Evaporites as gemstones

Many natural gemstone settings are not evaporite-associated, the following section discussed those that are. Interestingly, the common link between most natural gem-forming environments is a hydrous, typically saline to hypersaline, solution that is able to precipitate a crystal in a void from a fluid that contains elevated and unusual levels of particular constituents, including chromophores; hence pegmatites, volcanics and meta-evaporites are the most typical hosts for natural gemstones.


Precious stones and gems are rare by definition and need exceptional geologic conditions that give rise to gem-quality materials. These conditions may include the availability of sometimes uncommon major constituents, the presence of adequate chromophores, limited concentrations of undesirable elements, open space, an environment conducive to forming crystals of sufficient size and transparency.


This general statement of requirements to form a precious stone also encapsulates why some gems have meta-evaporitic associations. Halite and most other evaporite salts disappear or transform into other mineral phases by the early greenschist phase. As this happens the dissolution/transformation releases a pulse of hot basinal chloride waters that can leach/carry elements such as beryllium, chromium and vanadium (chromophores) from any adjacent organic-rich shales. Trace elements also tend to be enriched in the more evolved depositional brines that precipitated the latter minerals in any evaporite precipitates. At the same time as halite dissolves or transforms, anhydrite layers and masses typically remains into the more evolved portions of the metamorphic realm (amphibolite-granulite facies). The volume loss associated with the dissolution/transformation of meta-evaporites facilitates the formation of open space (sometimes pressurized) as veins and fractures that then allow the free growth of the precious stone or gem material.


Lapis lazuli

Lapis lazuli is made up not of a single mineral but an accumulation of minerals; it is mostly composed of lazurite (Na,Ca)8(AlSiO4)6(S,SO4,Cl)1-2), typically 30-40%. Lapis gemstones also contains calcite (white veins), sodalite (blue), and pyrite (gold flecks of colour). Dependent on metamorphic history and protolith chemistry, other common minerals in lapis include; augite, diopside, enstatite, mica, haüyanite, hornblende, and nosean. Some specimens also contain trace amounts of the sulphur-rich mineral lollingite (var. geyerite). Lazurite is a member of the sodalite group of feldspathoid minerals. Feldspathoids have chemistries that are close to those of the alkali feldspars, but are poor in silica. If free quartz were present at the time of formation it would react with any feldspathoid precursor to form a feldspar.


Natural lazurite contains both sulphide and sulphate sulphur, in addition to calcium and sodium, and so is sometimes classified as a sulphide-bearing haüyne. Sulphur gives lazurite its characteristically intense blue colour, it comes from three polysulphide units made up of three sulphur atoms having a single negative charge. The S3- ion in the sulphur has a total of 19 electrons in molecular orbitals and a transition among these orbitals produces a strong absorption band at 600 nm, giving a blue colour with yellow overtones. The intensity of the gem’s blue is increased with increasing sulphur and calcium content, while a green colour is the result of insufficient sulphur.


Other members of the sodalite group include sodalite and nosean. Sodalite is the most sodium-rich member of the sodalite group and differs from the other minerals of the group in that its lattice retains chlorine. Interestingly, sodalite can be created in the laboratory by heating muscovite or kaolinite in the presence of NaCl at temperatures of 500°C or more. In the literature, the commonly accepted origin of lazurite is through contact metamorphism and metasomatism of dolomitic limestone. Such a metasedimentary system also requires a source of sodium, chlorine and sulphur; the obvious source is interbedded evaporites in the protolith, as is seen in plots if its molecular constituents.


Colombian emeralds are meta-evaporitic gems formed via mixing of fault-focused hypersaline basinal solutions released during the subsurface dissolution and alteration of evaporites. Emeralds occur within carbonate-silicate-pyrite veins and breccias hosted in black shales and limestones, tied to occurrences of hydrocarbons/bitumens. Gem fields constitute two zones, the eastern and western emerald zones, some 80 km apart in the Eastern Cordillera of Columbia (Figure). The Columbian emerald fields supply more than 60% of the world emerald market.

Columbian emeralds


The western zone crops out in the core of the Villeta anticlinorium (A). The sedimentary series that enclose the emerald deposits (B) are a few hundred meters thick and are composed of, from bottom to top: (1) micritic, largely dolomitic limestones (Rosablanca Formation, Valanginian-Hauterivian); (2) calcareous black shales (Hauterivian); and (3) siliceous black shales that form the base (Hauterivian) of the thick mudstones (Barremian-Aptian) of the Paja Formation. Most of the emeralds are found in hydrothermal breccias or in carbonate-pyrite veins developed within both dolomitic limestones and calcareous black shales.


The deposits are hectometer-sized at most and display numerous folds, thrusts, and tear faults. All the tectonic contacts are marked by centimeter-to meter-thick hydrothermal breccias that are cataclasites with clasts of black shales and albitites (i.e., massively albitized shales) within a carbonate-albite-pyrite cement. These breccias derived from a fluid-rich pulp. Part of the overpressured fluids escaped and triggered intense hydraulic fracturing in surrounding rocks, especially along tear faults. In each deposit, there is evidence of complex deformation that resulted in polyphase duplex structures. For example, in the Coscuez deposit, the  ore-hosting N30°E verging folds and thrusts were a response to movement on the sinistral N20°E trending Coscuez tear fault (C).

In the eastern zone (D), Andean thick-skinned tectonics are responsible for the main deformation observed on regional cross sections through the Eastern Cordillera and adjacent Llanos foothills. This Andean deformation corresponds mainly to reverse faults (with an overall vergence to the southeast) and folds affecting the Palaeozoic basement and its Cretaceous-Tertiary sedimentary cover. Some of these faults are inverted Early Cretaceous growth faults. The Esmeralda fault likely represents the preserved part of an Early Cretaceous normal fault (D). As the result of Andean thick-skinned tectonics, the Chivor emerald deposit is located on a gently northwest dipping monocline situated on the western flank of a large, N30°E-trending, upright fold, devoid of cleavage.


The enclosing sedimentary series (E) correspond to the upper part of the Guavio Formation (Berriasian), which unconformably overlies the Palaeozoic basement and is overlain by the shales of the thick (2900 m) Valanginian Macanal Formation. The series hosting the emerald-bearing veins and associated hydrothermal breccias are composed of, from bottom to top: (1) shales and siltstones that are locally massively albitized (lower albitites); (2) a 1–10-m-thick, stratiform brecciated level, largely made of hydrothermal breccia after a former evaporite level; (3) an albitized and carbonatized sequence (upper albitites) that is white and initially contained anhydrite beds (as evidenced by phantom nodules, chicken-wire, and tepee structures); and (4) bioherms of micritic or shelly limestones grading vertically and laterally into black shales intercalated with calcareous pebbly mudstones and olistostromes.

Within the brecciated level, disrupted blocks of the hanging wall (albitites, black shales, limestones) and caving structures are evidence for the collapse of the roof. Thus, the dissolution of an evaporitic horizon (probably initially dominated by halite) appears to be a major process controlling the formation of the brecciated level. The stratiform association of evaporites, limestones, albitites, and the brecciated level is consistent across tens of kilometers around the Chivor deposit, owing to its folding by the Andean phase. All the emerald deposits of the Chivor mining district are located within or just above this regional level, and so define a stratigraphic emerald horizon (F,G).

Asian Rubies

Emeralds grew hydrothermally during thrusting and folding of black shales in a salt lubricated rauhwacke system. Highly reactive hypersaline brines, generated by circulation of hot basinal waters passing through and dissolving halokinetic evaporite beds, created chloride-rich brines that passed from thrust planes to drive wall-rock metasomatic alteration. By this process, Be, Cr and V were mobilised, and could then precipitate emeralds in extensional veins and hydraulic breccias (rauhwacke), which are also rich in calcite and pyrite.



Ruby is the red-pink transparent gem form of the mineral corundum (2[Al2O3]), with intensity of its red colour related to trace amounts of chromium oxide. The coloured gem varieties of corundum include; white sapphire (colourless), ruby (red), pigeon’s blood (deep red), sapphire (blue), golden sapphire (yellow), green sapphire (green), as well as olive green, purple and violet (violet sapphire). Corundum is found as highly valued rubies in Myanmar and Thailand, and as sapphires in Sri Lanka and Cashmere-India, as well as in lesser quality gem deposits in Australia, Tanzania, Cambodia, Russia, Thailand, Malagasy, Nepal, Malawi, and the USA.

Marble-hosted ruby deposits today represent the most important source of coloured gemstones in Central and South East Asia. These deposits are located in the Himalayan belt and formed during the Tertiary collision of the Indian plate in its northward movement into the Eurasian plate (H). SE Asian rubies are spatially related to granitoid intrusions and are hosted in evaporite-entraining platform carbonates that have experienced high-grade metamorphism. All occurrences are located close to major tectonic features active during Himalayan orogenesis, either directly in suture zones in the Himalayas, or in shear zones that guided extrusion of the Indochina block after the collision in South East Asia. Ar-Ar dating of micas syngenetic with the rubies, and U-Pb dating of zircon inclusions in the rubies, show that these deposits formed during Himalayan orogenesis and associated extensional tectonics active between the Oligocene and the Pliocene, in a belt from Afghanistan to Vietnam. These ruby-bearing marbles lie in the amphibolite facies (T = 610 to 790° C and P ≈ 6 kbar).


Fluid inclusions in the various gem rubies from the Jegdalek, Hunza and northern Vietnamese deposits indicate retrograde metamorphism; 620<T<670° C and 2.6<P<3.3 kbar (I). Whole rock analyses of non-ruby-bearing marbles show they contain enough aluminium and chromiferous elements to produce all the ruby crystals that they host. As protolith, these carbonates contained Al- and chromiferous-bearing detrital materials, probably clays, which were deposited on the platform as the same time as the mostly carbonate matrix along with locally enriched levels of organic matter.


In addition, the consistency of (C, O)-isotopic values of carbonate in the marbles shows marbles developed in a closed fluid system, and were not infiltrated by externally-derived fluids. The carbon isotopic composition of graphite in the marbles underlines its organic origin and that it exchanged C-isotopes with the adjacent carbonates during its metamorphic evolution. Earlier, less evolved examples of this isotopic exchange, with temperatures still in the upper diagenetic realms are seen in Permian Carbonates of the Saraburi region, Thailand (Warren et al., 2014). Moreover, the O-isotopic composition of ruby across SE Asia was buffered by metamorphic CO2, released during devolatilisation of marble, while the H-isotopic composition of associated mica is consistent with a metamorphic origin for water now in equilibrium with the micas. The boron isotopes from tourmalines associated with the rubies indicates a likely nonmarine evaporite source for the boron (J).


Ruby-bearing marbles from Nangimali, in the Azad-Kashmir deposits of Pakistan, contain, besides phengite, atypical mica intergrowths with paragonite, phlogopite and aspidolite (sodium phlogopite). Both phlogopites, although intimately linked and coexisting with paragonite, are fluorine rich, in contrast to the phengite and paragonite. The phengite is either associated with phlogopite or can be isolated. The presence of aspidolite in these ruby-bearing marbles, together with NaCl solid inclusions and the presence of anhydrite, further strengthen the notion that evaporites were involved in the genesis of these gem corundums.


In summary, Southeast Asian rubies formed during the retrograde portion of the metamorphic path, mainly by destabilization of muscovite or spinel. The metamorphic fluid system was rich in CO2 released from devolatilisation of carbonates, and in fluorine, chlorine and boron, probably released from molten and altering Cretaceous or Precambrian and even Permian salt masses (NaCl, KCl, CaSO4) caught up in the orogeny. Evaporites are key to explaining the formation of rubies and emeralds in this gem-rich region of southeast Asia. Molten salts mobilized the in situ Al and metal transition elements that were contained in the host marbles, leading to crystallization of the rubies.

Tsavorite and Tanzanite

Tsavorite, the vanadian variety of green grossular garnet, is a high value semi-precious gemstone. Currently it is hosted exclusively in metasedimentary and typically meta-evaporitic formations within the Neoproterozoic Metamorphic Mozambique Belt. Deposits are mined in Kenya, Tanzania and Madagascar and other occurrences in the metamorphic belt are located in Pakistan and East Antarctica. All are located within metasomatized graphitic rocks, such as graphitic gneiss and calc-silicates, intercalated with meta-evaporites. Tsavorite is found as primary deposits either in nodule (type I) or in quartz vein (type II), and in placers (type III). The primary mineralization (types I and II) are controlled by lithostratigraphy and/or structure (K).

For the African occurrences, the protoliths of the host-rocks are either anhydrite or baryte, both hosts were deposited at the beginning of the Neoproterozoic within a saline marine coastal sabkha environment, deposited at the margin of the Congo–Kalahari cratons. Subsequently, during the East African–Antarctican orogeny, these saline sediments experienced high amphibolite to granulite facies grade metamorphism, with the formation of the tsavorite occurring between 650 and 550 Ma. Nodules of tsavorite formed during prograde metamorphism, with calcium supplied by metamorphosing sedimentary sulphates and carbonates, whereas the alumina, silicates, vanadium and chromium come from a clay-chlorite protolith (K). Veins formed during deformation and shearing of the metasedimentary platform were partially filled with tsavorite and other metamorphic phases. Further metasomatism occurred subsequent retrograde metamorphism with ongoing reactions with pore waters.


All the metasedimentary sequences today are still characterized by the presence of evaporitic minerals, such as gypsum and anhydrite, and scapolite. Evaporites are essential to tsavorite as they provide calcium and facilitate the mobilization of all the chemical elements for tsavorite formation. The H2S–S8 metamorphic fluids characterized in primary fluid inclusions of tsavorites and the δ11B values of coeval dravite consistently around -20‰ confirm the marine evaporitic (high salinity) origin of the fluids. The V2O3 and Cr2O3 contents of tsavorite range respectively from 0.05 to 7.5 wt.%, while their δ18O values are in the range of 9.5–21.1‰.


The genetic model proposed for tsavorite is metamorphic, based on chemical reactions developed between an initial assemblage composed of gypsum and anhydrite, carbonates and organic matter deposited in a sabkha-like sedimentary basin.. Two kinds of nodules were likely precursors to tsavorite and both were hosted in an evaporitic protolith: (i) nodules Type NI were initially anhydrite concretions within silica-rich shales; (ii) nodules Type NII were initially baryte concretions located within calcareous shales. Both types of shales contained V(–Cr)-rich clays and organic matter (K).


At the beginning of the prograde metamorphism (stage 2), the hosting shales turned into schists, and V(–Cr)-rich clays and organic matter transformed into respectively V(–Cr)-rich micas and graphite. For the crystallisation of tsavorite in nodules Type NI, Si and Al come from the schist, and V and Cr from the clays and transported into the anhydrite core where it forms tsavorite utilising Ca from the evaporitic sulphate, following the equation;


3CaSO4 + 2Al3 + 3SiO2 + 6H2O -> Ca3Al2(SiO4)3+6O2+3H2S+6H+.


H2S is trapped in fluid inclusion cavities during the growth of tsavorite, while the sulphur is expelled into the schist to form pyrite. At the end of the prograde metamorphism, most of the anhydrite had been replaced by tsavorite, which is also present as small crystals scattered in the evaporitic rims. During retrograde metamorphism, these crystals are rehydrated and change in the rims into vanadian zoisite.

For the crystallisation of tsavorite in Type NII nodules, Ca and Mg came from the carbonates and Si is transported towards the central part of the nodule, replacing baryte with diopside. A minor amount of vanadium is sourced from the micas and had been incorporated into the diopside core. Ba and S were expelled into the schist, the latter forming pyrite. Then the diopside core increased until replacing all the baryte. Transport of Al, V and Cr from the schist into the core led to the formation of tsavorite following the reaction;


 3CaMgSi2O6 + 2Al3+ --> Ca3Al2(SiO4)3 + 3SiO2+3Mg2+


Tsavorite and diopside interacted then to form scapolite following the reaction;


2CaMgSi2O6 + 2Ca3Al2(SiO4)3 + 2SiO2+2CO2 + 4Al2O3 --> 2Ca4Al6Si6O24CO3 + 2MgO.

Hence, primary deposits of tsavorite are of two types, both with predictable occurrence positions in the field, which can be used to construct useful exploration paradigms (L). Type I nodules occur in meta-evaporite beds creating stratiform units, intercalated within the metamorphic series and more precisely in Ca-dominated rocks such as carbonate gneiss and calc-silicates with intercalations of marble (as in Kenya, Madagascar, part of Tanzania and Antarctic deposits). Type II associations occur where the metamorphic formations are affected by tight ductile isoclinal folding and shearing. The tsavorite occurs in quartz veins located at the hinges of the sheared isoclinal folds (Type IIA) and within ‘saddle reef’ structures (Type IIB). In this situation, deformation and shearing are associated with fluid circulation and percolation in the tectonically-opened structures. The metamorphic rocks undergo metasomatic phenomena such as pyritisation, graphitisation, silicification and carbonatisation, sometimes with precipitation of tsavorite (Type IIC). This type of deposit is found in the Lelatema area and Ruangwa in Tanzania, and in Pakistan.


Secondary tsavorite deposits (Type III) result from the erosion of the primary ones. Three secondary (reworked) sub-types are described: (i) sub-type IIIA corresponds to eluvial deposits such as those found in the Merelani area and in Kenya; (ii) sub-type IIIB is colluvial deposits such as those located in the Merelani area and in Kenya; (iii) sub-type IIIC is alluvial deposits forming the placers of Tunduru and Umba in Tanzania.


During retrograde metamorphism, hydration of both tsavorite and scapolite can lead to the formation of zoisite (4[Ca2Al3(O/OH/SiO4/Si2O7)]), a member of the epidote group, which can become another gem mineral, known as Tanzanite. In the Merelani region of Tanzania, veins containing tanzanite and tsavorite garnet have resulted from pegmatitic and hydrothermal fluids reacting with medium- to high-grade meta-evaporitic rocks along the crests of folds (L). The faceted crystal form of many of the tanzanites is a result of their growth into dissolution cavities in the anhydrites and gypsums, fed by undersaturated mineralizing solutions and free of interference from other minerals. The zone of mineralization with such excellent gems is only a few meters wide but extends for about 9 km and includes other spectacular gem growths such of chrome tourmaline and chrome diopside. To date the known occurrences of gem quality tanzanite are restricted to the Merelani area.

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