance of carbonate gangue; (2) Most types 1 and 2 FI have essentially the same melting and homogenization temperatures for CO2\mathrm{CO}_{2}, thus consistent with a well defined assemblage rather than indicating a series of unrelated fluid generations; (3) The inclusion characteristics of the FIA 1 are of low salinity ( ∼6.2wt%NaCl\sim 6.2 \mathrm{wt} \% \mathrm{NaCl} eq.) and such low salinity aqueous-carbonic fluids are widely reported as responsible for formation of orogenic gold deposits (Groves et al., 1998; Ridley and Diamond, 2000; Klein et al., 2000; Grandjean da Costas et al., 2019; Zobeir et al., 2019; Azeuda Ndonfack et al., 2021); (4) The disproportionate CO2−H2O\mathrm{CO}_{2}-\mathrm{H}_{2} \mathrm{O} ratios are consistent with heterogeneous fluid entrapment during unmixing and this mechanism is responsible for the majority of gold precipitation (Pal et al., 2019); and (5) The presence of an arsenic daughter mineral (arsenolamprite) within FIA 1 is significant because Au and As are commonly transported together as sulfide complexes in these crustaltype fluids (Goldfarb et al., 2005; Simmons et al., 2020). In fact, Wilkinson (2001) argued that the best evidence for a temporal genetic relationship between the ore and fluid is where a particular fluid contains daughter ore minerals. Furthermore, the Lower Lom Belt occurrence H2O−CO2−NaCl±N2±CH4FI\mathrm{H}_{2} \mathrm{O}-\mathrm{CO}_{2}-\mathrm{NaCl} \pm \mathrm{N}_{2} \pm \mathrm{CH}_{4} \mathrm{FI} are similar to those also reported as responsible for gold deposition elsewhere in the Central African Belt in Cameroon and Central African Republic (Kpeoua et al., 2020; Azeuda Ndonfack et al., 2021).

The constructed isochores for the aqueous-carbonic type 1 inclusions can help estimate the P-T trapping conditions based upon the microthermometric data (Wilkinson, 2001). Assuming final homogenization temperatures statistically approximate true trapping temperatures for type 1 inclusions, an estimate from most inclusions indicates trapping at 300 to 320∘C320^{\circ} \mathrm{C}. The construction of isochores (0.87−1.00 g.cm−3)\left(0.87-1.00 \mathrm{~g} . \mathrm{cm}^{-3}\right) for these type 1 inclusions could then provide an estimate of trapping pressures at those temperatures. In addition, it is known that the maximum temperature for the coexistence of tellurides (petzite + hessite) and gold is 313∘C313{ }^{\circ} \mathrm{C} (Bortnikov et al., 1988; Fig. 11). In our studied occurrences, petzite, hessite, and gold occur together as inclusions in pyrite and sphalerite grains (Fig. 7), therefore suggesting that the maximum temperature at which gold and tellurides were deposited was no higher than −313∘C-313{ }^{\circ} \mathrm{C}. The trapping temperature for gold mineralization was then assumed as being 310∘C310^{\circ} \mathrm{C}. This would indicate apparent trapping pressures between 1.4 and 3.5 kbar, reflecting fluctuating lithostatic to hydrostatic conditions during vein formation, similar to that argued by Boullier and Robert (1992). Assuming the lithostatic pressure to be representative of pre-faulting conditions, an apparent depth of ∼6−9 km\sim 6-9 \mathrm{~km} is estimated for the gold deposition event.

Fig. 8. Microphotographs showing petrographic characteristics of fluid inclusions. All of them were taken under the petrographic microscope at room temperature. At room temperature LCO2\mathrm{L}_{\mathrm{CO} 2} nucleated a VCO2\mathrm{V}_{\mathrm{CO} 2}-phase at T≤20C∘.LCO2=\mathrm{T} \leq 20 \mathrm{C}^{\circ} . \mathrm{L}_{\mathrm{CO} 2}= liquid carbonic phase, VCO2=\mathrm{V}_{\mathrm{CO} 2}= vapor carbonic phase, LH2O=\mathrm{L}_{\mathrm{H} 2 \mathrm{O}}= liquid aqueous phase, VH2O=\mathrm{V}_{\mathrm{H} 2 \mathrm{O}}= vapor aqueous phase, S; daughter mineral. Types 1 and 2 inclusions are characteristics of the fluid inclusion assemblage 1 , while type 3 is typical of fluid assemblage 2. a) Coexistence of type 1 and type 2 inclusions; b) Pseudosecondary type 1 and 2 inclusions; c) Cluster of type 1 fluid inclusions; d) Type 1 inclusion showing magnesite daughter mineral; e) Three-phase (LCO2+VCO2+LH2O)\left(\mathrm{L}_{\mathrm{CO} 2}+\mathrm{V}_{\mathrm{CO} 2}+\mathrm{L}_{\mathrm{H} 2 \mathrm{O}}\right) fluid inclusion; f) CO2\mathrm{CO}_{2}-rich type 2 inclusions coexisting with type 1 fluid inclusions; and g) Secondary aqueous fluid inclusions.

6.2. Sources of fluid volatile components

The data for the plot of calculated δ18Ofluid \delta^{18} \mathrm{O}_{\text {fluid }} values ( +5.0%+5.0 \% to +9.5%+9.5 \% ) versus δDfluid \delta \mathrm{D}_{\text {fluid }} measured values ( −50.6%-50.6 \% to −21.8%-21.8 \% ) partly overlap both the magmatic and metamorphic fluid fields (Sheppard, 1986; Fig. 12). These ratios are consistent with those from the granite-hosted Woum-bou-Colomine-Ketté orogenic gold district located to the south of this
study area (Azeuda Ndonfack et al., 2021; Fig. 12) and those for most worldwide orogenic gold deposits (Goldfarb et al., 2005). The uniformly heavy δ18O\delta^{18} \mathrm{O} values characterizing ore-forming fluids for such deposits (McCuaig and Kerrich, 1998; Grandjean da Costaa et al., 2019) typically range from +6.0%+6.0 \% to +11.0%+11.0 \% for Precambrian ores (McCuaig and Kerrich, 1998). However, the source of oxygen and hydrogen in the oreforming fluids is still debated due to several problems highlighted by

Fig. 9. Histograms show data of fluid inclusion microthermometry of the three studied gold occurrences. a) Melting points of carbon dioxide of type 1 inclusions ( ∼59.8\sim 59.8 to ∼56.6\sim 56.6 ); b) Salinities of type 1 inclusions calculated based on the melting temperatures of clathrate; c) Homogenization temperatures of carbon dioxide of type 1 inclusions ( 10−30∘C10-30^{\circ} \mathrm{C} ); d) Total homogenization temperatures of type 1 fluid inclusions ( 280−360∘C280-360^{\circ} \mathrm{C} ); e) Melting points of type 2CO22 \mathrm{CO}_{2}-rich inclusions ( ∼60.0\sim 60.0 to −56.6∘C-56.6^{\circ} \mathrm{C}; f) Homogenization temperatures of type 2CO22 \mathrm{CO}_{2}-rich inclusions ( 13−25∘C13-25^{\circ} \mathrm{C} ); g) Salinities of type 2CO22 \mathrm{CO}_{2}-rich inclusions calculated based on the melting points of clathrate; and h) Total homogenization temperatures of type 3 fluid inclusions ( 160−240∘C160-240^{\circ} \mathrm{C} ).

Goldfarb and Groves (2015), particularly because of the overlapping isotopic fields for magmatic and metamorphic origins. Moreover, some authors argued that measured δ18O\delta^{18} \mathrm{O} values will have undergone some shift by fluid-rock exchanges during fluid migration or at the site of fluid trapping, and thus may not be precisely indicative of original fluid source values (McCuaig and Kerrich, 1998; Bhattacharya et al., 2014; Groves and Santosh, 2016; Grandjean da Costaa et al., 2019). Nevertheless, these isotope data are consistent with those of most orogenic gold deposits and the association of the gold occurrences along the Bétaré Oya Shear Zone in metamorphosed rocks, rather than in spatial association with any particular outcropping intrusive body, leads us to
favor a metamorphic fluid source for the gold-transporting fluid in the Bétaré Oya district. The data presented here coupled with the existing data (Fig. 12) for eastern Cameroon are consistent with a metamorphic fluid forming the gold occurrences throughout the East Metallogenic Province of Cameroon and its extension into the Central African Republic.

The significant content of CO2\mathrm{CO}_{2} in FI in many gold deposits has attracted the attention of researchers and has led to significant debate over the past decades regarding the source for the non-aqueous volatile species in the lode gold hydrothermal fluids (Klein et al., 2006; Sarangi et al., 2012; Luders et al., 2015; Swain et al., 2018; Grandjean da Costaa

Fig. 10. Spectra from Laser Raman spectroscopy of vapor, liquid, and solid phases from fluid inclusions. a-b) Carbonate mineral (nahcolite) in types 1 and 2 inclusions; c-d) Ore mineral arsenolamprite in CO2\mathrm{CO}_{2}-phase of types 1 and 2 inclusions; e) Pure CO2\mathrm{CO}_{2} of type 1 inclusions; f) CO2\mathrm{CO}_{2} with dissolved nitrogen and methane in type 1 inclusions; g) Aqueous phase with magnesite daughter mineral in type 1 inclusions; and h) Spectrum of H2O\mathrm{H}_{2} \mathrm{O} liquid phase in type 1 inclusions. CO2\mathrm{CO}_{2} : carbon dioxide, N2\mathrm{N}_{2} : nitrogen, CH4\mathrm{CH}_{4} : methane.
et al., 2019). Some workers favor a mantle derivation for the CO2\mathrm{CO}_{2} (Sarangi et al., 2012; Swain et al., 2018), whereas others agree with a metamorphic origin (Luders et al., 2015). Measured δ13C\delta^{13} \mathrm{C} values between −10.0%-10.0 \% and 0%0 \% have been variably interpreted by most authors as the product of metamorphic devolatilization, crustal magmatism, or mantle degassing (Goldfarb et al., 2005). Nevertheless, McCuaig and Kerrich (1998) noted that many orogenic gold deposits are characterized by δ13Cfluid \delta^{13} \mathrm{C}_{\text {fluid }} values between −11.0%-11.0 \% and +2.0%+2.0 \%, and the carbon isotope
compositions extracted from inclusions for this study ranged from −5.8%-5.8 \% to −5.5%-5.5 \%. A similar range of δ13C\delta^{13} \mathrm{C} data ( −6.0%-6.0 \% to −5.5%-5.5 \% ) recorded from hydrothermal magnesite of meta-ultrabasic rocks in the Limpopo gold belt in South Africa was interpreted as representative of a deep-seated mantle source for the carbon (Van Schalkwyk and van Reenen, 1992; Van Reenen et al., 1994). Thus, although data are consistent with those from many orogenic gold deposits, whether such carbon was sourced simply from the metamorphic breakdown of organic

Table 2
Results of oxygen, hydrogen, carbon, and sulfur isotope measurements of the three studied gold occurrences.

T(∘C)\mathrm{T}\left({ }^{\circ} \mathrm{C}\right) : trapping temperature of the ore-forming fluid; δ18O5− SMOW (%)\delta^{18} \mathrm{O}_{5-\text { SMOW }}(\%) : oxygen isotope value measured from gold-bearing veins; δ18Ofluid (%)\delta^{18} \mathrm{O}_{\text {fluid }}(\%) : calculated oxygen isotope value of fluid based on the oxygen isotope value for gold-bearing vein and the trapping temperature of fluid inclusions, using the equation of Clayton et al. (1972).

Fig. 11. Diagram showing estimation of pressure and temperature trapping conditions using the isochores of H2O−CO2−NaCl\mathrm{H}_{2} \mathrm{O}-\mathrm{CO}_{2}-\mathrm{NaCl} type 1 fluid inclusions (calculated using the algorism of Steele-Maclinnis, 2018). Line A represents the temperature of tellurides (petzite + hessite) and gold precipitation, indicating the maximum temperature of tellurides (petzite + hessite) and gold coexistence at 313∘C313^{\circ} \mathrm{C} (Bortnikov et al., 1988). Line B represents the stability boundary of hornblende from metabasites (Elmer et al., 2006). Act (Actinolite); Chl (Chlorite); Hbl (Hornblende); Ab (Albite); Cz (Clinozoisite).
matter in the Lom Belt or whether mantle degassing played a role in oreforming event remains equivocal.

The δ34 S\delta^{34} \mathrm{~S} compositions of +5.3%+5.3 \% for galena and +8.2%+8.2 \% for pyrite from the Mbal occurrence are in the broad range that characterizes orogenic gold deposits ( −27.2%-27.2 \% to +24%+24 \%; McCuaig and Kerrich, 1998;

Fig. 12. Diagram showing oxygen and hydrogen isotopic compositions of fluid in equilibrium with ore-related hydrothermal veins in the three gold occurrences studied. Data of granite-hosted gold occurrences are from Azeuda Ndonfack et al. (2021). The δDfluid \delta \mathrm{D}_{\text {fluid }} was obtained by analyses of H2O\mathrm{H}_{2} \mathrm{O} from fluid inclusions and the δ18Ofluid \delta^{18} \mathrm{O}_{\text {fluid }} calculated using measured isotope composition of quartz and an estimated temperature of 310∘C310^{\circ} \mathrm{C}. Fields for metamorphic and magmatic water are from Sheppard (1986) and the meteoric water line is from Craig (1961).

Goldfarb et al., 2005; Chang et al., 2008). They are also consistent with δ34 S\delta^{34} \mathrm{~S} in data obtained for granite-hosted occurrences in the Woumbou-Colomine-Ketté gold district ( +6.5to+7.0%+6.5 \mathrm{to}+7.0 \%; Azeuda Ndonfack et al., 2021). Vishiti et al. (2017) suggested that δ34 S\delta^{34} \mathrm{~S} values between about

+3%+3 \% and +15%+15 \% for Bétaré Oya gold-related sulfides indicated mixing of magmatic and metamorphic sulfur. However, their analyses of samples of galena ( +2.8to+2.9%+2.8 \mathrm{to}+2.9 \% ) and pyrite ( +4.9%+4.9 \% ) from the Nkombokoro and Mborguéné occurrences were relatively similar to our measurements, whereas solely their pyrite values from Belikombone ( +14.7to+14.7 \mathrm{to} +14.9%+14.9 \% ) reflected the much heavier sulfur. The latter occurrence is about 10 km northeast of the earlier; it likely reflects a different sulfur source that could simply be a different lithologic unit (metavolcanic and metasedimentary rocks) within the belt that was devolatilized during the metamorphic event. The δ34 S\delta^{34} \mathrm{~S} values of about +3%+3 \% to +8%+8 \% for the Lower Lom Belt sulfides could be the result of S derived from magmatic fluids exsolved from some of the S-type intrusions outside of the belt (Vishiti et al., 2017). However, there is no supporting evidence to suggest any type of genetic relationship of the veins to any nearby intrusion. Furthermore, it is unclear whether a significant auriferous fluid volume could be unmixed from a crystallizing magma below our estimated depth of 9 km for vein formation; even considering a large CO2\mathrm{CO}_{2} component in the fluid exsolved from the melt, few economic gold deposits are recognized to have formed from a magmatic-hydrothermal fluid at such depths.

In summary, the participation of meteoric water was not significant during gold deposition and the ore-forming fluid was most likely a metamorphic fluid evolved in the belt between ca. 650 and 600 Ma . Such a fluid type is inherent to these types of tectono-metamorphic settings worldwide (Goldfarb et al., 2005). Based on the data obtained in this study, combined with those of Azeuda Ndonfack et al. (2021), there is no way to totally rule out a mantle carbon source, however, a crustal source is viewed as more favorable for sulfur.

6.3. Mechanism of gold precipitation

Complex hydrothermal processes are required to concentrate gold from crustal background levels into economic gold deposits. Gold transport and deposition depend on many variables that may include temperature, pressure, pH, F2\mathrm{pH}, \mathrm{~F}_{2}, water:rock ratios, mathrmCl2−concentration,andfugacityofconcentration,andfugacityofmathrmH2mathrmS˜ofthehydrothermalsystem.Thermodynamiccalculationshaveshownthatthegoldbisulfidecomplexes(suchasAuofthehydrothermalsystem.Thermodynamiccalculationshaveshownthatthegoldbisulfidecomplexes(suchasAuleft.(mathrmHS)2−right)\mathrm{Cl}_{2}{ }^{-}concentration,andfugacityofconcentration, and fugacity of \mathrm{H}_{2} \mathrm{~S}ofthehydrothermalsystem.Thermodynamiccalculationshaveshownthatthegoldbisulfidecomplexes(suchasAu of the hydrothermal system. Thermodynamic calculations have shown that the gold bisulfide complexes (such as Au \left.(\mathrm{HS})_{2}{ }^{-}\right)mathrmCl2−concentration,andfugacityofconcentration,andfugacityofmathrmH2mathrmS˜ofthehydrothermalsystem.Thermodynamiccalculationshaveshownthatthegoldbisulfidecomplexes(suchasAuofthehydrothermalsystem.Thermodynamiccalculationshaveshownthatthegoldbisulfidecomplexes(suchasAuleft.(mathrmHS)2−right)are the most suitable candidate for transporting gold in the range of temperatures of 270−400∘C270-400{ }^{\circ} \mathrm{C}, while the chloride complexes (such as AuCl2−\mathrm{AuCl}_{2}{ }^{-}) are dominant for transporting gold at higher temperatures (Mikucki, 1998). Our FI data indicate that the Bétaré Oya gold mineralization took place at a temperature of −310∘C-310{ }^{\circ} \mathrm{C} and thus we assume bisulfide was the dominant complexing agent. At such a temperature, chloride complexing would only be significant under acidic and oxidizing conditions (Simmons et al., 2020) and these conditions are not consistent with the observed mineralogy of the Bétaré Oya gold occurrences.

Gold deposition in the Bétaré Oya occurrences took place through chemical reactions that destabilized the sulfide complexes (Douring et al., 2004; Williams, 2007; Agangi et al., 2016; LaFlamme et al., 2018). The chemical processes that caused the gold deposition in the district may have involved fluid-phase unmixing and sulfidation reactions with iron-rich host rocks. The interaction of the fluid with iron-rich host rocks destabilized the gold bisulfide complexes via wallrock sulfidation. The presence of Ni - and Co-bearing sulfide minerals in mafic wallrocks confirms the importance of this process throughout the district. It helps explain the abundance of gold in hydrothermal alteration zones adjacent to gold-bearing quartz veins or as inclusions in pyrite and sphalerite (Figs. 5a-b, 7a). Simultaneously, gold deposition also occurred through fluid chemical changes caused by phase separation resulting from fluid cycling events (Boullier and Robert, 1992). Our FI investigations confirm that phase separation widely accompanied the Bétaré Oya gold event, almost certainly due to our estimated pressure fluctuation of at least 2000 bars, with resulting redox and/or pH changes leading to the decrease in gold solubility. The occurrence of free gold within the silicate matrix in the Bétaré Oya gold occurrences (Fon et al., 2012; Vishiti
et al., 2017) is evidence of the destabilization of gold bisulfide complexes during phase separation induced by pressure fluctuations. Fluid inclusion studies indicated similar drastic changes in pressure in the Woumbou-Colomine-Ketté district also located in eastern Cameroon (Azeuda Ndonfack et al., 2021).

6.4. Genetic model

Several features indicate that the primary gold occurrences of the Lower Lom Belt are consistent with the orogenic gold deposit model of Groves et al. (1998) and Goldfarb et al. (2005), such as: (1) The N- to NEtrending gold-bearing veins are structurally controlled and spatially related to deformation along the Bétaré-Oya Shear Zone; (2) Mineralized veins cut the Neoproterozoic metavolcanic-metasedimentary rocks of the Lower Lom Belt; (3) Mineralized quartz veins are late- to post-peak metamorphism as indicated by the lack of C-S structures that are welldeveloped within the metamorphic rocks, implying the veins formed at end of the D3 transpressional/transtensional deformation; (4) Hydrothermal minerals consist of quartz, calcite, ankerite, sericite, muscovite, chlorite, and sulfides and are identical to those observed in the other parts of the Central African Belt hosting orogenic gold deposits (Kpeoua et al., 2020; Azeuda Ndonfack et al., 2021); (5) The ore-forming fluid was a relatively low salinity H2O−NaCl−CO2±CH4⋅ N2\mathrm{H}_{2} \mathrm{O}-\mathrm{NaCl}-\mathrm{CO}_{2} \pm \mathrm{CH}_{4} \cdot \mathrm{~N}_{2} system; (6) Daughter minerals in FI are nahcolite, magnesite, and arsenolamprite; (7) δ18Oafluid (+5.0\delta^{18} \mathrm{Oa}_{\text {fluid }}(+5.0 to +9.5%+9.5 \% ), δDafluid (−50.6\delta \mathrm{Da}_{\text {fluid }}(-50.6 to −21.8%-21.8 \% ), δ13Cafluid (−5.8\delta^{13} \mathrm{Ca}_{\text {fluid }}(-5.8 to −5.5%-5.5 \% ), and δ34 S\delta^{34} \mathrm{~S} from galena and pyrite (+5.3%(+5.3 \% and +8.2%+8.2 \%, respectively) isotopic data of the gold-forming fluid fall into the ranges typical of orogenic gold deposits (McCuaig and Kerrich, 1998; Goldfarb et al., 2001, 2005); (8) Gold is geochemically associated with local enrichments in Ag,As,Bi,Cu,Mo,Pb,Te,W\mathrm{Ag}, \mathrm{As}, \mathrm{Bi}, \mathrm{Cu}, \mathrm{Mo}, \mathrm{Pb}, \mathrm{Te}, \mathrm{W}, and Zn (Freyssinet et al., 1989); and (9) Mafic dikes, perhaps lamprophyres, have been reported in the district studied as occur in many orogenic gold provinces worldwide and suggest deep crustal fault systems.

Based on the paleodepth of ore formation, ore mineralogy, fluid composition, P-T conditions of trapping fluid, and structural setting, the Lower Lom Belt gold occurrences clearly belong to the mesozonal orogenic gold deposit classification (Fig. 13). The mineralized veins are controlled by a well-developed sinistral brittle-ductile shear zone and their genesis is likely associated with metamorphic fluid produced during mineral reactions in medium-grade metamorphic rocks. Depending on the temperature of the more significant devolatilization reactions and the geothermal gradient, orogenic gold deposits can form over a variety of depths, from as shallow as 3 km to as deep as 20 km (Fig. 13), although it is important to note that any given deposit will not exceed 2−4 km2-4 \mathrm{~km} in vertical extent. In the Bétaré Oya district of the Lom Belt, the deposits formed near the brittle-ductile transition at temperatures slightly above 300∘C300^{\circ} \mathrm{C}. Orogenic gold typically forms during late orogenic shifts from compressional to transpressional/transtensional regimes (Goldfarb et al., 2005; Azeuda Ndonfack et al., 2021), such as is the case in the Lom Belt. Although magmatic sources are invoked by some workers for the formation of this deposit type (Vishiti et al., 2017), we see no evidence to support such for orogenic gold in eastern Cameroon. It is highly unlikely that granitoids below 9 km depth near these major shears are capable of exsolving large fluid volumes with significant amounts of gold from a crystallizing melt. Furthermore, the isotopic data collected in this study are inconsistent with a magmatic source for the various studied components. To reach the upper crust, a magmatic fluid would require a high fluid-rock ratio during upward flow along the main shear in the Lom Belt and under such flow, required large isotopic shifts in the fluid-dominant system are unlikely. Therefore, the metamorphic model is most compatible with the observed geological environment in eastern Cameroon (Fig. 13).

7. Conclusions

The results of this contribution show that economic amounts of gold

Fig. 13. Schematic representation of crustal environments of hydrothermal gold deposits in terms of paleodepth and tectonic setting (after Groves et al., 1998; Goldfarb et al., 2005) showing the approximate crustal location of formation of the Lom Belt gold occurrences. The oreforming fluids are interpreted to reflect metamorphic devolatilization below 9 km depth, fluid migration during transition to a more transpressional regime, and gold precipitation during fluid unmixing and wallrock sulfidation accompanying hydrofracturing and pressure fluctuations at about 6−9 km6-9 \mathrm{~km} depth. Magmatic and/or meteoric water contributions to the oreforming system are unlikely.
are associated with both quartz veins and altered wallrocks of the Lower Lom Belt. The mineralized veins are structurally controlled by the NErending sinistral strike-slip Bétaré-Oya Shear Zone and were deposited during the D3 deformational episode. The hydrothermal assemblage consists of quartz, sericite, muscovite, calcite, ankerite, chlorite, and sulfides. Two ore mineral assemblages are identified with the first composed of pyrite (Py 1), sphalerite, galena (Gn 1), chalcopyrite, pyrrhotite, petzite, hessite, wolframite, electrum, hematite, and gold, and the second consisting of pyrite (Py 2), galena (Gn 2), and greenockite. The supergene minerals (stage 3) include hematite, enargite, covellite, and goethite. The ore-forming fluid is composed of two FI assemblages of three inclusion types. They include assemblage 1 of H2O−\mathrm{H}_{2} \mathrm{O}- NaCl−CO2±CH4±N2\mathrm{NaCl}-\mathrm{CO}_{2} \pm \mathrm{CH}_{4} \pm \mathrm{N}_{2} (type 1) and CO2±CH4±N2\mathrm{CO}_{2} \pm \mathrm{CH}_{4} \pm \mathrm{N}_{2} (type 2) inclusions, some with arsenic- and carbonate-bearing solid phases, and assemblage 2 of H2O−NaCl\mathrm{H}_{2} \mathrm{O}-\mathrm{NaCl} (type 3) FI. The gold-associated inclusions of assemblage 1 are consistent with the Lom Belt mineralization having formed at −300∘C-300{ }^{\circ} \mathrm{C} and −6−9 km-6-9 \mathrm{~km} during fluid phase separation driven by significant pressure fluctuations and fluid-rock interaction. The δ18Ofluid (+5.0\delta^{18} \mathrm{O}_{\text {fluid }}(+5.0 to +9.5%+9.5 \% ), δDfluid (−50.6\delta \mathrm{D}_{\text {fluid }}(-50.6 to −21.8%-21.8 \% ), δ13Cfluid (−5.8\delta^{13} \mathrm{C}_{\text {fluid }}(-5.8 to -5.5 ), and δ34 S\delta^{34} \mathrm{~S} from galena and pyrite ( +5.3%+5.3 \% and +8.2%+8.2 \%, respectively) compositions suggest a metamorphic source for ore-forming fluid and sulfur.

The deformational features of the mineralized quartz veins (such as the laminated structures, subgrain boundaries, undulatory extinction), the ore mineralogy, the fluid inclusion characteristics, and O-H-C-S stable isotopic compositions indicate that the Lower Lom Belt gold
mineralization is best defined by the mesozonal orogenic gold model.

Declaration of Competing Interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

Acknowledgments

This paper is part of the Ph.D thesis of the first author and was funded by the Ministry of Science and Technology of the People’s Republic of China (State Key Research Plan, No. 2017YFC0601302). The first author received a grant from the Society of Economic Geologists to present the preliminary results of this work at the SEG-2019 conference, in Santiago, Chile. We are grateful to the traditional chiefs of the Dokayo, Mhambel, and Mborguéné villages for their kind support during the fieldwork. Special thanks to the technicians at the Beijing Institute of Uranium and Geology Beijing GeoAnalysis Co. Ltd. for their help during Laser Raman spectroscopy and C-H-O-S stable isotope analysis. XiaoYu, CuiKai, and YuChang are thanked for their support during SEM/EDS analysis and fluid microthermometric measurements. The authors are grateful to the two anonymous reviewers for theirs constructive comments and suggestions that significantly helped improve the quality of the manuscript. We sincerely acknowledge the Editor-in-Chief Prof. Franco Pirajno and

the Associate Editor Dr. Lebing Fu for their suggestions and editorial handling

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