Geochemical and reflectance spectroscopy data integration to characterize emerald deposits: the case of the Paraná deposit, Brazil.

The Paraná emerald deposit is one of the few occurrences of emerald, a rare beryl variety, in Borborema Province, northeastern Brazil. We characterized the Paraná deposit by combining field geology, petrography, whole-rock geochemistry, mineral chemistry, and reflectance spectroscopy. The mineralization is associated with phlogopite-, actinolite-phlogopite-, and phlogopite-phengite schists, mylonitic gneisses, and several acidic rocks (e.g. granitic pegmatites/aplites, quartz ± feldspar veins) along the Portalegre Shear Zone. Emerald can be found in quartz-feldspar and aplite veins and veinlets interleaved with phlogopite- or actinolite-phlogopite schists, or within the foliation of the schists. The presence of albitites and the compositional variation of the schists suggest a metasomatic origin for emerald with variations of the metasomatic process. All these different lithotypes can be readily identified through reflectance spectroscopy especially in the range of 2,150-2,450 nm, where the main mafic minerals show absorption features related to Al-OH (phengite), and Fe-OH and Mg-OH bonds (phlogopite/actinolite). Our study shows that possible mineralized phlogopite schists can be distinguished from other sterile rocks, although point spectral analysis does not separate emerald-bearing phlogopite schists from schists without emerald due to the dominance of major phlogopite absorption features rather than emerald features.


INTRODUCTION
Knowledge about gem-quality mineral deposits within the Borborema Province, northeastern Brazil, was concentrated for years in the Seridó Pegmatite Province (SPP, previously known as the Borborema Pegmatite Province; Scorza 1944), since the First World War, due to its large amount of industrial minerals (especially mica), until the end of the Second World War, as a result of the search for strategic materials such as tantalum and beryllium minerals (Silva et al. 1995, Santos et al. 2014).
Over the last two decades, a few other Brazilian gemological districts have been recognized outside the SPP, such as the Solonópole-Quixeramobim Pegmatite District (Ceará state;Vidal & Nogueira Neto 2005), the Cristais-Russas Pegmatite District (Ceará state; Vidal & Nogueira Neto 2005), and the Extreme Southwest Gemological District (Rio Grande do Norte state; Moraes 1999). The latter is known for amazonite and dark blue aquamarine mineralizations associated with granitic pegmatites in the Vieirópolis Pegmatite Field (Barreto et al. 2016), as well as emerald-bearing schists within the so-called Paraná -Marcelino Vieira -Francisco Dantas emerald belt (Moraes 1999).
This emerald belt stands out as one of the few occurrences of this rare dark green variety of beryl in Borborema Province (cf. Schwarz 1987, Giuliani et al. 1990, Zwaan et al. 2012, Santiago et al. 2018. The best emerald crystals have been recovered from the vicinity of Paraná city, specifically from the localities of Pitombeiras and Aroeira, which comprise the Paraná emerald deposit. Several technical reports have described emeralds from the Pitombeiras and Aroeira mines since the 1980s (e.g. Vasconcelos 1984, Moraes 1999, Medeiros 2008, Souza 2017. Araújo Neto et al. (2019) characterized the mineralogy and gemology of the emerald crystals, but detailed studies on the geology remain unveiled.
In this paper, we present a comprehensive geological description of the Paraná deposit by combining field geology, petrography, wholerock geochemistry, mineral chemistry and reflectance spectroscopy data to understand the relations between the main lithotypes and emerald mineralization.
Reflectance spectroscopy is a nondestructive method that studies the interaction of electromagnetic radiation with a material's surface in the visible to the short-wave infrared spectral range. For rocks and minerals, this interaction can unveil variations on the chemical composition and crystalline structure of each investigated sample. In the visible-near infrared range (VNIR: 400-1,000 nm), electronic transition processes commonly occur in transition metals (e.g. Fe, Cr, Ni, Co;Clark 1999, Thompson et al. 1999), generating diagnostic absorption features. Also, molecular vibrational processes (e.g. H-O-H, Al-OH, Fe-OH, C-O) cause absorption features in the short-wave infrared range (SWIR: 1,000-2,500 nm) (Clark 1999, Thompson et al. 1999. In general, the use of reflectance spectroscopy has been well-stated as an exploration method for metal deposits related to hydrothermal alteration systems (Bedell et al. 2009, Kerr et al. 2011, Swayze et al. 2014, Ramakrishnan & Bharti 2015. However, its applications are still very uncommon in gemstone deposits (cf. Turner et al. 2017).
Emerald mineralization is commonly related to mafic/ultramafic rocks containing ferroanmagnesian mica and amphibole. The emerald absorption features associated with electronic transitions in Cr 3+ and/or V 3+ are responsible for its green coloration (Wood & Nassau 1968). On the other hand, ferroan-magnesian mica and amphibole species can be identified by spectral responses in specific wavelength intervals in the VNIR-SWIR range due to their contents of Fe, Mg, Al and OH - (Pontual et al. 2008).
In this context, we introduce the use of reflectance spectroscopy data as a tool for the characterization of emerald host and associated rocks in a case study in NE Brazil.

Regional geological setting
The Borborema Province, located in northeastern Brazil, is a complex mosaic-like region folded during the Brasiliano Cycle (Almeida et al. 1981). This province covers an area of approximately 400,000 km², and it consists of local exposures of Archean nuclei and large outcrops of Paleoproterozoic gneiss-migmatite basement rocks covered and/or interleaved with Meso-to Neoproterozoic domains marked by supracrustal sequences. All these rock units were intensely affected by regional metamorphism and magmatism of the Brasiliano-Pan African orogeny (800-500 Ma, Brito Neves et al. 2014).
Many authors have conventionally subdivided this province into several domains or sub-provinces based on its structural configuration allied with geochronological and The Rio Grande do Norte Sub-province is a Rhyacian to Orosirian crustal block bordered to the south by the Patos Lineament and to the west by the Senador Pompeu Shear Zone (Brito Neves et al. 2000). According to Brito Neves et al. (2000) and Angelim (2006), three main domains compose the RNS: the São José do Campestre (SJD), the Rio Piranhas-Seridó (RPD), and the Jaguaribeano (JGD) domains. The studied area is situated between the Jaguaribeano (west) and Rio Piranhas-Seridó (east) domains, which are limited by the Portalegre Shear Zone (Figure 1b, Brito Neves et al. 2000), a crustal discontinuity that extends from the Patos Lineament to south to the Potiguar Basin to the north.
The Paraná emerald mineralization occurs within the Portalegre Shear Zone (Figure 1b), in a discontinuous trend with, at least 20 km length. The emerald crystals are found mostly inside quartz-feldspar and aplite veins and veinlets interleaved with phlogopite schists. These schists occur as lenticular bodies along the mylonitic fabric of the metavolcanosedimentary unit of the Caicó Complex, which also comprises spatially associated gneisses, amphibolites and quartzites (Moraes 1999, Araújo Neto et al. 2018, 2019.

MATERIALS AND METHODS
Lithological and structural data were collected from several outcrops in the Paraná emerald deposit, mainly from the Pitombeiras and Aroeira mines. The relation between schist and host rocks could be better studied in the Pitombeiras region due to a high number of shafts and trenches.
An Olympus BX51 microscope coupled with an Olympus DP26 camera system was employed for petrographic studies of 67 thin sections.
For whole-rock geochemical characterization, four schist samples were crushed and powdered in the laboratory facilities of the Stable Isotope Laboratory of the Nucleus of Geochemical Studies (NEG-LABISE) at the UFPE. These powders were sent to the Bureau Veritas Mineral Laboratories in Vancouver (Canada) for bulk chemical analyses. After lithium borate fusion, major element concentrations were obtained by inductively coupled plasma emission spectrometry (ICP-ES), while minor and trace elements were determined by inductively coupled plasma mass spectrometry (ICP-MS). Mineral compositions were determined for nine samples by electron probe micro-analysis (EPMA) using a JEOL JXA-8230 instrument in the Electron Microprobe Laboratory (LASON) of the University of Brasília (UnB). The microprobe is equipped with a scanning electron microscope (SEM), an optical microscope, and a CCD video camera for imagery, five wavelength dispersive X-ray spectrometers (WDS) for quantitative chemical analysis, and one energy dispersive X-ray spectrometer (EDS) for qualitative and/ or semi-quantitative analysis. The system was operated for standard silicate analysis, using LASON's internal standards for multistandard calibration: albite (Na), forsterite (Mg), topaz (F), microcline (Al, Si and K) andradite (Ca and Fe), vanadinite (Cl and V), MnTiO 3 (Ti and Mn), Cr 2 O 3 (Cr) and NiO (Ni). The following parameters were used: accelerating voltage of 15 kV, sample current of 10 nA, beam diameter of 1 μm, and a counting time on the peak of 10 s. Optical microscope images from the CCD video camera and backscattered electron images from SEM were used for selecting analysis spots and for avoiding mineral inclusions.
Reflectance spectroscopy was performed for 88 samples of different lithotypes, using a FieldSpec4 ® Standard Resolution spectroradiometer (Analytical Spectral Devices) from the University of Campinas (UNICAMP). The spectroradiometer records spectra over 2,151 channels, with wavelengths ranging from 350 to 2,500 nm, comprising the VNIR (350-1,000 nm) and SWIR ranges (1,000-2,500 nm). The spectral sampling (bandwidth) is 1.4 nm for the 350-1,000 nm range, and 1.1 nm for 1,001-2,500 nm (Malvern Panalytical 2020). A contact probe with an internal light source and a 20 mm spot size was used, and a white reference (Spectralon®) was employed for the instrument calibration. The samples were measured at least three times and an average reflectance spectrum was calculated for each sample.
The continuum removal technique (hull quotient) was applied for highlighting the absorption features, providing better visualization of the shape, symmetry, and depth of the absorption features centered at specific wavelengths. This technique consists of normalizing the reflectance spectrum using a mathematical function that defines a convex hull, which fits the spectrum curve. The continuum removal is the ratio of the reflectance values by the continuum line (Clark & Roush 1984, Clark et al. 2003.

The Paraná emerald deposit
In the Pitombeiras mine, the basement rocks are grey-colored biotite gneiss of the Caicó Complex, fine-to medium-grained, with a protomylonitic to mylonitic texture reflecting the intense deformation of the Portalegre ductile strike-slip shear zone. The mylonitic gneisses host lenses of mafic schists, commonly phlogopite schists and actinolite-phlogopite schists, that occur as vertical to sub-vertical bodies within the mylonitic foliation with a NE-SW trend, varying from N15E to N70E (Figure 2a).
The schists are metric-sized, greenishblack colored, fine-to medium-grained, with a lepidoblastic or nematolepidoblastic texture. They host several recrystallized veins and veinlets with granitic compositions and aplitic to pegmatitic textures. These granitic bodies occur as small boudins composed of quartz ± potassium feldspar ± plagioclase, and minor biotite ( Figure 2b). Emerald crystals are found (i) within the schistosity planes, (ii) inside the small granitic bodies, or (iii) at the schist-vein contact (Figure 2c). At the regional scale, several granite pegmatites occur along the Portalegre Shear System. In the Pitombeiras region, centimeter-to meter-sized granite pegmatites occur as lenses and veins in association with the emeraldbearing schists (Figure 2d). These pegmatites are often composed of massive potassium feldspar with minor quartz and muscovite, and rarely garnet. Additionally, in the main Pitombeiras shaft, a meter-sized albitite dike occurs adjacent to the basement gneisses ( Figure 2e) and the host schists. These albitites are tabular sheared pegmatite bodies, composed of albite and minor pale green muscovite.
In the Aroeiras region, the emerald has been mined within a single shaft, specifically from a gallery 16 meters deep. The phlogopite schist is N30E trending with a dip varying from 60° to 24° to SE and occurs interleaved with quartzfeldspar pegmatite that sometimes seems boudinated. Strongly fractured biotite gneisses of the Caicó Complex host both pegmatite and schist ( Figure 2f).

Petrography
Most of the basement rocks are biotite gneisses with protomylonitic to mylonitic textures. On the other hand, the schists can be petrographically divided into three types: phlogopite schist, actinolite-phlogopite schist, and rare phlogopite-phengite schist. No significant differences between samples from the Pitombeiras and Aroeira mines were recorded, although phlogopite-phengite schists were found exclusively outcropping near the main Pitombeiras shaft. The igneous bodies that occur as dikes, veins and veinlets comprise fundamentally quartz and/or feldspar compositions to alkali feldspar granites.
The plagioclase crystals are anhedral to subhedral and typically twinned on the albite law. Altered plagioclases are very common, with sericite formation along the twinning planes ( Figure 3a). Anhedral microcline with tartan twinning and perthitic microcline can occur as megacrysts presenting plagioclase with myrmekite around its edges. The interstitial quartz crystals are anhedral with undulatory extinction or vermicular forming myrmekite intergrowths ( Figure 3b).
Biotite occurs as oriented subhedral lamellar crystals and shows pleochroism from greenish-brown to pale yellowish-brown. The minor minerals are titanite, apatite, and allanite, which are anhedral to subhedral, with the allanite showing epidote rims and anomalous interference colors due to metamictization ( Figure 3c). Epidote varies from small euhedral crystals to anhedral crystals when bordering allanite cores. The opaque minerals show subhedral irregularly shaped rhomb-like sections. Hornblende is rare and occurs as euhedral to subhedral crystals with dark green to bluish-green colors; sections with two cleavage traces are common.
Mylonitic to protomylonitic facies are characterized by quartz and feldspar neoblasts (Figure 3d). Recrystallization and neoblast formation through subgrain rotation are also frequent.

Phlogopite schists
The phlogopite schists are black colored, fine-to coarse-grained, with texture ranging from lepidoblastic to granolepidoblastic. These schists are composed of phlogopite (Figure 4a), making up about 75% of the total volume, and subordinate quartz, microcline, and plagioclase, which can occur as sigmoids with protomylonitic texture due to intense recrystallization and subgrain formation ( Figure  4b). Minor minerals are apatite and rare emerald crystals. The phlogopite crystals are brown to yellow colored and euhedral to subhedral in shape. Quartz, microcline and plagioclase occur in variable proportions but, generally, represent 20-25% of the total volume. Quartz is anhedral with undulatory extinction and often occurs as subgrains. Microcline and plagioclase are anhedral and may be present as coarse crystals with lobe-shaped grain boundary and intense recrystallization at the edges. Apatite exhibits subhedral shape in prismatic and basal sections ( Figure 4a). The emerald is subhedral and occurs mainly as basal sections oriented parallel to the schistosity ( Figure 4c).

Actinolite-phlogopite schists
The actinolite-phlogopite schists are greenishblack colored, medium-to coarse-grained, with nematolepidoblastic texture. The main mineralogy is phlogopite and actinolite with interstitial quartz and feldspar. Minor minerals are apatite, titanite, and allanite. A few samples also show epidote and emerald as minor constituents. In general, actinolite-phlogopite schists show intrafolial folds and crenulation cleavage in the form of chevron folds, evidenced by deformed phlogopite lamellae with symmetrical limbs (Figure 4d).
There is great variation in the phlogopite and actinolite proportions in each sample. The average content of phlogopite is 54%, with a 16-76% range. Phlogopite crystals are eu-to subhedral, with pale yellow to brown color. Actinolite content averages 24% (3-62%) and comprises euhedral to subhedral prismatic crystals with a faint pleochroism (pale green to colorless), which suggests low Fe contents. Some actinolites show relic phlogopite crystals (Figure 4e). Quartz and feldspars represent about 21% of the total volume; they are anhedral and occur as stretched or minute grains due to recrystallization. Weathered samples often show sericitized alkali feldspar. Apatite occurs as short and prismatic subhedral crystals. Titanite and allanite are rare and anhedral.

Phlogopite-phengite schists
The phlogopite-phengite schists are relatively rare rocks, fine-to coarse-grained, with lepidoblastic texture. Phengite, phlogopite, quartz, and feldspar are the main minerals. Phengite is anhedral, colorless and comprises about 57% of the rock. Some relic phlogopite crystals are observed (Figure 4f). The phlogopite is subhedral, brown to pale yellow colored, representing about 34% of the volume of the rock. Quartz and feldspar are anhedral and represent 9% average of the total volume. These rocks comprise quartz, potassium feldspar, and plagioclase, and can range from pure quartz compositions to granitic ones or even plagioclase dikes (albitites, Figure 5a). They are grouped here due to their few mineralogical variations, leucocratic and acidic natures. The minor minerals are frequently biotite and muscovite, and/or emerald in some cases (Figure 5b), associated with quartz-feldspar boundins and veinlets within phlogopite schists. Garnet is rare and can be observed in some potassium feldspar veins (Figure 5c) or in the albitite. In common, these dikes, veins and veinlets often contain altered feldspar ( Figure  5d) and protomylonitic texture represented by recrystallization and grain boundary migration of quartz and feldspar (Figure 5e and 5f).

Whole-rock geochemistry
The whole-rock geochemistry analyses of four samples of the phlogopite schist are shown in Table I  Beryllium contents average 144.5 ppm but reach up to 300 ppm in sample EM21, which contains millimetric emerald crystals.

Micas
Several dark-colored and light-colored mica crystals were analyzed by EPMA from six samples of phlogopite schist, one sample of phlogopitephengite schist, and one sample of albitite (Tables II and III).
In the emerald-hosting schists from Paraná Deposit, the dark-colored micas are phlogopite and ferroan phlogopite, according to the octahedral Fe tot +Mn+Ti-Al VI versus octahedral Mg-Li binary diagram proposed by Tischendorf et al. (2001) (Figure 6) (Table II).
The light-colored dioctahedral micas from the phlogopite-phengite schist are ferroan muscovites (Figure 7), a nomenclature approved by the International Mineralogical Association  (Table III).

Amphiboles
One thin section of a representative amphibolebearing schist (sample EM88) from the Pitombeiras mine was also chemically analyzed by EPMA (Table IV) (2012) for calcium amphiboles, these magnesium-rich amphiboles are in the field of tremolite (Figure 8). In this case, however, the term "actinolite" is more suitable since it was retained for petrological reasons in the Hawthorne et al. (2012) classification for amphiboles of the tremolite-ferro-actinolite series with a compositional range that extends  from Mg < 4.5 apfu and Fe 2+ > 0.5 apfu to Mg ≥ 2.5 apfu and Fe 2+ ≤ 2.5 apfu.

Reflectance spectroscopy
The spectral characterization was carried out using representative samples from the Pitombeiras and Aroeira mines. Phlogopite-, actinolite-phlogopite-, and phlogopite-phengite schists from the Pitombeiras mine show broad absorption features in the VNIR and SWIR range (400-1,600 nm) due to the presence of Fe 2+ in the structure of ferromagnesian micas and amphiboles. Subtle absorption features at ~900 nm are also attributed to Fe 2+ , while features near 700 nm are related to Fe 3+ -Fe 2+ charge transfer (cf. Hunt 1977, Dyar 1990, Pontual et al. 2008) ( Figure  9a). For the Pitombeiras' schists, the SWIR region is marked by a subtle absorption at ~1,395 nm, associated with OHand/or water molecular vibrational processes, and an asymmetric absorption feature at 1,910 nm due to the vibration of water molecules (cf. Hunt 1977, Clark 1999, Pontual et al. 2008) (Figure 9a). The 2,150 to 2,450 nm range shows the best diagnostic features for each schist type, reflecting the mineralogical composition (Figure 9b). The absorption feature at ~2,250 nm occurs in all schist samples due to Fe-OH vibrational processes in both phlogopite and/ or actinolite (Hunt 1977, Clark et al. 1990, Pontual et al. 2008. The phlogopite schists are also marked by absorptions at ~2,330 and ~2,388 nm related to the Mg-OH bond in phlogopite (Clark et al. 1990, Pontual et al. 2008). In the actinolitephlogopite schists, the spectral signature of actinolite is predominant and it is marked by the double Mg-OH feature at ~2,297 and ~2,320 nm, although a secondary Mg-OH feature at ~2,388 nm can also be observed (Pontual et al. 2008). In the phlogopite-phengite schists both Mg-OH features of phlogopite are masked by the presence of typical phengite Al-OH absorption at ~2,350 nm and 2,450 nm (Scott & Yang 1997). These schists are also marked by an expressive absorption feature at ~2,218 nm associated with Al-OH bonds diagnostic for white micas like phengite (Scott & Yang 1997, Pontual et al. 2008). On the other hand, Al-OH absorption bands can be expected as well in weathered phlogopite schists due to the presence of clay minerals and/or sericite (e.g. double OHfeatures at ~1,390 and ~1,412 nm, and double Al-OH features at ~2,167 and ~2,208 nm in sample EM21, likely related to the presence of kaolinite; Hunt 1977, Pontual et al. 2008 (Figure 9b). Spectral interpretations of other rocks from the Pitombeiras mine show that the broad Fe 2+ absorption feature of biotite in the VNIR region allows for the swift distinction between biotite gneisses and rocks of essentially quartzfeldspar compositions (Figure 10a). In the SWIR range (Figure 10b), biotite gneisses show absorptions at ~2,250 nm (Fe-OH) and ~2,330 and ~2,384 nm (Mg-OH) related to biotite/phlogopite (Hunt 1977, Clark et al. 1990, Pontual et al. 2008). The absorption feature at ~2,200 nm (Al-OH) is common due to potassium feldspar alteration to sericite or clay minerals (Pontual et al. 2008).
The analysis of an orange-colored potassium feldspar vein shows atypical Al-OH absorption (2,205 nm) and OHand/or water absorptions at 1,438 and 1,929 nm (Figure 10b). These absorption features likely originated from alteration processes such as sericitization. In the VNIR region, a deep absorption feature at 440 nm is probably related to the presence of Fe 3+ in the feldspar crystal structure, which is responsible for the orange color (cf. Faye 1969) (Figure 10a).
Also, a sample of muscovite-bearing albitite is marked by the spectral signature of muscovite (Hunt 1977, Pontual et al. 2008, with a deep Al-OH absorption feature at 2,195 nm, and secondary ones at 2,343 and 2,433 nm ( Figure  10b). The presence of deep OHand/or water absorptions at 1,407 and 1,910 nm suggests a contribution from illite and/or sericite, common products of weathering.
In the Aroeira mine, the spectral analysis of a lithological profile parallel to the dip direction of both schist and pegmatite (Figure 11a), shows the influence of Al-OH, Fe-OH, and Mg-OH absorption features for characterizing the different lithotypes of the emerald deposit. In the SWIR region (Figure 11b), the mylonitic biotite gneiss (sample EM111) shows the relatively deep Fe-OH (~2,250 nm) and Mg-OH (~2,333, ~2,388 nm) absorption features of biotite, and a subtle Al-OH absorption at 2,200 nm which is probably due to some alkali feldspar alteration. This feature at 2,200 nm is deeper in sample EM112, a quartz-feldspar pegmatite near the contact with the biotite gneiss, probably due to a higher concentration of altered alkali feldspar (Figure 11b). On the other hand, a quartz-feldspar pegmatite sample (EM113) at the contact with the phlogopite schists, shows a spectral signature dominated by the phlogopite absorption features of Fe-OH and Mg-OH bonds. The Al-OH feature at 2,200 nm disappears in the phlogopite schist (sample EM114) since phlogopite concentrations subdue those of altered feldspar. However, if the phlogopite schist contains alkali feldspar veinlets and/or megacrysts (sample EM115), it may be characterized by well-pronounced absorption features derived from Fe-OH, Mg-OH, and Al-OH molecular vibrational process ( Figure  11b).
The emerald-bearing schist spectra Araújo Neto et al. (2019) previously investigated rough emerald crystals spectra from the Paraná deposit. In brief, these emeralds are marked by Cr 3+ absorption features in the VNIR range at ~430 nm and ~630 nm, and a broad Fe 2+ absorption band at ~850 nm. In the SWIR range, the emeralds show water absorption features at ~1,150, ~1,410, ~1,900 nm, and other four non-identified features at 2,072, 2,158, 2,205, and 2,329 nm.
The spectra of emeralds that are within the host schists show a mixture of spectral responses from different minerals ( Figure 12). Sample ML01 contains several prismatic emeralds in a quartzfeldspar vein inside the phlogopite schist. Spectral analysis in these crystals shows sharp and well-defined absorptions features typical for the expected emerald spectrum. Nevertheless, if the emerald crystal is hosted within the schist foliation, the Cr 3+ and water absorptions appear attenuated by the broad absorption of Fe 2+ in phlogopite. In this case, phlogopite absorption features are well marked at 2,252 nm (Fe-OH), and 2,328 and 2,388 nm (Mg-OH), and the diagnostic emerald features are hard (if not impossible) to distinguish. This attenuation occurs even in large centimetric emerald crystals with few mica flakes at the surface, such as observed in sample ML05 (Figure 12).

DISCUSSION
The emeralds from the Paraná deposit are associated with quartz-feldspar veins and veinlets interleaved with basic schists (average of 48.96 wt.% SiO 2 ) that are rich in magnesium (average of 13.33 wt.% MgO) and iron (average of 8.56 wt.% FeO). The schist mineralogical composition varies from the common phlogopite schist to actinolite-phlogopite schist, and phlogopitephengite schist, all hosted by mylonitic gneisses within the Portalegre Shear Zone. The absence of ultrabasic and/or metaultrabasic rocks in this region (e.g. dunites, pyroxenites, talcserpentinites, and serpentinites) suggests that the metabasic lenses of amphibolite, which occur in proximity to the Portalegre Shear Zone, are the best candidates to be the source of iron and chromium in emerald (Souza 2017). This hypothesis is corroborated by the high iron and low chromium and vanadium content in the Paraná phlogopite schists. This pattern is similarly reflected in the emerald crystals of the Paraná deposit, which contains higher Fe, and lower Cr and V in comparison to other deposits associated with ultrabasic rocks (Araújo Neto et al. 2019). However, until the development of this research, the field relations between amphibolite and schist could not be observed since all amphibolite lenses were found a few kilometers far from the Pitombeiras and Aroeira mines.
The pegmatite injections across the Portalegre Shear Zone are probably the source of Be, and the presence of albite-rich pegmatites suggests metasomatic modification (desilication) of previous granitic pegmatites (Walton 2004), with Si leaving the pegmatite to probably form metasomatic schists and the numerous quartz veins. Desilication provides a relative Al enrichment in the pegmatite that results in recrystallization of Al-rich minerals such as muscovite and garnet (Černý 2002). Additionally, the compositional variation of schists, from phlogopite-to phlogopitephengite-and actinolite-phlogopite schists, with phlogopite relicts in both actinolite and phengite, implies possible variations in the metasomatic process. Nevertheless, the spatial distribution of these schists is not in welldefined zones, but an erratic disposition across the deposit. Emerald crystals are found in both phlogopite and actinolite-phlogopite schist, but so far, no phlogopite-phengite schist has contained emerald.
Reflectance spectroscopy proved to be a powerful tool on the investigation of the different lithotypes that compose the Paraná emerald deposit. Because of its rapid and nondestructive properties, reflectance spectroscopy can quickly distinguish mineralogical variations among basic schist samples. Phlogopite-, phlogopite-phengite-, and actinolite-phlogopite schists can be recognized mainly by distinctive absorption features in the 2,210-2,400 nm spectral range, unveiling both mineralogical content and chemical composition of these rock types (Figure 13a).
The relation between the wavelength of minimum reflectance and the MgO and FeO contents (wt.%) for representative schists is shown in figures 13b and 13c, respectively. The minimum reflectance for phengite-bearing schist is at 2,218 nm (Al-OH), while pure phlogopite schists have the minimum reflectance centered at 2,330 nm (Mg-OH), shifting to 2,325 nm with the increasing of MgO content in phlogopites ( Figure 13b). Lastly, the actinolite-bearing schists also have a minimum reflectance associated with Mg-OH vibrational process but centered at 2,315 nm (Figure 13a).
Samples of emerald crystals hosted by schist could not be appropriately identified using reflectance spectroscopy due to the attenuation of the VNIR absorption features of emerald by the broad iron absorption of phlogopite and amphibole. This attenuation restricts the role of the point spectral analysis as a prospective tool in the identification of emerald signatures on local and regional scale. However, reflectance spectroscopy can still be used as a prospective tool in the investigation and identification of the host basic schists.

CONCLUSIONS
The geological setting of the Paraná emerald deposit shows characteristic features of the classic schist-type deposit (see Sinkankas 1981, Schwarz 1987, Walton 2004, Groat et al. 2008, with the formation of emerald-bearing quartz-feldspar veins and veinlets hosted in phlogopite-and actinolite-phlogopite schists probably due to metasomatic interaction between granitic pegmatites and (meta)basic/ ultrabasic rocks. This interaction and the mobilization of Be and chromophores elements (Cr, V and Fe) most-certainly happened during a Neoproterozoic strike-slip shearing represented by the Portalegre Shear Zone.
The presence of muscovite albitite and the schist compositional variation are indicative of fluid interactions and metasomatic modification (desilicated pegmatites). The uncommon high iron, and low chromium and vanadium content in Paraná schists and emerald crystals, associated with the absence of (meta)ultrabasic rocks in the region suggest a metabasic source of chromophore elements, with the Caicó amphibolite lenses as the main candidates to be this source.
Despite being an unusual method for emerald exploration, reflectance spectroscopy proved to be useful in the spectral characterization and individualization of the different lithotypes found in the deposit, especially using the SWIR range of 2,200-2,400 nm, where the main mafic minerals show absorption features derived from Al-OH bonds (phengite), and Fe-OH and Mg-OH bonds (phlogopite and actinolite).
Typical potential host rocks such as phlogopite schists and actinolite-phlogopite schists can be readily distinguished using reflectance spectroscopy from sterile rocks like phlogopite-phengite schists even though by visual analysis they are almost identical. Still, the point spectral analysis is not able to fully distinct emerald-bearing phlogopite schists from phlogopite schists without emerald due to masking of emerald absorption features by major phlogopite features. This masking can be surpassed in future evaluations by changing point spectral analysis for hyperspectral imaging.
Notwithstanding this restriction, spectral reflectance studies can be performed by using systematic reflectance and imaging spectroscopy in drill cores, as well as high resolution airborne or orbital optical sensors to target potential host rocks.