Salty Matters

The Blog is written by me, John Warren. Once every three or four weeks or so I will post an article or two on an evaporite topic that has piqued my interest. On the Saltwork Publications webpage (under "the Works") there is a growing library of pdfs and epubs based on these blogs. These articles on the website have much higher resolution extractable graphics in than in the blog. There is also a link to this set of pdfs and epubs on the home page (www.saltworkconsultants.com).

Lapis Lazuli: A metamorphosed evaporite

John Warren - Friday, November 13, 2015

Introduction 

Precious stones and [1]gems are rare by definition; hence need exceptional geologic conditions to give rise to gem-quality materials. A nexus across most natural gem-forming environments is the requirement for hydrous typically saline to hypersaline solutions, apt to precipitate euhedral crystals in a void or a pressure shadow, from fluids that contain elevated and unusual levels of particular constituents, including chromophores; hence pegmatites, volcanics and meta-evaporites are commonplace hosts for natural gemstones. Fluids promoting the growth of gem-quality crystals typically include the availability of uncommon major constituents, along with the presence of adequate chromophores [2] , as well limited concentrations of undesirable elements. Another need is open fluid space in an environment conducive to growing 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.

We know that depositional units of evaporite salts typically disappear or transform into other mineral phases by the early greenschist phase (Warren, 2016). As this happens the dissolution/transformation releases a pulse of hot basinal chloride waters that can carry gold and base metals (topics for another blog) as well as leaching and carrying elements such as beryllium, chromium and vanadium (chromophores) from adjacent organic-rich shales. Trace elements also tend to be enriched in the more evolved depositional brines that precipitate in later minerals in a primary evaporite precipitative series. Later, at the same time as halite dissolves or transforms on entry into the metamorphic realm, anhydrite layers and masses typically remain 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 fluid space (sometimes pressurized) in veins and fractures so favoring sites that then allow the free growth of precious stones and euhedral ruby, tourmaline and emerald gemstones.

I am not saying all precious stones and gems are associated with evaporites, many natural gemstone settings are not, but the lapis lazuli of Afghanistan, the melon-sized perfect rubies of Myanmar, the prolific emerald fields in Columbia, and the tsavorite deposits of east Africa likely are (Garnier et al., 2008; Giuliani et al 2005; Feneyrol et al., 2013). In this article I will focus on lapis lazuli, for a discussion of other semi-precious stones and gems that are meta-evaporites see Chapter 14 in Warren (2016).


 

Lapis lazuli

Lapis Lazuli is the metamorphic remnant of a sodic-rich, quartz-absent, evaporite mineral assemblage. It is composed of an accumulation of minerals not a single mineral (unlike rubies and emeralds); it is mostly lazurite (Na, Ca)8(AlSiO4)6(S, SO4,Cl)1-2), typically at levels of 30-40%. Lapis gemstone also contains calcite (white veins), sodalite (blue), and pyrite (gold flecks of color). 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 (Table 1). 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 have reacted with any feldspathoid precursor to form feldspar not lazurite. Natural lazurite contains both sulphide and sulphate sulphur, in addition to calcium and sodium, and so is sometimes classified as a sulphide-bearing haüyne (Figure 1). Sulphur gives lazurite its characteristically intense blue color, which 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 strong blue color, with yellow overtones. The intensity of the gem’s blue is increased with increasing sulphur and calcium content, while a green color is the result of insufficient sulphur (O’Donoghue, 2006, p. 329).

 

Other members of the sodalite group include sodalite and nosean (Table 1). 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 of its molecular constituents (Aleksandrov and Senin, 2006).


Lapis lazuli from the Precambrian of Baffin Island, Canada (Figure 1), and from Edwards, New York, are meta-evaporites with evaporite remnants (anhydrites) remaining in the same series, as are the lapis lazuli deposits at Sar-e-Sang in the Kokcha valley, Afghanistan and the lapis deposits in Liadjuar-Dara region (“River of Lazurite”) at an altitude of 5000 m in the Pamir Mountains, Tajikistan (Webster, 1975). Throughout history its bright blue color has made lapis, mostly from Sar-e-Sang, a valued gem commodity. First mined 6000 years ago, the Sar-e-Sang lapis was transported to Egypt and present day Iraq and later to Europe where it was used in jewelry and for ornamental stone[3]. Europeans even ground down the rock into an expensive powdered pigment for paints called “ultramarine”.

Lapis deposits in Lake Harbor on Baffin Island and in the Edwards Mine, New York, were produced by high-grade metamorphism of a sulphate-halite-marble protolith (Hogarth and Griffin, 1978). The anhydrites preserved near Balmat are remnants of this sequence. On Baffin Island the two main lapis lazuli lenses, some 1.6 km apart, lie at the structural top of two sequences of dolomitic marble, the thicker lens being approximately 150 m across (Figure 1b). The elongation of both lenses parallels the local layering and foliation and shows a well-developed layering parallel to the regional foliation, giving additional evidence of its sedimentary protolith to the deposits. The Main and Northern bodies constitute diopside–lazurite rocks of variable gem quality and are localized in marbles among biotite gneisses. The Main (Southern) occurrence is as long as 170 m and 6 m thick. In these deposits, sheets of high-quality lazurite (up to 1 m thick) contain variable amounts of relict diopside and plagioclase, as well as newly formed haüyne, nepheline, or phlogopite. The quantitative proportions of these minerals define the color of the rock, which changes to a more intense blue with heating. The Northern occurrence (25×36 m in size) is less rich than the Main occurrence and consists of small (no more than 1 m) lenses showing disseminated lazurite, which imparts a bluish green color to the polished surface of the rock. Chlorine and sulphur in the various lazurites, accessory pyrite, and pyrrhotite were derived from metamorphosed gypsum-, anhydrite-, and evaporitic-carbonate protoliths (Hogarth and Griffin, 1978).


In the Lake Baikal lazurite occurrences, there is once again a strong association between marble of the Perval’na Group and lazurite occurrence (Figure 2a). For example, the Slyudyanka deposit is hosted in diopside skarns and spinel–forsterite calciphyres, developed from metamorphically-evolved evaporitic dolomites (Aleksandrov and Senin, 2006). The Slyudyanka deposit shows clearly pronounced metasomatic zoning, which was associated with the prograde magnesian skarn stage and was overprinted by retrograde postmagmatic assemblages, that formed together with lazurite-bearing rocks under the influence of saline alkaline S–Cl-bearing hydrothermal solutions. These solutions also caused microclinization of blocks of leucocratic granite with the formation of lazurite in the some of the inner skarn zones. Potassium solutions caused phlogopitization of the host rocks.

Likewise, scapolite and magnesian whiteschists are typically saline mineral phases in the classic deposits of the Sar-e-Sang District (Figure 2b; Faryad, 2002). There, the lapis is composed of a combination of lazurite, diopside, calcite and pyrite and occurs in beds and lenses up to 4 meters thick within a scapolitic magnesian-marble skarn near the center of the Hindu Kush granitic massif. It is typically interlayered with, or forms veins and lenses within a gneissic and pegmatitic host. Lens-shaped lodes are typically hosted in orthoclase–microcline–perthite hornfels containing albite and quartz (Figure 2b). Lazurite bodies at the Sar-e-Sang deposit are associated with diopside metasomatites bearing nepheline, pale blue haüyne, and blue lazurite, and some lazurite-rich zones can contain up to 40-90 vol% lazurite. The rocks also contain diopside, haüyne, afghanite, and nepheline, as well as disseminated pyrite replaced by pyrrhotite. Pockets of near pure lapis lazuli can be up to 40m across and occasionally up to a meter.

Lapis lazuli in the North Italian Mountains of Colorado occurs in impure marbles in a meta-evaporitic skarn near the contact with the Eocene-age quartz monzonite and quartz diorites of the Italian Mountain stocks (Hogarth and Griffin, 1980; Mauger, 2007). There, near vertical Pennsylvanian black shales and carbonates along the west margin of the intrusive have been converted to phlogopite-diopside-andalusite hornfels and scapolite-diopside skarns with minor analcime. Compared to Sar-e-Sang, lapis in this skarn deposit is of inferior quality. It forms as deep blue lazurite granules in fine-grained forsterite-Ti phlogopite-calcite skarns and calcite marbles with diopside, Ti-phlogopite and pyrite. The hosting sediments (Mississippian limestones and Devonian sandstones) define along the NE margin of the pluton, while the NaCl came from dissolution of once nearby halite or dissolution-derived saline surface waters and shallow groundwaters moving south from the Eagle Basin.

High quality lapis is also mined from a limestone-granite skarn contact in the Chilean Andes (3500 m elevation) in the headwaters of the Cazadero and Vias River, Ovalle, Coquimbo, Chile. The lapis there is good quality, although somewhat paler than Sar-e-Sang and, like the Baikal lapis deposits of Russia, is associated with wollastonite not diopside, making it a less attractive gem. The Chilean lapis occurs in an association of phlogopite, sodalite, calcite and pyrite (Coenraads and Canut de Bon, 2000).

Meta-evaporites in the Sar-e-Sang region of Afghanistan exhibit mosaic equilibria across small volumes (in the cm3 range) within a talc-kyanite schist (whiteschist) host. The microscale mineral variations are characterized by variations in mineral assemblages conventionally attributed to vastly different pressure/temperature conditions during regional metamorphism.

On the basis of petrographic and microprobe studies, these assemblages are attributed to three consecutive stages of metamorphism of a chemically exceptional rock with a composition that falls largely into the model system MgO-Al2O3-SiO2-H2O (Figure 3; Schreyer and Abraham, 1976). Stage 1, typified by Mg chlorite-quartz -talc and some paragonite, was followed during stage 2 by talc-kyanite, Mg [4]gedrite-quartz and the growth of large dravites (magnesian tourmalines). Microprobe analyses of the phases, gedrite and talc, indicate variable degrees of sodium incorporation into these phases according to the substitution NaAl—>Si. In stage 3, pure Mg cordierite formed with or without corundum and/or talc, and the kyanite was partly converted into sillimanite. Pressure and temperature during this final stage of metamorphism was near 5-6 kb and 640°C.


Schreyer and Abraham (1976) concluded that chemical variations in the metamorphic fluids were generated by progressive metamorphism and mobilization of an evaporite deposit. Relict anhydrite and gypsum(rehydrated anhydrite) still occur in the Sar-e-Sang area. Whiteschists and the associated lapis lazuli deposits of the region are part of a highly metamorphosed evaporitic succession. Salts have largely vanished due to ongoing melting and volatilizations. The preservation of the three stage succession of mineral assemblages, across such small scales and yet related to each other through isochemical reactions, means that the main factors governing the metamorphic history of this whiteschist were compositional changes of the coexisting fluids with time. Under this scenario any pressure-temperature variations were subordinate and the chemistry of the fluids evolved as the evaporites underwent metasomatic alteration.

The sedimentary pelitic layers of this precursor evaporitic sequence first underwent a period of metamorphism in which fluid pressures approached lithostatic (stage 1). Subsequently at higher metamorphic grades, with the beginning of mobilization of the salts, the metamorphic fluids became increasingly enriched in ions such as Na+, Mg2+, Cl-, SO42-, BO33-, etc., so that water fugacity dropped considerably. This period is represented by stage 2 of the whiteschist metamorphism and was characterized by strong metasomatism that led, for example, to the growth of dravite and the amphibolite, gedrite. The physical and chemical character of stage 3 is less clearly defined. Kyanite/sillimanite inversion requires an increase in temperature or a decrease in pressure, or both; but changes in the composition of a coexisting gas phase may have played an additional role in the formation of cordierite.

Unlike classic metamorphic associations, the meta-evaporite-derived assemblage in Afghanistan may in a single thin section entrain mineral assemblages that conventionally would be assigned to the greenschist facies, the hornfels facies, and to a high pressure (amphibolite) regime. The assemblages are in effect mosaic equilibria that reflect changes in fluid composition generated from a metamorphosing evaporite pile over time and only to a lesser degree by regional evolution of total temperature and pressure. Once again, evaporites generate unusual responses compared to the general responses of metasediments.

In a refinement paper discussing the likely relationships between evaporites and whiteschists, Franz et al., 2013 note that whiteschist mineral assemblages are stable up to pressures of more than 4 GPa, but may already form at pressures of 0.5 GPa. Their formation largely depends on the composition of the protolith and requires elevated contents of Al and Mg as well as low Fe, Ca, and Na contents, as otherwise chloritoid, amphibole, feldspar, or omphacite are formed instead of kyanite or talc. They go on to note that the stability field of a whiteschist mineral assemblage strongly depends on XCO2 and fO2: at low values of XCO2, CO2 binds Mg to carbonates strongly reducing the whiteschist stability field, whereas high fO2 enlarges the stability field and stabilizes yoderite [Mg(Al,Fe3+)3(SiO4)2O(OH)].

The scarcity of whiteschist is not necessarily due to unusual P–T conditions, but to the restricted range of suitable protolith compositions and the spatial distribution of these protoliths: (1) continental sedimentary rocks and (2) hydrothermally and metasomatically altered felsic to mafic rocks. They argue continental sedimentary rocks that may produce whiteschist mineral assemblages typically have been deposited under arid climatic conditions in closed evaporite basins and may be restricted to relatively low latitudes. These rocks typically contain large amounts of palygorskite and sepiolite. Franz et al., (2013) conclude whiteschist assemblages typically are only found in settings of continental collision or where continental lacustrine fragments were involved in subduction.

In my opinion, the mosaic signature of the precursor mineral phases in the typical Sar-e-San lapis lazuli is a metamorphically-evolved response to the combination of precursor permeability and stability contrasts typical of variably-cemented halite mosaic sediments in what were likely haloturbated and variably cemented saline continental lacustrine precursors.

References

Aleksandrov, S., and V. Senin, 2006, Genesis and composition of lazurite in magnesian skarns: Geochemistry International, v. 44, p. 976-988.

Coenraads, R., and C. C. de Bon, 2000, Lapis Lazuli from the Coquimbo Region, Chile: Gems & Gemology, v. 36, p. 28-41.

Faryad, S. W., 2002, Metamorphic Conditions and Fluid Compositions of Scapolite-Bearing Rocks from the Lapis Lazuli Deposit at Sare Sang, Afghanistan: Journal of Petrology, v. 43, p. 725-747.

Feneyrol, J., G. Giuliani, D. Ohnenstetter, A. E. Fallick, J. E. Martelat, P. Monié, J. Dubessy, C. Rollion-Bard, E. Le Goff, E. Malisa, A. F. M. Rakotondrazafy, V. Pardieu, T. Kahn, D. Ichang'i, E. Venance, N. R. Voarintsoa, M. M. Ranatsenho, C. Simonet, E. Omito, C. Nyamai, and M. Saul, 2013, New aspects and perspectives on tsavorite deposits: Ore Geology Reviews, v. 53, p. 1-25.

Franz, L., R. L. Romer, and C. Capitani, 2013, Protoliths and phase petrology of whiteschists: Contributions to Mineralogy and Petrology, v. 166, p. 255-274.

Garnier, V., G. Giuliani, D. Ohnenstetter, A. E. Fallick, J. Dubessy, D. Banks, H. Q. Vinh, T. Lhomme, H. Maluski, A. Pecher, K. A. Bakhsh, P. Van Long, P. T. Trinh, and D. Schwarz, 2008, Marble-hosted ruby deposits from Central and Southeast Asia: Towards a new genetic model: Ore Geology Reviews, v. 34, p. 169-191.

Giuliani, G., A. E. Fallick, V. Garnier, C. France-Lanord, D. Ohnenstetter, and D. Schwarz, 2005, Oxygen isotope composition as a tracer for the origins of rubies and sapphires: Geology, v. 33, p. 249-252.

Hogarth, D. D., and W. L. Griffin, 1978, Lapis lazuli from Baffin Island; a Precambrian meta-evaporite: Lithos, v. 11, p. 37-60.

Mauger, R. L., 2007, Contact metamorphism-metasomatism associated with the latest Eocene northern Italian Mountain granite intrusion, Gunnison County, Colorado: Abstracts with Programs - Geological Society of America, v. 39, p. 394.

O'Donoghue, M., 2006, Gems; Their Sources, Descriptions and Identification (6th Edition): Amsterdam, Elsevier, 873 p.

Schreyer, W., and K. Abraham, 1976, Three-stage metamorphic history of a whiteschist from Sar e Sang, Afghanistan, as part of a former evaporite deposit: Contributions to Mineralogy & Petrology, v. 59, p. 111-130.

Von Rosen, L., 1990, Lapis lazuli in archaelogical contexts, in P. Aströms, ed., Studies in Mediterranean Archaeology and Literature, v. 93, Partille, Sweden.

Warren, J. K., 2016, Evaporites: A compendium (ISBN 978-3-319-13511-3) Released Dec. 2015: Berlin, Springer, 1600 p.

Webster, R., 1975, Gems, their sources, descriptions and identification: London, Newnes, Butterworths.

Wood, J., 1841, John Wood. A Personal Narrative of a Journey to the Source of the River Oxus by the Route of the Indus, Kabul, and Badakhshan, Performed under the Sanction of the Supreme Government of India, in the Years 1836, 1837, and 1838. Avaialble as Elibron Classics, 2001, 458 pages. Replica of 1841 edition by John Murray, London.



[1] The Ancient Greeks, distinguished between precious and semi-precious stones; similar distinctions were made in other ancient cultures. In modern usage, the precious stones are diamond, ruby, sapphire and emerald, with all other gemstones, including lapis lazuli, being semi-precious.

[2] Chromophore; is the part of a gem lattice responsible for its color. A gem’s color arises when a molecule absorbs certain wavelengths of visible light and transmits or reflects others. It is a structural feature in the lattice indicative of the presence of a gem-specific electron configuration of the ions in its crystal lattice; such as transition metal ions (Cr, V or Fe) occupying several different coordination sites. For example, ferrous iron (Fe2+) or ferric iron (Fe3+), the ferrous ion in peridot causes the green color and ferric ion causes the yellow color in chrysoberyl. This color effect has important uses in heat-treatment gemstones such as blue color in heat-treated sapphire.

[3] Lapis is the Latin word for “stone” and lazuli is the genitive form of the Medieval Latin lazulum meaning blue, which was taken originally from the Persian lāžaward, the name of a place where lapis lazuli was mined. Taken as a whole, lapis lazuli originally meant “stone of Lāzhward.” With time, the name of the place came to be associated with the stone mined there and, eventually, with its bright blue color.

Lapis lazuli’s use as jewelery can be traced back to the 5th millennium B.C.E. with the discovery of beads at a cemetery outside the temple walls of Eridu (Sumer) in southern Babylonia (Von Rosen, 1990) and was used as glyptic from then until now in the manufacture of jewels, amulets, seals and inlays. To the ancient Egyptians, it was considered a gem representing the skies or heaven, thus was thought to denote light, truth and wisdom. It was thus often shaped into eye-shaped gems and was worn by judges in ancient Egypt. A lapis amulet graced the brow of Ra. Lapis is noted in Revelations in Christian mythology as a stone in the Breastplate of Aaron. In China, lapis was worn during the Manchu dynasty for services in the Temple of Heaven. The Romans and Greeks used it as a cure for fever and melancholy. It also glazed the bricks that formed the spectacularly blue “Gate of Kings” or Ishtar entryway to the ancient city of Babylon (≈1800 BC).

The Sar-e-Sang region has supplied much of the gem quality lapis to the world. One of the first European explorers to the region (Wood, 1841) described mining methods in use at that time. Camel-thorn and tamarisk twigs were collected from the valley below and carried up the steep path to the mine. When sufficient fuel had been collected, it was piled against the rock face and a fire was lit. When the rock was hot, cold water, which also had to be carried up the steep 350 m ascent from the valley floor, was thrown onto it. The rock cracked and split, enabling further work to be done with the primitive tools available (pick, hammer and chisel) in order to extract the lapis lazuli from its marble host rock.

[4] Gedrite is a silicate mineral of the amphibole group with formula: (Mg;Fe2+)2[(Mg;Fe2+)3Al2](Si6Al2)O22(OH)2

Seawater chemistry (2 of 2): Precambrian evolution of brine proportions

John Warren - Wednesday, August 26, 2015

We saw in the previous Salty Matters article (part 1 of 2) that ionic proportions of major ions in seawater and oceanic salinity have changed through the Phanerozoic and so influenced the make-up of bittern precipitates once the lower salinity salts (carbonates, gypsum and halite) had precipitated. In the Phanerozoic, seawater was dominantly a Na-K-Mg-Ca-Cl (Ca-rich) brine that changed periodically to a Na-K-Mg-Cl-SO4 (SO4-rich) type, as in the modern ocean. This oscillation across 600 million years forces  number of questions, for example, do similar oscillations in ocean chemistry extend back across the Precambrian? How consistent is the chemistry of the world’s oceans since the early Archean? Does the evaporite evidence in Precambrian sediments support a notion of a primordial reducing atmosphere and/or higher levels of bicarbonate in an early Archean ocean?

Some authors postulate that there have been no significant changes in the major ion proportions in seawater and hence the evaporation mineral series for the past 4 Ga (Morse and Mackenzie, 1998). Others assert that the Archean was dominantly a time of little or no atmospheric oxygen and that ocean waters were reducing anoxic fluids and so sulphate levels were low and sulphide levels high in evaporative marine waters (Krupp et al., 1994). Yet others propose that the bicarbonate to calcium ratio was so high in Archean and Palaeoproterozoic seawater compared to today that all the calcium was used up in widespread abiotic marine aragonite and Mg-calcite precipitates (Sumner and Grotzinger, 2000). In this case trona or nahcolite are likely marine evaporites in the early Archean bitterns (see Figure 1 in part 1). Still others have theorised cyclic changes in oceanic chemistry occurred across much of the Precambrian were similar to those of the Phanerozoic. Such changes were perhaps related to changes in styles and rates of sea floor spreading-hydrothermal circulation in midoceanic ridges (Channer et al., 1997) and the development of tonalitic continents (Knauth, 1998). 

Given that the world's oldest known halites occur in the Bitter Springs Formation in the Amadeus Basin of Australia and that they were deposited some 840 Ma, we can only extend a halite chevron inclusion-based study of ocean chemistry back to that time. These brines were sulphate-depleted, while recrystallised halite from the uppermost Neoproterozoic Salt Range Formation (ca. 545 Ma) in Pakistan, contains solitary inclusions indicating SO4-rich brines (Kovalevych et al., 2006). This supports a similar late Neoproterozoic ocean chemistry to today, as do proportions derived from primary fluid inclusions from the Neoproterozoic Ara Formation of Oman (ca. 545 Ma). It seems that  SO4-rich seawater existed during latest Neoproterozoic time. In contrast while recrystallised halite from the somehat older Bitter Springs Formation contains brine inclusions that are entirely Ca-rich, implying ambient basin brines and the mother seawater were Ca-rich some 830-840 Mas. These combined data, supported by the timing of aragonite and calcite seas, as preserved in various marine carbonates, suggest that during the Neoproterozoic, significant oscillations of the chemical composition of marine brines, and seawater occurred over the last 250 million years of the NeoProterozoic, and that the end-members were similar to those of the Phanerozoic oceans. It seems that Ca-rich seawater dominated for a substantial period of Late Precambrian time (more than 200 Ma) from 850 Ma, until some 650 Ma, this was replaced by SO4-rich seawater, returning to Ca-rich seawater at 530 Ma. 

The detail for much of the remaineder of the Precambrian back to 4 Ga is far less precise than when modelling inclusion chemistries based on actual halites. The oldest documented chevron halite is 850Ma and the oldest bedded anhydrite is 1.2Ga, beyond that, only evaporite pseudomorphs are available to study. So, beyond the 850 Ma record established by halite inclusions in the Bitter Springs Fm., can other Precambrian evaporites especially the calcium sulphates with a record that extends back patchily to the Mesoproterozoic, give indirect clues as to a chemical scenario for the world’s paleo-oceans and brine?

 

Pseudomorphs, especially of halite hoppers, occur in marine rocks as old as Archean, but are far more common, as are the actual salts, in Proterozoic strata (Figure 1; Warren, 2016). Halite or its pseudomorphs characterise areas of widespread marine chemical sedimentation from the Archean to the present. CaSO4 pseudomorph distribution is more enigmatic. In the 1980s and 1990s, the oldest documented CaSO4 pseudomorphs were thought to cm-sized growth-aligned barytes and cherts in 3.45 Ga metasediments in the Pilbara/North Poleregion of Western Australia. They were interpreted as replacing primary bottom-nucleated gypsum (Figure 2; Barley et al., 1979; Lowe, 1983; Buick and Dunlop, 1990). These barytes and cherts occur in volcaniclastics in association with what are possibly the world’s oldest stromatolites (Hofmann et al., 1999; Allwood et al., 2007). Similar growth-aligned baryte crystals, which initially were also interpreted as likely primary gypsum pseudomorphs, occur in the Nondweni greenstones in South Africa, some 3.4 Ga (Wilson and Versfeld, 1994).

 

Sequences in both regions are now completely silicified or barytised. At the time they were first documented, the recognition of what were considered shallow-water Early Archean gypsum pseudomorphs at North Pole, Pilbara Craton, caused a re-evaluation of models of a totally reducing Archean atmosphere (Dimroth and Kimberley, 1975; Clemmey and Badham, 1982). The presence of free sulphate in surface brines of the Archean world was thought to imply an at least locally oxygenated hydrosphere. Gypsum precipitating in Archean ocean waters also meant calcium levels in the ocean waters were in excess of bicarbonate, as is in the modern oceans. The presence of free-standing gypsum on the seafloor is incompatible with any model of the Early Archean ocean as a “soda lake.”

However, in both the Pilbara and the South African sequences there are no actual calcium sulphate evaporites preserved, only growth-aligned crystal textures, now preserved as baryte or chert. Textures in baryte ore from Frasnian sediments in Chaudfontaine, Belgium, are near identical to those observed at North Pole, Australia. The Belgian barytes are primary shallow subsea-bottom precipitates with no precursor mineral phase (Figure 2 inset; Dejonghe, 1990). Some workers in the Pilbara feel that the growth-aligned Archean baryte in this region is also a primary seafloor precipitate, formed in the vicinity of hydrothermal vents (Vearncombe et al., 1995; Nijman et al., 1999; Runnegar et al., 2001). As such, it is not secondary after gypsum. A similar hydrothermal discharge model has been developed for aligned barytes in the Barberton Greenstone belt (de Ronde et al., 1994, 1996). 

Based on this more recent analysis, levels of Archean sulphate in the world ocean were probably less than a few percent of the current levels and probably remained so until the evolution of a widespread oxygen-producing biota into the Proterozoic (Figures 3, 4; Habicht and Canfield, 1996; Kah et al., 2004). Barium sulphate is highly insoluble in modern oxygenated seawater. To carry large volumes of barium or sulphur (as sulphide) in seawater solution to the precipitation site required anoxic conditions. If the aligned baryte crystals are primary, their formation still requires sulphate to be locally present on the seafloor, at least in the vicinity of the depositional site. A possible source for local sulphate production in the shallow waters that characterised the North Pole site was shortwave ultraviolet photoxidation of volcanic SO2, indicating an inorganic association (Runnegar et al., 2001). Within barytes in the same 3.47-Ga-old barytes there are microscopic sulphides. These sulphide inclusions show a d34S of 11.6‰, possibly indicating microbial sulphate reduction with H2 as electron donor in what was an anoxic seafloor (Canfield et al. 2004; Shen et al., 2009).

According to Nijman et al. (1999) the occurrence of the North Pole baryte in sedimentary mounds atop growth faults meant sulphate was locally derived via boiling of escaping hydrothermal vent waters enriched in Ba, Si and sulphide. As these hydrothermal waters vented beneath marine water columns perhaps 50 metres deep, they boiled or violently degassed. Consequent mixing with normally stratified seawater, caused instantaneous oxidization of sulphide into sulphate that then, on cooling, combined with the Ba to precipitate as growth-aligned baryte crystals on the seafloor. Conflicting notions (replaced gypsum versus primary baryte) mean that at this stage of our understanding, the bedded baryte evidence cannot be reliably used to support an evaporite paragenesis of gypsum and so infer an Archean ocean with ionic proportions similar to those of today.

Archean and Proterozoic distributions of gypsum have been further complicated by the misidentification of primary aragonite splays and pinolitic siderite marbles as gypsum replacements (Warren 2016; Chapter 15). When these misidentifications are removed from the record it is obvious that calcium sulphate precipitating directly from Archean seawater to form widespread beds did not occur, and that precipitation of aragonite as thick crusts on the sea floor was significantly more abundant than during any subsequent time in earth h istory. In contrast to gypsum, halite pseudomorphs are found throughout the Precambrian (Figure 1;e.g. Boulter and Glover, 1986). 

Grotzinger and Kasting (1993) argue that high levels of atmospheric CO2 meant HCO3/Ca ratios were much higher in the Archean and the Palaeoproterozoic oceans than today. All the calcium in seawater was deposited as marine cementstones and other alkaline earth precipitates well before bicarbonate was depleted and there was no Ca left over to precipitate as gypsum. The early Archean ocean was perhaps a Na–Cl–HCO3 sea, and not the Na–Cl ocean of today (Kempe and Degens, 1985; Maisonneuve, 1982). This early Archean hydrosphere had a chemistry similar to that found in modern soda lakes like Lake Magadi and Lake Natron (pathway I brines) and hence the term “soda-lake oceans” (see Figure 1 in part 1) This rather different marine brine chemistry would have precipitated halite and trona/nahcolite, not halite/gypsum. It probably meant that if gypsum did ever precipitate from Archean seawater it did so only in minor amounts well after the onset of halite precipitation. Excessive sodium in the ocean may help explain the ubiquity of stratiform albitites in much of the Archean. They would have formed throughout the marine realm as early diagenetic replacements of labile volcaniclastics/zeolites in volcanogenic/greenstone terranes).

A case for nahcolite (NaHCO3) as a primary evaporite, along with halite, in the 3.42 Ga rocks of the Barberton greenstone belt was documented by Lowe and Fisher-Worrell (1999). Sugitani et al. (2003) reported silicified nahcolite (the high CO2 form of sodium carbonate salts) in ≈3.2 Ga rocks in the northern part of the Eastern Pilbara block, Western Australia. Coarse, upward-radiating, silicified evaporite crystals in the ca. 3.47–3.46 Ga Strelley Pool Chert (Lowe, 1983) show the same habit, geometry, and environmental setting as nahcolite in the Barberton belt and also probably represent silicified NaHCO3 precipitates (Lowe and Tice, 2004).


Marine nahcolite in the 3.5-3.2 Ga sedimentary record is thought to be evidence of surface temperatures around 70±15°C (Figures 3b, c, 4; Lowe and Tice, 2004). Contemporary early Archean nahcolite (NaHCO3) as a primary evaporitic mineral in a very aggressive weathering regime, in the absence of land vegetation, is best explained by a mixed CH4 and CO2 atmospheric greenhouse. CH4/CO2 ratios were <<1 and pCO2 was at least 100-1000 times the present value, perhaps as high as several bars (Kaufman and Xiao, 2003). The formation of large areas of continental crust at 3.2-3.0 Ga, including the Kaapvaal and Pilbara cratons, resulted in the gradual depletion of atmospheric CO2 through weathering and a lack of marine nahcolite since the early Archean. By 2.9-2.7 Ga, declining pCO2 was associated with climatic cooling and siderite-free soils. 

Transitory CH4/CO2 ratios of ~1 may have resulted in the sporadic formation of organic haze from atmospheric CH4, and are reflected in one or more isotopic excursions involving global deposition of abnormally 13C-depleted organic carbon in sediments of this age. Surface temperatures of <60°C after 2.9 Ga may have allowed an increase in the distribution and productivity of oxygenic photosynthetic microbes (and a decrease in sulphur dependent thermophiles). Eventual lowering of newly formed continental blocks by erosion, reduced loss of atmospheric CO2 due to weathering, and continued long-term tectonic recycling of CO2 resulted in rising pCO2 and decreasing CH4/CO2 ratios in the later Archean and eventual re-establishment of a mainly CO2 greenhouse. Similar events may have been repeated in the latest Archean and earliest Proterozoic, but gradually rising production of O2 effectively kept CH4/CO2 ratios to <<1.

 

By 2.2-2.0 Ga and perhaps as early as 2.5 Ga, reliable examples of pseudomorphs after primary marine-sourced calcium sulphate first appear in the rock record, but aside from the Karelian beds associated with the Lomagundi Event (LE), widespread stratiform sulphate beds of anhydrite do not appear until 1.2 Ga (Figure 5a). Undeniable CaSO4 nodular and lenticular pseudomorphs are widespread in latest NeoArchean of South Africa and Palaeoproterozoic to Mesoproterozoic sediments of the McArthur Basin, Northern Territory, Australia, and in rocks of Great Slave Lake in northern Canada. For example, in the Malapunyah Formation (1.65 Ga) of the Northern Territory, Australia, the outer portions of numerous decimetre to metre-diameter silicified anhydrite nodules still retain outlines of felted anhydrite laths (pers. obs). The oldest reliable sulphate pseudomorphs after anhydrite and gypsum in Australia come from Palaeoproterozoic cherts in the 2.0-2.2 Ga Bartle Member of the Killara Formation, western Australia (Pirajno and Grey, 2002). These cherts locally retain small amounts of anhydrite (verified by XRD, as well as appearing as highly birefringent flecks in thin sections). Other widespread but younger sulphate pseudomorphs occur in the 1.2 Ga Amundsen Basin in the Canadian Arctic Archipelago. Actual CaSO4 beds outcrop in the 1.2 Ga Society Cliff Formation in Baffin and Bylot Islands of the Canadian Archipelago (Kah et al., 2001, 2004). Sulphate evaporite pseudomophs and nodules in all these Neoproterozoic basins are hosted in sedimentary layers up to tens of metres thick and with lateral extents measured in hundreds of square kilometres. All were laid down in shallow marine, coastal, and alluvial environments under an increasingly oxygenated Meso- to Neoproterozoic atmosphere (Jackson et al., 1987; Walker et al., 1977). After passing from the Archean, by the Mesoproterozoic the hydrosphere contained free sulphate and Ca/HCO3 ratios were lower, leading to a decrease in molar-tooth, herringbone and other carbonate textures indicative of widespread inorganic calcium carbonate saturation in shallow oceanic waters (Figure 6). However, oceanic mother brines for these now-widespread calcium-sulphate evaporites were largely H2S rich with only moderate levels of oxygen in the atmosphere until some 800 Ma (Figure 3a).

The work of Kah et al. (2004) shows that prior to 2.2 Ga, when oxygen began to accumulate in the Earth’s atmosphere, sulphate concentrations in the world’s oceans were low, <1 mM and possibly <200 μM (Figure 5). By 0.8 Ga, oxygen and thus sulphate levels had risen significantly. Sulphate levels were between 1.5 and 4.5 mM, or 5–15% of modern values, for more than a billion years after initial oxygenation of the Earth’s biosphere some 2.2-2.4 Ga and mid -ocean depth waters were anoxic for most of that time (Brocks et al., 2005). Marine sulphate concentrations probably remained low, no more than 35% of modern values, for nearly the entire Proterozoic. A significant rise in biospheric oxygen, and thus oceanic sulphate, may not have occurred until the latest Neoproterozoic (0.54 Ga), just before the Cambrian explosion, when sulphate levels may have reached 20.5 mM, or 75% of present day levels. This is a time when thick sulphate platforms first characterised the salt basins of Oman, prior to that most actual calcium sulphate is in the form of nodules or relatively thin beds.

In a refinement of the sulphate model, Bekker and Holland (2012) note that free sulphate bottom-nucleated sulphate evaporites and not just pseudomorphs were present during the Lomagundi Event (2.22 to 2.06 Ga), and then became relatively scarce once more until some 1.2 Ga. For example, there is a 200 m thick stratigraphic interval of sulphate evaporites of Lomagundi-age, preserved in a shallow-water open-marine siliciclastic and carbonate succession (Lower Jatuli informal group) of Karelia, Russia (Morozov et al., 2010). The Lomagundi Event defines the most extreme and longest lasting isotope excursion of carbon in the world’s marine carbonate record. Bedded gypsum pseudomorphs in the Malmani Group some 2.5 Ga (Gandin and Wright, 2007; Eriksson and Warren, 1983) implies that elevated oceanic sulphate levels that typify the Lomagundi Event may have extended a little further back in time, at least locally (Figure 5).

At the same time as the Lomagundi event, the average ferric iron to total iron (expressed as Fe2O3/Fe|Fe2O3|) ratio of shales increased dramatically. At the end of the Lomagundi Event (LE), the first economic sedimentary phosphorites were deposited, and the carbon isotope values of marine carbonates returned to ≈0.0‰VPDB (Figure 2.50). Thereafter marine sulphate evaporites and phosphorites again became scarce, while the average Fe2O3/Fe|Fe2O3| ratio of shales decreased to values intermediate between those of the Archean and Lomagundi-age shales.

In support of this notion of an “oxygen overshoot,” sulphur isotope work by Reuschel et al. (2012) on the 2.1 Ga dolomitic Tulomozero Fm, which entrains abundant CaSO4 pseudomorphs, concluded that there was a minimum level of 2.5 mM sulphate in the world ocean at that time (Figure 5).

Bekker and Holland (2012) argue the short appearance of sulphate evaporites in Logamundi and the other associated events can be regarded as a ca. 200 Ma “glitch” in the gradual oxidation of the atmosphere–ocean system. It was driven by a positive feedback between the rise in atmospheric O2, the oxidation of pyrite in rocks undergoing weathering, a decrease in the pH of soil and ground water, and an increase in the phosphate flux to the oceans. This sequence led to a major increase in the rate of organic matter burial, a rise in atmospheric oxygen, a large increase in the 13C value for marine carbonates, the deposition of marine evaporites containing gypsum and anhydrite, and the formation of the first commercially important phosphorites. The end of the LE was probably brought about by the weathering of sediments deposited during the LE.

In yet another proposal of hydrosphere-atmosphere evolution, Huston and Logan (2004) argue that the presence of relatively abundant bedded sulphate deposits before 3.2 Ga (as the contentious Archean barytes and chert mentioned earlier) and after 1.8 Ga (as CaSO4 salts), and the peak in banded iron formation abundance between 3.2 and 1.8 Ga, and the aqueous geochemistry of sulphur and iron, when taken together suggest that the redox state and the abundances of sulphur and iron in the hydrosphere varied widely during the Archean and Proterozoic. They propose a layered hydrosphere prior to 3.2 Ga in which sulphate was enriched in an upper oceanic layer, whereas the underlying layer was reduced and sulphur-poor. The sulphate was produced by atmospheric photolytic reactions with volcanic gases in a reducing atmosphere. Mixing of the upper and lower water masses allowed the banded barytes to form prior to 3.2 Ga and created an ocean chemistry where nahcolite was a marine evaporite. Between 3.2 and 2.4 Ga, decreasing volcanogenesis and sulphate reduction removed sulphate from the upper layer, producing broadly uniform, reduced, sulphur-poor and iron-rich oceans.

Whatever the origin of the early Archean baryte and chert, around 2.2 - 2.4 Ga, as a result of increasing atmospheric oxygenation, the flux of sulphate into the hydrosphere by oxidative weathering was greatly enhanced, producing layered oceans, with sulphate-enriched, iron-poor surface waters and reduced, sulphur-poor and iron-rich bottom waters. Gypsum evaporites were increasingly likely as marine precipitates. The rate at which this process proceeded varied between basins depending on the size and local environment of the basin. By 1.8 Ga, the hydrosphere was relatively sulphate-rich and iron-poor throughout. Gypsum was now a widespread marine evaporite. Variations in sulphur and iron abundances suggest that the redox state of the oceans was buffered by iron before 2.4 Ga and by sulphur after 1.6 to 1.8 Ga (Figure 1).

Gypsum in combination with halite was the marine evaporite association from then until now. Seawater was predominantly a Na-Cl±SO4 ocean. Neoproterozoic stratiform sulphates along with widespread halokinetic halite, occur in the Bitter Springs Formation of the Amadeus basin, central Australia (0.8 Ga), its equivalents in the Officer Basin, the Callana beds of the Flinders Ranges and the younger Infracambrian salt basins of the Arabian (Persian) Gulf (≈0.545 Ga; Wells, 1980; Cooper, 1991; Mattes and Conway-Morris, 1990; Edgell, 1991).


The transition to calcium sulphate textures in evaporite pseudomorphs mirrors a marked change in the style of marine carbonates that began around 2.2 to 2.3 Ga when herringbone calcite and precipitated carbonate beds become much less common and the precipitation mode shifted from the seafloor to the water column (Figure 6; Sumner and Grotzinger, 1996, 2000). The boundary also corresponds to the “rusting” of the oceans when oxygen levels became high enough to precipitate widespread banded iron deposits on the seafloor. Microdigitate stromatolites cross this boundary with little effect, suggesting the marked decrease in dissolved iron exerted little influence on them.

The relative scarcity of actual Pre-Phanerozoic salts, not pseudomorphs, especially in the Archean has been used by some to argue that conditions were less favourable for widespread evaporite deposition in the early Precambrian (Cloud, 1972). Others, myself included, feel that the relative scarcity of preserved evaporites in older sequences reflects the greater likelihood of fluid flushing, evaporite dissolution and metasomatism in progressively older rocks. It is likely that oceanic calcium-sulphate evaporites were less common in the Archean, and that sodium carbonates mixed with halite were dominant evaporite salts in the seawater-fed saline giants in appropriate tectonic seepage depressions of the Early Archean. But widespread evaporite deposition from sodium-dominated brines did occur throughout the Archean in large drawdown basins isolated from a surface connection with the ocean. A paucity of preserved bedded evaporite salts in the Precambrian reflects an increased probability of partial or complete evaporite dissolution, remobilization and metasomatism with increasing geological age (see meta-evaporite).

In what is an inclusion study of oldest actual halite, Spear et al., (2014) characterised marine brine chemistry using brine inclusions in the 830 Ma salt of the Browne Formation, Officer Basin, Australia (equiv. to Bitter Springs Fm.). It seems that concentrations of the major ions in these inclusions, except K+ and possibly SO42−, fall within the known range of Phanerozoic seawaters. This ananlysis suggests that mid-Neoproterozoic marine sulphate concentrations were lower (≈90%) than modern values. By the terminal Neoproterozoic, fluid inclusions in halite and evaporite mineralogy from the Khewra Salt of Pakistan and the Ara salt in Oman indicate seawater sulphate levels had risen significantly, to 50%-80% of modern concentrations, which parallels increases in atmospheric and oceanic oxygen.

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Seawater chemistry (1 of 2): Potash bitterns and Phanerozoic marine brine evolution

John Warren - Tuesday, August 11, 2015

The significance of evaporites as indicators of the chemical evolution of seawater across time and in relation to potash bitterns is considered in the next two Salty Matters articles. This article focuses on Phanerozoic seawater chemistry, where actual salts are widespread and the proportions of potash bittern salts are a useful pointer to the chemical makeup of the mother brine. Throughout both articles, the term “lower salinity” refers to marine brines with salinities between one and ten times that of ambient seawater. The second article considers seawater chemistry based on Precambrian evaporites, where much of the evidence of mother brine composition comes from salt pseudomorphs, rather than remnants of actual salts. In the second article we shall see that atmospheric conditions in the Early Precambrian were reducing and hotter than today, so that seawater was more saline, warmer, anoxic, with higher levels of calcium and bicarbonate compared to Phanerozoic seawater. Gypsum (CaSO4.2H2O), which requires free sulphate, was a rare precipitate during concentration of Archean seawater. Changing atmospheric proportions of CO2, CH4 and O2 meant sodium carbonate salts were significant lower-salinity early Archean marine-brine precipitates. Yet today, sodium carbonate salts, such as trona (NaHCO3.Na2CO3), nahcolite (NaHCO3) and shortite (2CaCO3.Na2CO3) cannot precipitate from a brine with the ionic proportions of modern seawater. The presence of sodium carbonate salts in any evaporite succession across the Phanerozoic is a reliable indicator of a nonmarine mother brine (Figure 1).


A Phanerozoic dichotomy: evolving marine potash bitterns

Consistently across the last 550 million years, halite and gypsum (mostly converted to anhydrite in the subsurface) are the dominant lower-salinity marine salts. But potash-bittern evaporite associations plotted across the same time framework define two end-members (Figure 1):

1) Sulphate-enriched potash deposits, with ores typically composed of halite (NaCl) with carnallite (MgCl2.KCl.6H2O) and lesser sylvite (KCl), along with varying combinations of MgSO4 salts, such as polyhalite (2CaSO4.MgSO4.K2SO4.H2O), kieserite (MgSO4.H2O) kainite (4MgSO4.4KCl.11H2O) and langbeinite (2MgSO4. K2SO4); and

2) Sulphate-depleted potash deposits are composed of halite with sylvite and carnallite, and entirely free or very poor in the magnesium-sulphate salts. The sulphate-depleted association typifies more than 65% of the world’s exploited Phanerozoic potash deposits. Sylvite ores with this association have properties that are easier to process cheaply (Warren 2016; see also blog 4 of 4 in the Salty Matters Danakhil articles).

The sulphate-enriched group of ancient potash salts contains a bittern mineral suite predicted by the evaporation and backreaction of seawater with proportions similar to modern marine brine. In contrast, the sulphate-depleted group of bittern salts must have precipitated from Na-Ca-Mg-K-Cl brines with ionic proportions quite different from that of concentrated modern seawater. The separation between the two bittern associations is defined by brine evolution across the gypsum divide. That is, once gypsum (CaSO4.2H2O) and halite (NaCl) have precipitated in the lower salinity spectrum, are the remaining brines enriched in sulphate or calcium (Figure 1)? The greater suitability for potash utilisation of the sulphate depleted bitterns makes understanding and hence predicting occurrences of the sulphate-depleted association in time and space a useful first-order potash exploration tool.

Why the dichotomy?

In the older literature dealing with Phanerozoic salt chemistry, MgSO4-depleted potash evaporites were often explained as diagenetically-modified marine evaporite brines, thought to result from backreactions during burial diagenesis of normal marine waters (Borchert, 1977; Dean, 1978; Wilson and Long, 1993). If so, then the mother seawater source across the Phanerozoic had ionic proportions like those of today, but diagenetically altered via; a) dolomitisation, b) sulphate-reducing bacterial action, c) mixing of brines with calcium bicarbonate-rich river water, or d) rock-fluid interaction during deep burial diagenesis. As another option, Hardie (1990) suggested MgSO4-depleted potash bitterns formed by the evaporative concentration of sulphate-depleted nonmarine inflow waters seeping into an evaporite basin via springs and faults. Such springs were sourced either from CaCl2-rich hot hydrothermal brines or via cooling of deep basinal brines. Such fault-fed deeply-circulating CaCl2 brines source the various springs feeding the Dead Sea, the Qaidam Basin, the Salton Sea and the Danakil Depression. In all these cases, the elevated salinities of inflow waters are related, at least in part, to the dissolution of buried evaporites. Upwelling of brines in these regions is driven either by thermally-induced density instabilities, related to magma emplacement, or by the creation of tectonically-induced topographic gradients that force deeply-circulated basinal brines to the surface. Ayora et al. (1994) demonstrated that such a deeply-circulating continental Ca–Cl brine system operated during deposition of sylvite and carnallite in the upper Eocene basin of Navarra, southern Pyrenees, Spain.

Today, a more widely accepted explanation for SO4-enriched versus SO4-depleted Phanerozoic potash bitterns, is that seawater chemistry has evolved across deep time. Background chemistry of the marine potash dichotomy is simple and can be related to brine evolution models published by Hardie more than 30 years ago (Hardie, 1984). He found that the constituent chemical proportions in the early stages of concentration of any marine brine largely controls the chemical makeup of the subsequent bittern stages. These ionic proportions control how a brine passes through the lower salinity CaCO3 and gypsum divides (Figure 1). That is, a marine brine’s bittern make-up is determined by the ionic proportions in the ambient seawater source. It determines the carbonate mineralogy during the precipitation of relatively insoluble evaporitic carbonates (aragonite, high-magnesium calcite, or low-magnesium calcite) which in turn controls its constituent chemistry as it attains gypsum saturation. These two stages are called the CaCO3 and gypsum divides. Hence, the chemical passage of a bittern is controlled by the ionic proportions in the original ambient seawater. The CaCO3 divide kicks in when a concentrating seawater brine attains a salinity around twice that of normal seawater (60‰). The gypsum divide occurs when brine concentrations are around 4-5 times that of normal seawater (140-160‰). Normal seawater has a salinity around 35-35‰ and the various potash bittern salts precipitate when concentrations are around 40-60 times that of the original seawater (Figure 2 – lower part).

As seawater concentrates and calcium carbonate mineral(s) begin to precipitate at the CaCO3 divided then, depending on the relative proportions of Ca and HCO3 in the mother seawater, either Ca is used up, or the HCO3 is used up. If the Ca is used up first, an alkaline brine (pH>10) forms, with residual CO3, along with Na, K, SO4 and Cl, but no remaining Ca (Figure 1). With ongoing concentration this brine chemistry will then form sodium bicarbonate minerals, it cannot form gypsum as all the Ca is already used up. Such an ionic proportion chemistry likely defined oceanic waters in the early Archaean but is not relevant to seawater evolution in the Proterozoic and Phanerozoic, as evidenced by widespread gypsum (anhydrite) or pseudomorphs in numerous post-Archean marine-evaporite basins. At higher concentrations, early Archean marine brines would have produced halite and sylvite bittern suites, but with no gypsum or anhydrite (Figure 1).

If, instead, HCO3 is used up during initial evaporitic carbonate precipitation, as is the case for all Phanerozoic seawaters, the concentrating brine becomes enriched in Ca and Mg, and a neutral brine, depleted in carbonate, is formed. Then the ambient Mg/Ca ratio in a concentrating Phanerozoic seawater will control whether the first-formed carbonate at the CaCO3 divide is aragonite (Mg/Ca>5) or high-magnesium calcite (2>Mg/Ca>5), or low-Mg calcite(Mg/Ca>2). The latter Mg/Ca ratio is so low it is only relevant to concentrating Cretaceous seawaters. Elevated Mg/Ca ratios favouring the precipitation of aragonite over high Mg-calcite typify modern marine seawater brines, which have Mg/Ca ratios that are always >5 (Figure 2). At lower salinities, modern marine brines are Na-Cl waters that with further concentration and removal of Na as halite evolve in Mg-SO4-Cl bitterns (Figure 2)

 

The next chemical divide reached by concentrating marine Phanerozoic brines (always depleted in HCO3 at the carbonate divide) occurs when gypsum precipitates at around 4-5 times the concentration of the original seawater (Gypsum divide in Figure 1). As gypsum continues to precipitate, either the Ca in the brine is used up, or the SO4 in the brine is used up. If the Ca is depleted, a calcium-free brine rich in Na, K, Mg, Cl, and SO4 will be the final product and the diagnostic bittern minerals will include magnesium sulphate minerals. This is the pathway followed by modern seawater bitterns. If, however, the sulphate is used up via gypsum precipitation, the final brine will be rich in Na, K, Mg, Ca, and Cl. Such a sulphate-depleted brine precipitates diagnostic potassium and magnesium chloride minerals such as sylvite and carnallite. If calcium-chloride levels are very high, then diagnostic (but uncommon) minerals such as tachyhydrite (CaCl2.2MgCl2.12H2O), and antarcticite (CaCl2.6H2O) can precipitate from this brine. But both these assemblages contain no sulphate bittern minerals, making potash processing relatively straightforward (Warren, 2016). In Phanerozoic marine salt assemblages, tachyhydrite, which is highly hygroscopic, is present in moderate quantities only in Cretaceous (Aptian) marine sylvite-carnallite associations in the circum-Atlantic potash basins and the Cretaceous (Albian) Maha Sarakham salts of Thailand, along with its equivalents in Laos and western China. The CaCl2-entraining bittern mineral assemblages of these deposits imply ionic proportions of Cretaceous seawater differ from those of today.

Inclusion evidence

Based on a study of brine inclusion chemistry preserved in halite chevrons, from the Early Cretaceous (Aptian, 121.0–112.2 Ma) of the Sergipe Basin, Brazil, the Congo Basin, Republic of the Congo, and the Early to Late Cretaceous (Albian to Cenomanian, 112.2–93.5 Ma) of the Khorat Plateau, Laos and Thailand, Timofeeff et al. (2006) defined a very different chemical makeup for Cretaceous seawater, compared to that of today. Brine proportions in the fluid inclusions in these halites indicate that Cretaceous seawaters were enriched several fold in Ca, depleted in Na and Mg, and had lower Na/Cl, Mg/Ca, and Mg/K ratios compared to modern seawater (Table 1). 


Elevated Ca concentrations, with Ca>SO4 at the gypsum divide, allowed Cretaceous seawater to evolve into Mg–Ca–Na–K–Cl brines lacking measurable sulphate. Aptian seawater was extreme in its Ca enrichment, more than three times higher than present day seawater, with a Mg/Ca ratio of 1.1–1.3. Younger, Albian-Cenomanian seawater had lower Ca concentrations, and a higher Mg/Ca ratio of 1.2–1.7. Cretaceous (Aptian) seawater has the lowest Mg/Ca ratios so far documented in any Phanerozoic seawater from fluid inclusions in halite, and lies well within the range chemically favourable for precipitation of low-Mg calcite ooids and cements in the marine realm.


Likewise, a detailed analysis of the ionic make-up of Silurian seawater using micro-inclusion analysis of more than 100 samples of chevron halite from various Silurian deposits around the world was published by Brennan and Lowenstein (2002), clearly supports the notion that ionic proportions in the world’s Silurian oceans were different from those of today (Figure 3). Samples were from three formations in the Late Silurian Michigan Basin, the A-1, A-2, and B Evaporites of the Salina Group, and the Early Silurian in the Canning Basin (Australia) in the Mallowa Salt of the Carribuddy Group. The Silurian ocean had lower concentrations of Mg, Na, and SO4, and much higher concentrations of Ca relative to the ocean’s present-day composition (Table 1). Furthermore, Silurian seawater had Ca in excess of SO4. Bittern stage evaporation of Silurian seawater produced KCl-type potash minerals that lack the MgSO4-type late stage salts formed during the evaporation of present-day seawater and allowed sylvite as a primary precipitate. In a similar fashion, work by Kovalevych et al. (1998) on inclusions in primary-bedded halite from many evaporite formations of Northern Pangaea, and subsequent work using micro-analyses of fluid inclusions in numerous chevron halites (Lowenstein et al., 2001, 2003), shows that during the Phanerozoic the chemical composition of marine brines has oscillated between Na-K-Mg-Ca-Cl and Na-K-Mg-Cl-SO4 types. The former does not precipitate MgSO4 salts when concentrated, the latter does (Figure 3). A recent paper by Holt et al. (2014), focusing on chevron halite inclusions from various Carboniferous evaporite basins, further refined the transition from the Palaeozoic CaCl2 high Mg-calcite sea into a MgSO4-enriched aragonite ocean of the Permo-Carboniferous, so showing CaCl2 oceanic chemistry (and sylvite-dominant bitterns) extend somewhat further across the Palaeozoic than previously thought (Figure 4).

 

More recent work has shown varying sulphate levels in the Phanerozoic ocean rather than Mg/Ca variations are perhaps more significant in controlling aragonite versus calcite at the CaCO3 divide and the associated evolution of MgSO4-enriched versus MgSO4-depleted bittern suites in ancient evaporitic seaways than previously thought. Bots et al. (2011) found experimentally that an increase in dissolved SO4 decreases the Mg/Ca ratio at which calcite is destabilized and aragonite becomes the dominant CaCO3 polymorph in an ancient seaway (Figure 5). This suggests that the Mg/Ca and SO4 thresholds for the onset of ancient calcite seas are significantly lower than previous estimates and that Mg/Ca levels and SO4 levels in ancient seas are mutually dependent. Rather than variations in Mg/Ca ratio in seawater being the prime driver of the aragonite versus calcite ocean chemistries across the Phanerozoic, they conclude sulphate levels are an equally important control.


Mechanisms

There is now convincing inclusion-based evidence that the chemistry of seawater has varied across the Phanerozoic from sulphate-depleted to sulphate-enriched, what is not so well understood are the various worldscale processes driving the change (Figure 4). Spencer and Hardie (1990) and Hardie (1996) argued that the level of Mg in the Phanerozoic oceans has been relatively constant across time, but changes in the rate of seafloor spreading have changed the levels of Ca in seawater. This postulate is also supported in publications by Lowenstein et al. (2001, 2003). Timing of the increase of Ca in the world’s oceans was likely synchronous with a decrease in the SO4 ion concentration, which at times was as much as three times lower than the present.

Simple mixing models show that changes in the flux rate of mid-oceanic hydrothermal brines can generate significant changes in the Mg/Ca, Na/K and SO4/Cl ratios in seawater (Table 1). Changes of molal ratios in seawater have generated significant changes in the type and order of potash minerals at the bittern stage. For example, Spencer and Hardie’s (1990) model predicts that an increase of only 10% in the flux of mid-ocean ridge hydrothermal brine over today’s value would create a marine bittern that precipitates sylvite and calcium-chloride salts, as occurred in the Cretaceous instead of the Mg-sulphate minerals expected during bittern evaporation of modern seawater. Such Ca-Cl potash marine bitterns correspond to times of “calcite oceans” and contrast with the lower calcium, higher magnesium, higher sulphate “aragonite oceans” of the Permo-Triassic and the Neogene (Figure 3; Hardie, 1996; Demicco et al., 2005).

Ocean crust, through its interaction with hydrothermally circulated seawater, is a sink for Mg and a source of Ca, predominantly via the formation of smectite, chlorite, and saponite via alteration of pillow basalts, sheeted dykes, and gabbros (Müller et al., 2013). Additional removal of Mg and Ca occurs during the formation of vein and vesicle-filling carbonate and carbonate-cemented breccias in basalts via interaction with low-temperature hydrothermal fluids. Hence, changing rates of seafloor spreading and ridge length likely influenced ionic proportions in the Phanerozoic ocean and this in turn controlled marine bittern proportions.

According to Müller et al., 2013, hydrothermal ocean inputs are and the relevant ionic proportions in seawater are driven by supercontinent cycles and the associated gradual growth and destruction of mid-ocean ridges and their relatively cool flanks during long-term tectonic cycles, thus linking ocean chemistry to off-ridge low-temperature hydrothermal exchange. Early Jurassic aragonite seas were a consequence of supercontinent stability and a minimum in mid-ocean ridge length and global basalt alteration. The breakup of Pangea resulted in a gradual doubling in ridge length and a 50% increase in the ridge flank area, leading to an enhanced volume of basalt to be altered. The associated increase in the total global hydrothermal fluid flux by as much as 65%, peaking at 120 Ma, led to lowered seawater Mg/Ca ratios and marine hypercalcification from 140 to 35 Ma. A return to aragonite seas with preferential aragonite and high-Mg calcite precipitation was driven by pronounced continental dispersal, leading to progressive subduction of ridges and their flanks along the Pacific rim.

Holland et al. (1996), while agreeing that there are changes in ionic proportion of Phanerozoic seawater and that halite micro-inclusions preserve evidence of these changes, recalculated the effects of changing seafloor spreading rates on global seawater chemistry used by Hardie and others. They concluded changes in ionic proportions from such changes in seafloor spreading rate were modest. Instead, they pointed out that the composition of seawater can be seriously affected by secular changes in the proportion of platform carbonate dolomitised during evaporative concentration, without the need to invoke hydrothermally driven changes in seawater composition. In a later paper, Holland and Zimmermann (2000) suggest changes in the level of Mg in seawater were such that the molar Mg/Ca ratio of the more saline Palaeozoic global seawater (based on dolomite volume) was twice the present value of 5.

Using micro-inclusion studies of halites of varying ages, Zimmermann (2000a, b) has proposed that the evolving chemistry of the Phanerozoic ocean is more indicative of changing volumes of dolomite than it is of changes in the rates of seafloor spreading . Using halite inclusions, she showed that the level of Mg in seawater has increased from ≈38 mmol/kg H2O to 55 mmol/kg H2O in the past 40 million years (Figure 6). This increase is accompanied by an equimolar increase in the level of oceanic sulphate. Over the longer time frame of the Palaeozoic to the present the decrease in Mg/Ca ratio corresponds to a shift in the locus of major marine calcium carbonate deposition from Palaeozoic shelves to the deep oceans, a change tied to the evolution of the nannoplankton. Prior to the evolution of foraminifera and coccoliths, some 150 Ma, the amount of calcium carbonate accumulating in the open ocean was minimal. Since then, a progressively larger portion of calcium carbonate has been deposited on the floor of the deep ocean. Dolomitization of these deepwater carbonates has been minor.

 

In a study of boron isotopes in inclusions in chevron halite, Paris et al. (2010) mapped out the changes in marine boron isotope compositions over the past 40 million years (Figure 7). They propose that the correlation between δ11BSW and Mg/Ca reflects the influence of riverine fluxes on the Cenozoic evolution of oceanic chemical composition. Himalayan uplift is a major tectonic set of events that probably led to a 2.5 times increase of sediment delivery by rivers to the ocean over the past 40 m.y. They argue that chemical weathering fluxes and mechanical erosion fluxes are coupled so that the formation of the Himalaya favoured chemical weathering and hence CO2 consumption. The increased siliciclastic flux and associated weathering products led to a concomitant increase in the influx levels of Mg and Ca into the mid to late Tertiary oceans. However the levels of Ca in the world’s ocean are largely biologically limited (mostly by calcareous nannoplankton and plankton), so leading to an increase in the Mg/Ca ratio in the Neogene ocean.

 

a study of CaCO3 veins in ocean basement, utilising 10 cored and documented drilled sites, Rausch et al. (2013) found for the period from 165 - 30 Ma the Mg/Ca and the Sr/Ca ratios were relatively constant (1.22-2.03 mol/mol and 4.46-6.62 mmol/mol respectively (Figure 8). From 30 Ma to 2.3 Ma there was a steady increase in the Mg/Ca ratio by a factor of 3, mimicking the brine inclusion results in chevron halite. The authors suggest that variations in hydrothermal fluxes and riverine input are likely causes driving the seawater compositional changes. They go on to note that additional forcing may be involved in explaining the timing and magnitude of changes. A plausible scenario is intensified carbonate production due to increased alkalinity input to the oceans from silicate weathering, which in turn is a result of subduction-zone recycling of CO2 from pelagic carbonate formed after the Cretaceous slow-down in ocean crust production rate. However, world-scale factors driving the increase in Mg in the world’s oceans over the past 40 million years are still not clear and are even more nebulous the further back in time we look.

 

Changes in Phanerozoic ocean salinity

As well as changes in Mg/Ca and SO4, the salinity of the Phanerozoic oceans shows a fluctuating but overall general decrease from the earliest Cambrian to the Present (Figure 9; Hay et al. 2006). The greatest falls in salinity are related to major extractions of NaCl into a young ocean (extensional continent-continent proximity) or foreland (compressional continent-continent proximity) ocean basins (Chapter 5). Phanerozoic seas were at their freshest in the Late Cretaceous, some 80 Ma, not today. This is because a substantial part of the Mesozoic salt mass, deposited in the megahalites of the circum-Atlantic and circum-Tethyan basins, has since been recycled back into today’s ocean via a combination of dissolution and halokinesis. Periods characterised by marked decreases in salinity (Figure 9) define times of mega-evaporite precipitation, while periods of somewhat more gradual increases in salinity define times when portions of this salt were recycled back into the oceans (Chapter 5).


The last major extractions of salt from the ocean occurred during the late Miocene in the various Mediterranean Messinian basins created by the collision of Eurasia with North Africa. This was shortly after a large-scale extraction of ocean water from the ocean to the ice cap of Antarctica and the deposition of the Middle Miocene (Badenian) Red Sea rift evaporites. Accordingly, salinities in the early Miocene oceans were between 37‰ and 39‰ compared to the 35‰ of today (Figure 9). The preceding Mesozoic period was a time of generally declining salinity associated with the salt extractions in the opening North Atlantic and Gulf of Mexico (Middle to Late Jurassic) and South Atlantic (Early Cretaceous) and the earliest Cambrian oceans also had some of the highest salinities in the Phanerozoic. Recently, work by Blättler and Higgins (2014) utilising Ca isotopes studies of selected Phanerozoic evaporites has confirmed the dichotomous nature of Phanerozoic ocean chemistry that was previously defined by micro-inclusion studies of chevron halite (Figure 3).

So what?

In summary, based on a growing database of worldwide synchronous changes in brine chemistry in fluid inclusions in chevron halite, echinoid fragments, vein calcites at spreading centres and Ca isotope variations, most evaporite workers would now agree that there were secular changes in Phanerozoic seawater chemistry and salinity. Ocean chemistries ranged from MgSO4-enriched to MgSO4-depleted oceans, which in turn drove the two potash endmembers What is not yet clear is what is the dominant plate-scale driving mechanism (seafloor spreading versus dolomitisation versus uplift/weathering) that is driving these changes.

In terms of marine bitterns controlling favourable potash ore associations, it is now clear that the variation in ionic proportions in the original seawater controls whether or not potash-precipitating bitterns are sulphate enriched or sulphate depleted. A lack of MgSO4 minerals as co-precipitates in a sylvite ore makes the ore processing methodology cheaper and easier (Warren, 2016). Understanding the ionic proportion chemistry of Phanerozoic seawater is a useful first-order exploration tool in ranking potash-entraining evaporite basins across the Phanerozoic.

References

Ayora, C., J. Garciaveigas, and J. Pueyo, 1994, The chemical and hydrological evolution of an ancient potash-forming evaporite basin as constrained by mineral sequence, fluid inclusion composition, and numerical simulation: Geochimica et Cosmochimica Acta, v. 58, p. 3379-3394.

Blättler, C. L., and J. A. Higgins, 2014, Calcium isotopes in evaporites record variations in Phanerozoic seawater SO4 and Ca: Geology, v. 42, p. 711-714.

Borchert, H., 1977, On the formation of Lower Cretaceous potassium salts and tachyhydrite in the Sergipe Basin (Brazil) with some remarks on similar occurrences in West Africa (Gabon, Angola etc.), in D. D. Klemm, and H. J. Schneider, eds., Time and strata bound ore deposits.: Berlin, Germany, Springer-Verlag, p. 94-111.

Bots, P., L. G. Benning, R. E. M. Rickaby, and S. Shaw, 2011, The role of SO4 in the switch from calcite to aragonite seas: Geology, v. 39, p. 331-334.

Brennan, S. T., and T. K. Lowenstein, 2002, The major-ion composition of Silurian seawater: Geochimica et Cosmochimica Acta, v. 66, p. 2683-2700.

Dean, W. E., 1978, Theoretical versus observed successions from evaporation of seawater, in W. E. Dean, and B. C. Schreiber, eds., Marine evaporites., v. 4: Tulsa, OK, Soc. Econ. Paleontol. Mineral., Short Course Notes, p. 74-85.

Demicco, R. V., T. K. Lowenstein, L. A. Hardie, and R. J. Spencer, 2005, Model of seawater composition for the Phanerozoic: Geology, v. 33, p. 877-880.

Hardie, L. A., 1984, Evaporites: Marine or non-marine?: American Journal of Science, v. 284, p. 193-240.

Hardie, L. A., 1990, The roles of rifting and hydrothermal CaCl2 brines in the origin of potash evaporites: an hypothesis: American Journal of Science, v. 290, p. 43-106.

Hardie, L. A., 1996, Secular variation in seawater chemistry: an explanation for the coupled secular variation in the mineralogies of marine limestones and potash evaporites over the past 600 m.y.: Geology, v. 24, p. 279 - 283.

Hay, W. W., A. Migdisov, A. N. Balukhovsky, C. N. Wold, S. Flogel, and E. Soding, 2006, Evaporites and the salinity of the ocean during the Phanerozoic: Implications for climate, ocean circulation and life: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 240, p. 3-46.

Holland, H. D., J. Horita, and W. Seyfried, 1996, On the secular variations in the composition of Phanerozoic marine potash evaporites: Geology, v. 24, p. 993-996.

Holland, H. D., and H. Zimmermann, 2000, The Dolomite Problem Revisited: Int. Geol. Rev., v. 42, p. 481-490.

Holt, N. M., J. García-Veigas, T. K. Lowenstein, P. S. Giles, and S. Williams-Stroud, 2014, The major-ion composition of Carboniferous seawater: Geochimica et Cosmochimica Acta, v. 134, p. 317-334.

Kovalevych, V. M., T. M. Peryt, and O. I. Petrichenko, 1998, Secular variation in seawater chemistry during the Phanerozoic as indicated by brine inclusions in halite.: Journal of Geology, v. 106, p. 695-712.

Lowenstein, T. K., L. A. Hardie, M. N. Timofeeff, and R. V. Demicco, 2003, Secular variation in seawater chemistry and the origin of calcium chloride basinal brines: Geology, v. 31, p. 857-860.

Lowenstein, T. K., M. N. Timofeeff, S. T. Brennan, H. L. A., and R. V. Demicco, 2001, Oscillations in Phanerozoic seawater chemistry: Evidence from fluid inclusions: Science, v. 294, p. 1086-1088.

Müller, R. D., A. Dutkiewicz, M. Seton, and C. Gaina, 2013, Seawater chemistry driven by supercontinent assembly, breakup, and dispersal: Geology, v. 41, p. 907-910.

Paris, G., J. Gaillardet, and P. Louvat, 2010, Geological evolution of seawater boron isotopic composition recorded in evaporites: Geology, v. 38, p. 1035-1038.

Rausch, S., F. Böhm, W. Bach, A. Klügel, and A. Eisenhauer, 2013, Calcium carbonate veins in ocean crust record a threefold increase of seawater Mg/Ca in the past 30 million years: Earth and Planetary Science Letters, v. 362, p. 215-224.

Spencer, R. J., and L. A. Hardie, 1990, Contol of seawater composition by mixing of river waters and mid-ocean ridge hydrothermal brines, in R. J. Spencer, and I. M. Chou, eds., Fluid Mineral Interactions: A Tribute to H. P. Eugster, v. 2: San Antonio, Geochem. Soc. Spec. Publ., p. 409-419.

Timofeeff, M. N., T. K. Lowenstein, M. A. M. da Silva, and N. B. Harris, 2006, Secular variation in the major-ion chemistry of seawater: Evidence from fluid inclusions in Cretaceous halites: Geochimica et Cosmochimica Acta, v. 70, p. 1977-1994.

Warren, J. K., 2016, Evaporites: A compendium (ISBN 978-3-319-13511-3) Released November 2015: Berlin, Springer, 1600 p.

Wilson, T. P., and D. T. Long, 1993, Geochemistry and isotope chemistry of Ca-Na-Cl brines in Silurian Strata, Michigan Basin, USA: Applied Geochemistry, v. 8, p. 507-524.

Zimmermann, H., 2000a, On the origin of fluid inclusions in ancient halite - basic interpretation strategies, in R. M. Geertmann, ed., Salt 2000 - 8th World Salt Symposium Volume 1: Amsterdam, Elsevier, p. 199-203.

Zimmermann, H., 2000b, Tertiary seawater chemistry - Implications from primary fluid inclusions in marine halite: American Journal of Science, v. 300, p. 723-767.

Saline Clays

John Warren - Thursday, July 23, 2015

When discussing evaporites we typically focus on the formation and alteration of the various evaporite salts and their diagenetic evolution, but the same evolving saline hydrologies can also drive the formation and alteration of clays (Table 1). Many authigenic clay minerals formed in hypersaline settings are enriched in magnesium (Fisher, 1988), but authigenic clays do not make up the greater volumes of clay in modern or ancient salt lakes. Most of the clays in salt lakes and playas are detrital and reflect compositions of older argillaceous formations in the palaeodrainage areas. Illite, kaolinite, chlorite, dioctahedral smectite and a number of mixed-layers clays are commonplace detrital clay minerals in saline formations (Figure 1; Calvo et al., 1999). Widespread flocculation of clays is an effective sedimenter of suspended clay wherever freshwater runoff and streams flood an area of standing saline water. Thus the composition of initial clay sediments in a playa largely reflects that of the minerals carried as suspended load into the lacustrine depression.


The magnitude of detrital clastic input is thought to be a significant factor in the relative volume of authigenic clay. Regions with rapid deposition of clays, tied to high detrital inputs, tend to be areas where the authigenic clay component is swamped by the high detrital input. Clay authigenesis in evaporitic basins is favoured in marginal playa areas where rates of detrital clay input are low (Figure 1). This encompasses interdunal depressions, peripheral sandflats and muddy carbonate flats. In these low sedimentation areas the transformation of precursor clays is more effective, driven by episode surface inflow and groundwater discharge (Calvo et al., 1999). Highly reactive nearsurface and surface conditions are favoured by inherently large variations in pore water salinity, pH and pCO2 levels.

 

Clay authigenesis in many saline depressions is driven by pedogenesis, especially in the marginal areas where sedimentation rates are low and subaerial exposure dominates at the sedimentation surface. Below the surface episodic wet-dry cycles means neoformed clays are the byproduct of complex reactions between Na and Mg-rich interstitial brines and detrital silicates. Pedogenic processes account for the formation of widespread lake margin palygorskite and sepiolite, typically in association with the creation of calcretes, dolocretes and silcretes. In cases where palygorskite dominates the soil profile, they are sometimes described as palycretes. Zeolites can also form from saline groundwaters in saline lake-margin pedogenic settings (Figure 2). Artesian and phreatic groundwater discharge through springs into the lake margin areas also plays a significant role in the formation of other authigenic clays, as in saline lakes at the foot of Mt Kilimanjaro, in Tanzania and Kenya (Hay et al., 1995).

 

Hypersaline brines in modern, marine-edge evaporite basins can also enhance clay authigenesis even in settings where thermal and saline stresses keep both organic and inorganic carbon concentrations in the sediments unusually low relative to coastal marine environments with lower salinities (Martini et al., 2002). This is the case in Salina Ometepec where sediment pore waters exhibit little microbial sulphate reduction, and dissolved inorganic C contents are also very low. Instead of carbonate alteration (dolomitisation) in the Mg brine, authigenic K-rich Mg-smectite (saponite) formation is occurring, driven by the concurrent processes of brine concentration, selective dissolution of K- and Mg-bearing salts, and dissolution of detrital aluminosilicates. Salina Ometepec pore waters at a depth of 1 m have 87Sr/86Sr ratios that require input of Sr that is less radiogenic than that of Gulf of California seawater. This Sr is likely derived from weathering and leaching of detrital aluminosilicates from nearby volcaniclastic sources. Although rare in Holocene successions, similar Mg-rich authigenic clay assemblages are well documented in Palaeozoic evaporite basins (Bodine, 1983; Janks et al., 1992; Andreason,1992).

Once precipitated in an evaporite basin, authigenic clays can be retransported further out into the saline depression and in more humid climatic stages may even end up on the floor of freshwater lakes (Figure 1). This situation is seen in lacustrine sequences from the Miocene formations of the Madrid Basin (Bellanca et al., 1992) where significant amounts of palygorskite and sepiolite occur as either mud chips or clay aggregates in the basal part of a fresher water lacustrine unit. Eolian transport of saltating clay pellets or dust suspensions may also contribute to the transport of authigenic clays from marginal to more central areas. This sometimes leads to problems of interpretation of detrital versus authigenic in ancient lacustrine successions subject to oscillations in climate, especially when detrital clays are partially or fully inherited from arid soils.

Sepiolite, interstratified Mg-Smectite and palygorskite form authigenic phases in the Quaternary sediments of the Double Lakes Formation, Texas (Webster and Jones, 1994). The dominance of each of these minerals in separate horizons represents evaporative shifts in salinity at the time they precipitated. Sepiolite is thought to indicate a brackish lake, while Mg-smectite indicates more saline conditions. Palygorskite is interpreted as a saline pore water precipitate in the arid soils of the playa stage. Likewise Jones (1986) interpreted authigenic Mg-smectites (e.g. stevensite) as requiring higher salinity than sepiolite. Mg-silicates also define saline lake clays in Great Salt Lake (Spencer, 1983) and some Bolivian salars (Badaut and Risacher, 1983). In Bolivia, the authigenic Mg-smectite replaces the biogenic silica in diatom frustules and requires a pH in excess of 8.2. Authigenic stevensite occurs in unconsolidated muds underlying saline crusts in the interdunal depressions of northern Lake Chad and as small aragonite-associated oolites on the lake floor (Gac 1980, Darragi and Tardy, 1987). Similar stevensite oolites have been found in the Eocene Green River lacustrine basin. Stevensite is also an early authigenic phase in the modern carbonate thrombolites in the hyposaline Lake Clifton, Australia (Burne et al., 2014). Authigenic sepiolite associated with calcite, gypsum and dolomite occurs about the margin of Saline Valley Playa, California and the edges of saline pans in the Kalahari of southern Africa (Hardie, 1968; Kautz and Porada, 1976). Palygorskite, sepiolite and authigenic smectite are commonplace precipitates in calcretes of groundwater discharge playas in inland Australia (Arakel et al., 1990).

Clearly, palygorskite and sepiolite (both two-chain structure fibrous clays) occur worldwide as authigenic phases in the soils and palaeosols of arid and semi-arid regions, but the mode of precipitation is still not well understood (Singer, 1979). Both minerals are common in environments with elevated levels of magnesium and silica. Hence they form in alkaline lakes and caliche, as well as in deep sea sediments and Hydrothermal alteration products; Folk and Rasbury (2007) argue there may also be a microbial association to their formation, at least in some Texan caliches. Jones (1986) concluded sepiolite in the calcic soils of southwest Nevada required percolation of high salinity groundwaters. Magnesium and silica solutes were supplied by the weathering of nearby pyroclastics and carbonates. Sepiolite has replaced magnesite pebbles, from the edges in, during freshened highstand intervals in Miocene Lake Eskisehir in Turkey (Ece, 1998). Palygorskite in calcic soils is thought to be the result of incongruent dissolution of pre-existing clays (Jones and Galán, 1988). Fibrous clays degrade when climate becomes more humid and alter to smectite. Paquet and Millot (1972) conclude that the transformation takes place when mean rainfall exceeds 300 mm and Calvo et al. (1999) suggested the transformation can be used as a palaeoclimatic indicator.

Alunite (KAl3(SO4)2(OH)6) is a common clay product in acid saline lacustrine settings, but can also form diagenetically in regions where sulphate reduction is occurring. It is thought to be derived by the reaction of clay minerals with sulphuric acid created by oxidation of sulphides or H2S at a redox boundary. It is a common product where clays are present in zones of sulphate reduction and examples have been documented in the Middle Miocene gypsums of the Gulf of Suez (Rouchy et al., 1995) and the Upper Miocene gypsums of the Lorca Basin in Spain (Rouchy et al., 1998).

Even the smectite to illite transformation, which is used as an indicator of diagenetic intensity and clay transformations occurring at higher temperatures may be influenced by salinity. This makes illite crystallinity a less reliable indicator of diagenetic stage in environments with saline pore fluids (Honty et al., 2004). Turner and Fishman (1991) found illite-smectite mixed layer clays having a range of expandabilities in altered tuff beds in a Jurassic lake in the Morrison Formation (Eastern Colorado Plateau, USA). The observed clays did not experience deep burial, and did not undergo hydrothermal alteration. The illite content generally increases from the lake margin (100–70% smectite) to the lake centre (30–0% smectite) and follows a lateral hydrogeochemical gradient, which was characterized by increasing salinity and alkalinity (Figure 3). It seems that in a saline depositional setting, solution chemistry is a principal factor controlling the smectite to illite proportion. Illite-smectite can form from smectite at low temperatures in several ways (see Honty et al., 2004), but forms best in saline environments subject to wetting and drying cycles, which is a hydrology exemplified in salt lakes and playas. In the presence of K+ ions, alternating wetting and drying leads to irreversible fixation of K and the formation of illite layers. Illite-smectite clays forming at elevated pH may not even require wetting and drying cycles.


References

Andreason, M. W., 1992, Coastal siliciclastic sabkhas and related evaporative environments of the Permian Yates Formation, North Ward-Estes field, Ward County, Texas: American Association of Petroleum Geologists Bulletin, v. 76, p. 1735-1759.

Arakel, A. V., G. Jacobson, and W. B. Lyons, 1990, Sediment-water interaction as a control on geochemical evolution of playa lake systems in the Australian arid interior: Hydrobiologia, v. 197, p. 1-12.

Badaut, D., and F. Risacher, 1983, Authigenic smectite on diatom frustules in Bolivian saline lakes: Geochemica et Cosmochimica Acta, v. 47, p. 363-375.

Bellanca, A., J. P. Calvo, P. Censi, R. Neri, and M. Pozo, 1992, Recognition of lake-level changes in Miocene lacustrine units, Madrid Basin, Spain. Evidence from facies analysis, isotope geochemistry and clay mineralogy: Sedimentary Geology, v. 76, p. 135-153.

Bodine Jr, M. W., 1983, Trioctahedral clay mineral assemblages in Paleozoic marine evaporite rocks: Sixth international symposium on salt, v. 1, p. 267-284.

Burne, R. V., L. S. Moore, A. G. Christy, U. Troitzsch, P. L. King, A. M. Carnerup, and P. J. Hamilton, 2014, Stevensite in the modern thrombolites of Lake Clifton, Western Australia: A missing link in microbialite mineralization?: Geology, v. 42, p. 575-578.

Calvo, J. P., M. M. Blanc-Valleron, J. P. Rodriguez Arandia, J. M. Rouchy, and M. E. Sanz, 1999, Authigenic clay minerals in continental evaporitic environments: Special Pulication Internation Association Sedimentologists, v. 27, p. 129-151.

Darragi, F., and Y. Tardy, 1987, Authigenic trioctohedral smectites controlling pH, alkalinity silica and magnesium concentrations in alkaline lakes: Chemical Geology, v. 63, p. 59-72.

Deer, A., R. Howie, W. S. Wise, and J. Zussman, 2004, Rock Forming Minerals. vol. 4B. Framework Silicates: Silica Minerals, Feldspathoids and the Zeolites: London, The Geological Society.

Ece, O. I., 1998, Diagenetic transformation of magnesite pebbles and cobbles to sepiolite (Meerschaum) in the Miocene Eskisehir lacustrine basin, Turkey: Clays & Clay Minerals, v. 46, p. 436-445.

Fisher, R. S., 1988, Clay minerals in evaporite host rocks, Palo Duro Basin, Texas Panhandle: Journal of Sedimentary Petrology, v. 58, p. 836-844.

Folk, R., and E. Rasbury, 2007, Nanostructure of palygorskite/sepiolite in Texas caliche: Possible bacterial origin: Carbonates and Evaporites, v. 22, p. 113-122.

Gac, J. Y., 1980, Géochimie du bassin du Lac Tchad: Travaux et Documents ORSTOM, v. 123, p. 54 pp.

Hardie, L. A., 1968, The origin of the Recent non-marine evaporite deposit of Saline Valley, Inyo County, California: Geochimica et Cosmochimica Acta, v. 32, p. 1279-1301.

Hay, R. L., R. E. Hughes, K. T. K., H. D. Glass, and J. Liu, 1995, Magnesium-rich clays of the Meerschaum Mines in the Amboseli, Tanzania and Kenya: Clays & Clay Minerals, v. 43, p. 455-466.

Honty, M., P. Uhlik, V. Sucha, M. Caplovicova, J. Francu, N. Clauer, and A. Biron, 2004, Smectite-to-illite alteration in salt-bearing bentonites (The East Slovak Basin): Clays and Clay Minerals, v. 52, p. 533-551.

Janks, J. S., M. R. Yusas, and C. M. Hall, 1992, Clay mineralogy of an interbedded sandstone, dolomite, and anhydrite; The Permian Yates Formation, Winkler County, Texas, Origin, Diagenesis, and Petrophysics of Clay Minerals in Sandstones, v. 47, SEPM (Society for Sedimentary Geology), p. 145-157.

Jones, B. F., 1986, Clay mineral diagenesis in lacustrine sediments, in F. A. Mumpton, ed., Studies in diagenesis, v. 1578, US Geological Survey Bulletin, p. 291-300.

Jones, B. F., and E. Galán, 1988, Sepiolite and palygorskite, in S. W. Bailey, ed., Hydrous Phyllosilicates (Exclusive of Micas), v. 19: Washington, Mineralogical Society of America, Reviews in Mineralogy, p. 631-674.

Kautz, K., and H. Porada, 1976, Sepiolite formation in a pan of the Kalahari: Nues Jahrb. Mineral Monatsh., v. 12, p. 545-559.

Martini, A. M., L. M. Walter, T. W. Lyons, V. C. Hover, and J. Hansen, 2002, Significance of early-diagenetic water-rock interactions in a modern marine siliciclastic/evaporite environment: Salina Ometepec, Baja California: Geological Society America Bulletin, v. 114, p. 1055-1069.

Paquet, H., and G. Millot, 1972, Geochemical evolution of clay minerals in the weathered products of Mediterranean climate: Proceedings, International Clay Conference, Madrid, 21-24 June, 1972, p. 199-206.

Rouchy, J. M., D. Noel, A. M. A. Wali, and M. A. M. Aref, 1995, Evaporitic and biosiliceous cyclic sedimentation in the Miocene of the Gulf of Suez; depositional and diagenetic aspects: Sedimentary Geology, v. 94, p. 277-297.

Rouchy, J. M., C. Taberner, M. M. Blanc-Valleron, R. Sprovieri, M. Russell, C. Pierre, E. Di Stefano, J. J. Pueyo, A. Caruso, J. Dinares-Turell, E. Gomis-Coll, G. A. Wolff, G. Cespuglio, P. Ditchfield, S. Pestrea, N. Combourieu-Nebout, C. Santisteban, and J. O. Grimalt, 1998, Sedimentary and diagenetic markers of the restriction in a marine basin: the Lorca Basin (SE Spain) during the Messinian: Sedimentary Geology, v. 121, p. 23-55.

Singer, A., 1979, Palygorskite in sediments: detrital, diagenetic or neoformed - A critical review: Geologische Rundschau, v. 68, p. 996-1008.

Spencer, R. J., 1983, The geochemical evolution of Great Salt Lake, Utah: Doctoral thesis, Johns Hopkins University.

Turner, C. E., and N. S. Fishman, 1991, Jurassic Lake T'oo'dichi': A large alkaline, saline lake, Morrison Formation, eastern Colorado Plateau: Geological Society of America Bulletin, v. 103, p. 538-558.

Turner, C. E., and N. S. Fishman, 1991, Jurassic Lake T'oo'dichi': A large alkaline, saline lake, Morrison Formation, eastern Colorado Plateau: Geological Society of America Bulletin, v. 103, p. 538-558.

Webster, D. M., and B. F. Jones, 1994, Paleoenvironmental implicationsof lacustrine clay minerals from the Double Lakes Formation, southern Great Plains, Texas., in R. W. Renaut, and W. M. Last, eds., Sedimentology and geochemistry of modern and ancient saline lakes, v. 50: Tulsa, Society Economic Paleontologists and Mineralogists Special Publication, p. 661-686.

Salt's uses across human history

John Warren - Wednesday, May 13, 2015

Salt's uses across human history

Until the 19th Century and the advent of refrigeration, salt's main uses was as a preservative and as a much sought after flavouring for foodstuffs. Even today, most people think of salt in terms of sprinkling it on their food, or in colder climates they may also think in terms of road de-icing. These are in fact lesser modern usages of halite; its main use is as a feedstock in the chemical industry, a topic we shall discuss in a later blog. This essay focuses on salt's importance to humanity prior to the chemical age.


Sodium chloride (halite), the most common industrial evaporite salt and is used in some form by virtually every person in the world. The human body contains about 110 gm of salt. Salt is essential to all living creatures and even many plants. Since the body cannot manufacture it, salt is an "essential" nutrients, and as an electrolyte, we lose it every time we sweat. Without enough salt, muscles won't contract, blood doesn’t circulate, food goes undigested, and ultimately the heart ceases to beat.

Halite, along with other salts, has long played a very important role in human affairs. Early hominids lived on the edge of the saline Lake Olduvai (Hay and Kyser, 2001) and salt was part of their diet. In ancient Greece it was so valuable that the slave trade involved an exchange of salt for a slave and gave rise to the expression, “not worth his salt.” Some 4,700 years ago the Peng-Tzao-Kan-Mu was published in China. It is probably the earliest known treatise on pharmacology, with detailed discussions of the palliative and curative powers of more than 40 kinds of salts, including descriptions of two methods of recovering usable salts from brine. There are more than 14,000 reported usages of halite and more than 30 references to salt in the Bible. Some 3,200 years ago, near Hallstatt in Austria, Bronze-Age miners were extracting salt, from a network of several kilometres of galleries up to 300 metres below the surface. Two thousand seven hundred years later some ten pages deal with salt in the “De Re Metallica” by Georgius Agricola (Georg Bauer), both mining it and producing it from seawater or brine springs. Published in 1556, Agricola’s was the first book on mining to be based on field research and observation.

 

Salt as a preservative

Historically, whether food was hunted, gathered, or grown and harvested, food supply was rarely available year-round to all members of a society. Yet, effective, year-round, reliable food storage was vital, especially for non-nomadic agricultural societies. Today, to maintain reliable food supplies to our ever-expanding urban populations, we refrigerate, freeze-dry or can our food. Food preservation problems seem trivial to most consumers in the developed world, outside of the world’s war zones. But, prior to the 19th century, effective food storage often made the difference between life and death to large segments of the world's human population.

In arid climates, food can effectively be stored by drying. But in more humid temperate climates, fungus and bacteria rapidly destroy stored and cellared food. Even where food can be stored in winter ice, it quickly rots when spring thaws set in. Documents from northern Europe, give some clues about the severity of the problem, and its solution. In medieval societies, with relatively poor transportation systems, villages and counties had to be close to self-sufficient in food. If a bad harvest occurred, to mitigate a potential disaster there had to be enough food stored to tide over until the next harvest. Medieval Europe offers an example of the way in which agricultural societies dealt with food security. Good-quality arable land was scarce, and had to be used for crops. That meant that grazing and foraging animals, mainly cattle and pigs, were turned out into the local woodlands for the summer to forage for grass, roots, and nuts. Any relative shortage of winter fodder in turn meant that surplus animals would best be butchered before cold weather drove them indoors. In medieval England the annual slaughter was traditionally around Martinmas, St Martin's Day (10 November), but it was earlier in the colder climes of Sweden. In turn, that meant that fresh meat was readily available only at that time, and that fresh protein in the form of milk and butter was only available in winter from cows kept in shelter. In addition, taxes were often paid in kind rather than money, and that meant that the landlord had to be given foodstuffs that could be stored.

The response of the Swedes and most northern Europeans, was to preserve almost all their food, and they used salt to do so. Beef and pork were salted and dried as joints, hams, and sausages. Butter was salted. Typically it took a pound of salt to preserve 10 pounds of butter (salt was sufficiently costly that housewives removed salt before they used stored butter). Fish, whether freshwater or from the sea, were salted and dried, and bread was salted and hung to dry. Surviving records from 1573 show the servants of King Gustavus Vasa of Sweden ate some 102 kg (224 pounds) of beef and pork, but 99 kg of it (218 pounds) was salted and dried. They almost never had fresh meat. The King issued an order to release 3-year-old butter from the tax stores for some men hired to work at the castle, and ordered the sale of 4-year-old (barley) malt because it was starting to get weevils in it. He ordered the peasants to store their butter and meat in the fall, after the annual slaughter, but he also ordered them not to eat any of it for 12 months (as they should be eating the previous year's food during that time).

 

Outside of salted food storage, the ancient Egyptians are famous for their perfection of the art of mummification. A key ingredient in the process was natrun, which is a natural mixture of halite, trona and sodium sulphate (Edwards et al., 2007). The ancients knew its preservative properties as it readily absorbs water, making it an excellent desiccant/preservative of organic material. Natrun is found in large quantities in the beds of several Egyptian playa lakes (e.g. Wadi Natrun and El Kab, as well as Behiera in the nearby Libyan desert and in Lake Natron in the African Rift valley; Figure 2). It has been mined and traded from such localities for thousands of years. Writings as old as the reign of Rameses III (1198-1166 B.C.) refer to natrun deposits. Its preservative qualities must have been immediately apparent to the ancient Egyptians from its effects on any wild life, which had died in these lakes (Figure 2). There is some evidence that the ancient Egyptians artificially precipitated natrun by isolating shallow basins of salt lake waters for faster evaporation, as is still done in parts of the Faiyum depression today. For purification and preservation, natrun was preferred over pure halite as it chemically attacks and destroys grease and fat, and so is a superior drying agent (as is sodium borate). Its residues are found not only in tombs and in pits, along with other discarded embalming materials, but also forms nodules and residues in the mummies themselves.

There is some popular debate over the method in which the natrun was used for mummification by the ancient Egyptians. Some argue it was used in a way similar to the contemporary method for “salting” fish. Dry natrun would be sprinkled over the body, perhaps with sawdust, or spread with linen cloth wraps. Others with a more starry-eyed bent, believe the body was immersed in vats containing a natrun solution. Such a wet method would have been odiferous and accelerated putrefaction, thus counterproductive to the preservation of the body, although it makes for good Hollywood images. A dry body is also more readily bandaged as well as being more amenable to the attachment of amulets and other jewellery. Although mummification has supernatural trappings in popular culture and ancient religions, its basis is rooted in simple chemistry and processes as mundane as salting fish.

 

The mummies of some Buddhist monks (Sokushinbutsu sect) in Japan resulted from the practise of nyūjō, which ultimately aimed to cause their own death and mummification by encasement in salt. This ritual took years to complete and involved starvation and dehydration. During the first three years, an ascetic monk significantly decreased his body fat by eating only nuts, seeds, and berries, while he increased his physical activity. Towards the end of the ritual the monk reduced his food intake even further by only consuming bark, roots, and sometimes stones. Post-mortem preservation was further aided by consumption of toxic herbs and tea that eliminated bodily fluids and killed bacteria that aid in decomposition. Japanese Sokushinbutsu monks were known to drink a tea made from the urushi tree, also known as the Chinese lacquer tree because it’s sap is used to lacquer tableware, instruments, and jewellery.

After years of starvation and dehydration, when the monk felt like he was close to death, his fellow monks arranged his body in the lotus position inside a coffin or a tomb. Then they surrounded the dying man with salt, wood, paper, or lime to pull more moisture away from the body and further prevent post-mortem decay. Only a small opening for air was allowed when the tomb was closed. The monk then chanted, meditated and occasionally rang a bell until he died.

Once his fellow monks heard silence they completely sealed the tomb. After several years, the monks exhumed the body to see if the self-mummification ritual was successful. Like some Eastern Orthodox religions, these Buddhists believed that an incorrupt body, a body having delayed decomposition, indicated a monk’s holiness. If the body was incorrupt after exhumation then the corpse was placed in a temple, adorned, and tended to by followers. However, if a tomb was opened and the body had decayed, then the corpse was left behind and the tomb was resealed. That monk’s efforts were respected, but his body was not given the deference of a religious relic. Japan banned unburying in 1879 and assisted suicide, including religious suicide is now illegal. In a similar vein, in 1933, the Dalai Lama was buried sitting up in a bed of salt.

Mummification can occur naturally if a body is encased in halite, and natrun is not necessary, although it improves preservation. In 1593 AD, and again in 1616 AD, several tombs encased in salt were exposed by natural salt weathering and collapse in the Hazel Mountains. When the coffins were opened by the local people of Hallein and Hallstatt, there was astonishment that the bodies inside had very well preserved soft tissues. It was the result of the hyperarid encasement in a Neolithic salt mine, but frightened religious locals, encouraged by the local clergy, insisted on prompt reburial, along with additional religious efforts to lessen ambient sin levels (in part in the form of alms to mother church) and hence more prayer to create more effective seals. There was a similar popular response in 1734 AD when the salt preserved body of a man wearing mountain clothing (likely a salt miner) was discovered. Fearful locals, once again encouraged by the local clergy, insisted on immediate reburial with no further scientific study or observations on the remains (Aufderheide, 2011). 

In Iran, first in the winter of 1993 and later in 2004, in the modern Chehr Abad Salt Mine, near Hamzehloo, Zanjan Province, a total five salt-preserved male bodies were found in a collapsed tunnel of a former salt mine, which was active around 400 BC. The first discovery in the winter of 1993 was a salt encased bearded head and some artefacts, the later discovery, beginning in November 2004, was of the remaining bodies. It is likely all five men died in earthquake induced collapses in the salt mine (Pollard et al., 2008). Encasement in the hyperarid atmosphere of the collapsed salt mine tunnel led to natural mummification of the bodies.

Salt and war

Salt’s historical use as a food preservative, along with its medicinal use, made it a valuable commodity with political and military significance. The earliest recorded war over access to a supply of salt was over a salt lake in China in 3000 BC. In 2200 BC the Chinese emperor Hsia Yua declared that Shandong Province must supply the Imperial Court with salt. An ancient Chinese philosopher once called salt “the sweetest thing on earth.” The words, “war” and “peace” originate from the words for salt and bread in ancient Hebrew and Arabic, while from the Latin "sal," came words such as "sauce" and “sausage."

As an example of salt’s military import let’s look at the significance of a reliable salt supply to the army of the Old South in the 1860s during the American Civil War. Each Confederate soldier was provided with starch (26 pounds of coarse meal, 7 pounds of flour or biscuit, 3 pounds of rice), protein (10 pounds of bacon), and salt (one and a half pounds). Bacon was the meat of the South, and every pound of it required salt. As well as military personnel, horses also need salt in their diet. The Confederacy also needed this precious mineral to treat wounds, tan leather and dye cloth for uniforms. Last century, the historian Ella Lonn (1933) devoted an entire book to the problem of reliable salt supply for the Confederacy during the Civil War. We know that the Confederate soldiers were hungrier than the Northerners throughout the war. We shall never know whether the hogs that were not slaughtered because there was no salt to preserve them took the edge off the Confederate troops, or whether the salt that was not available for the horses took the edge off the cavalry. "What hogs we have to make our meat, we can't get salt to salt it," wrote Mrs Sarah Brown to Governor Pettus of Mississippi in December 1861. In 1862, Governor Brown of Georgia wrote that only half of the meat of the State could be saved for the 1862-1863 season.

 

That most intelligent and brutally efficient of the Northern Generals, Sherman, had no doubt about salt’s importance to any army and its morale, he considered it as important as gunpowder, he declared. "Without salt they cannot make bacon and salt beef," and, "Salt is eminently contraband, because [of] its use in curing meats, without which armies cannot be subsisted." Sherman sent a captain for trial on a charge of aiding the enemy, because he had allowed salt through the lines to the Confederates. The Union forces were sent orders to destroy salt stores and salt works wherever they were found (Figure 4). Throughout the American Civil War the South’s salt production facilities in Saltville, Va., Virginia's Kanawha Valley and Avery Island, Louisiana, were targets of the Union Army. The North fought for 36 hours to capture Saltville, Va., where the salt works were considered so crucial that Confederate President Jefferson Davis offered to waive military service to anyone willing to tend coastal salt kettles and so supply the South's war effort.

In November 1863, General Burnside noted in a despatch to Grant that Lee had placed a strong defensive force in front of Saltville [Virginia]. Grant understood the significance of the deployment. In December 1863 he wrote to General Foster, "If your troops can get as far as Saltville and destroy the works there, it will be an immense loss to the enemy." In the event, the Confederates guarded the works so well that the Union Army did not take (and destroy) the salt works until December 1864. General Burbridge boasted that the loss of Saltville would be "more felt by the enemy than the loss of Richmond." Meanwhile the North, even with salt sources of its own, imported 86,208 tons of salt from England in 1864 alone.

Likewise, thousands of Napoleon’s troops died during his retreat from Moscow, because for lack of salt their wounds would not heal.

Ancient salt production and its taxable value

Around 6,000 BC on the margins of Lake Yun Cheng in Northern China’s we see the first evidence of an industry designed to harvest and produce salt, via the evaporation of lake brines in purpose-built salt pans. In Europe, the first recorded industrial production of pan salt took place in Italy some 2500 years ago when Ancus Martius, one of the early Roman kings, began letting sea water into an enclosed basin, then allowing the sun to evaporate the water to create a salt residue. The importance that Rome attached to the salt works and port at Ostia was such that the main highway along which the salt was carried to Rome was called the Via Salaria. Like Venice after it, the city of Rome based much of its early commerce on trading salt. Special salt rations paid to early Roman soldiers were known as “salarium argentum”, the forerunner of the English word “salary.” With a near monopoly on supply to Rome, the traders in the port of Ostia raised the salt price so high that the state was forced to take over the industry in 506 BC.

When Julius Caesar landed in Britain in 55 BC, he brought his salinators with him, but found that even the backward Britons were extracting salt by pouring brine on to hot stones. The Romans, however, used iron pans in which they boiled the brine, and Caesar established a brine-based salt works in Cheshire and subsequently in other localities where ancient salt occurred at shallow depths. The towns in Britain where salt was made from brines extracted from shallow buried ancient salt beds can be distinguished to this day by the termination “wich”, an Anglo-Saxon descriptor for a place where salt was made and includes towns like Greenwich, Ipswich, Droitwich, Northwich and Middlewich. Likewise, within regions of shallow salt and brines in Austria and Germany, names containing "salz" and "halle," such as Salzburg ("salt city"), Salzkammergut, Reichenhall, Halle, Hallein, and Hallstatt, as well as the old Austrian/Polish province of Galicia, identify some of the salt-bearing areas.

Merchants in 12th-Century Timbuktu in Africa, the gateway to the Sahara Desert was the seat of renowned scholars, who valued salt extracted from salt lakes to north in the vicinity of Taoudenni, Mali, as highly as books or gold. The Taoudenni mines are located on the bed of an ancient erg-edge salt lake and have been actively quarried for more than a 1,000 years. Today, the miners use crude axes to dig pits that usually measure 5 m by 5 m with a depth down to around 4 m. The miners first remove up to 1.5 m of red clay overburden, in contrast to the salt miners in the Danakil who work at the active pan surface (see Salty Matters, 19 April, 2015). Then several layers of poor quality salt are removed before reaching three layers of high quality salt. The salt is cut into slabs that are 110 cm x 45 cm by 5 cm in thickness and weigh around 30 kg. Two of the high quality layers are of sufficient thickness to be split in half, so that 5 slabs can be produced from the three layers. Having removed the salt from the base area of the pit, the miners excavate horizontally to create galleries from which additional slabs can be obtained.  As each pit is exhausted, another is dug so there are now thousands of pits spread over a wide area on the lake. Over the centuries salt has been extracted from three distinct areas of the lake depression, with each successive area located further to the southwest. The  areas can be clearly seen on satellite photographs (22.606519°E, 4.030660°S). Until recently salt was transported south by huge camel trains, now more and more salt is carried out by 4-wheel drive trucks, south to Timbucktu and on to the river port of Mopti (Figure 5).  Among the some of the nomadic tribes of the Sahara and Ethiopia's Danakil Plains, salt carried by camel trains and is still used occasionally as money or bartered for a cash equivalent. When the camel trains of Mali carried the salt, each animal typically carried 4 blocks of salt. On reaching the salt market, three blocks sold off the back of each animal went to the camel train owners, and the profit of the sale of the remaining block to the salt miner. 

 

In Tibet, Marco Polo noted that tiny cakes of salt, manufactured from salt lakes in the high plains of Tibet were pressed with images of the Grand Khan and used as coins.The ancient Maya made salt at Salinas de los Nueve Cerros, Guatemala, an area where natural salt springs flowed into a river gully, giving easy trading access to downstream customers (Figure 6). This site was the only large-scale source of salt for the interior Lowland Maya. Maya technology included solar evaporation and firing of brine from salt springs in special large ceramic bowls that are the largest receptacles ever found in any Maya sites.

The highly organized salt trade of China was observed by Marco Polo, who recorded that the major item of trade on the Yangtze River was salt, shipped upstream from the coast (especially from the city of Hangchou) to the interior cities. The Chinese produced salt by many methods: they evaporated it, boiled seawater, and pumped brine from wells drilled into salt beds. Modern oil-drilling traces its roots back to Chinese methods of bamboo-based drilling technology that originally evolved for salt production from ancient subsurface brine sources (this will be the subject of a later blog).

 

Salt production, politics and taxes

Salt's economic value has meant it has been taxed by governments from the ancient Chinese and Romans to governments of late medieval Europe to those of France, even up to the late 1940s. In 2200 BC the Chinese Emperor Hsia Yu levied a salt tax, which was one of the world's first documented state taxes.

The Mediterranean and the rise of Venice

The great trading ports of the Mediterranean dealt in salt as well as spices and textiles. Not surprisingly, the greater of them, Genoa and Venice not only traded in salt, but fought for supremacy over the trade. Because of the hot dry summers and mild wet winters, salt can be made in a saltern or pan in almost any suitable seashore flat or plain in the Mediterranean. So although it is possible to envisage a trader's cartel from a specific geological region of shallow buried salt in Austria or England, it is much more difficult to control the production of salt in coastal saltpans. So, in hindsight, it is surprising how effectively Genoa and especially Venice, managed to take control of Mediterranean salt production, as well as trading, across the 13th to 16th centuries. Genoa was positioned in the Western Mediterranean and Venice at the head of the Adriatic. Each used all its political and military strength to consolidate its local salt trade, and to encroach as far as possible on that of its rival. However, Venice was more organised politically, which translated into more ruthless and effective use of state power. And Venice made a conscious decision to concentrate on the salt trade, whereas to the Genoese it was just one of a set of potentially profitable cargoes. Where the two came into conflict over salt, the Venetians tended to win.

Venice managed to make a business out of control of the Adriatic salt trade. Venice owed some of its early wealth to the salt trade from salt works in its lagoon, and had a number of contracts with inland Italian cities in the 13th century to supply them with salt. The more that Venice came to control the salt trade in the Adriatic, the more the resulting profits were used by the city to subsidise other trading activities. Venetian traders delivering salt to the city were given bank credits, for example, allowing them to buy goods quickly. As the historian S. A. Adshead has written, "For the Venetians, salt was not a commodity among commodities... it greased the wheels of all the working parts and fuelled its motor". The salt trade allowed Venetian traders to compete very effectively with their rivals across the board. Salt was "il vero fondamento del nostro stato” (The true foundation of our state).


Always, from their beginnings in the 5th Century, the Venetians were willing to exercise raw power to foster their control of salt. Prior to the rise of the Venetian State, the Roman salt-making center in the Adriatic was at Comacchio, a little north of Ravenna. After the fall of Rome, records of the 8th-century Lombard King Luitpold show that Comacchian salt was being shipped to all the major inland cities of Lombardy, through Ferrara, at least as far inland as Parma, Lodi, and Brescia. By 523 AD Venice was producing salt and in 932 AD the Venetians destroyed Camacchio. They burned the citadel, massacred the inhabitants, and carried off the survivors to Venice, where they had to swear an oath of loyalty to the Doge before they were released. The Venetians began to construct salt works on their own lagoon, and around 1028, we find the Doge of Venice giving permission for Chioggia to build more salines on the Venetian lagoon. However, it turned out that it was not as easy to build salt works in the relatively exposed, storm-prone lagoon of Venice as it had been at Comacchio, and it took a long time before salt production became really successful at Chioggia. Meanwhile, the city of Cervia, south of Ravenna, filled the salt production vacuum left by the destruction of Comacchio and Cervia was in full production at least by 965- ­975 AD.

Around 1180, it was clear that Cervia and Chioggia were rivals for salt production, under the protection of Ravenna and Venice respectively. The Archbishop of Ravenna and the Doge of Venice now began exerting political pressure on the Adriatic salt market. Venice declared it illegal for Chioggia salt to be sold or shipped without a Venetian certificate, and Ravenna did the same for Cervia. The salt market was now out of the hands of merchants and in the hands of the politicians and the Catholic church. By 1234, war between Venice and Ravenna ended with a ban on any Ravenna (Cervia) salt being shipped northward, and Venetian galleys enforced the treaty.

Then, the Venetians went one logical step further: for all practical purposes they gave up trying to be salt producers, and instead concentrated on being (monopoly) salt traders. Between 1250 and 1280, they came more and more to be the dominant buyers of salt, which they then warehoused, shipped and sold (Figure 7). By the 1350s, no salt could move on a ship in the Adriatic unless it was a Venetian ship bound to or from Venice.

A golden rule of Venetian policy was that all trade goods under their control must pass through Venice. As late as 1590 they were making an 81% mark-up on salt sold inland. But that was not always the case, sometimes, if it would foster trade in higher-value goods that would yield more profit, Venice sold salt at less than normal rates. All this activity was planned and supervised by a special State body, the Collegio del Sal. The rewards were staggering, and help to justify the tenacity and ruthlessness with which the Venetians operated the salt business. Typically, Venetian merchants bought salt for 1 ducat a ton, and it cost them about 3 ducats a ton to ship it to Venice. There they received a State subsidy of 8 ducats a ton. The State collected a tax as the salt left Venice, and after shipping to the customer, the selling price was roughly 33 ducats a ton. That was a profit worth fighting for! And it was not only the merchants who profited. Some of the State profits went to the architecture, sculpture, and paintings that remain today and make Venice so magnificent (Figure 7).

The Venetians had different methods for maintaining their trading monopoly. On the island of Pag, they would buy up all the salt that was not needed locally. It would then be shipped to Venice, warehoused and sold (at very high prices) to customers. At Muggia and Capodistria, the Venetians were given a fraction (about 10%) of the salt produced (presumably as protection money), but the locals were allowed to sell the other 90% only as long as it was carried overland, effectively limiting its value and the sales area.

As late as 1578, the Venetians destroyed the salt works at Trieste, and in the following twenty years were making an 80% profit on salt sold inland on the Lombardy plain. But around 1600, paradoxically with the defeat of the Turks at sea, shipping intensity in the Adriatic became too great for the Venetians to be able to maintain their monopoly by force. Their source of riches in the spice trade had also been cut off as the trade routes to India now passed around Africa, and so their shipping power and wealth declined.

Salt and wealth in inland Europe and the UK

Much of the salt supply of inland central and northern Europe came from the mining of shallowly buried ancient salt (Permian) or associated brines. The great salt extraction centre at Reichenhall, in southern Bavaria, was first operated in Roman times, but was destroyed later, possibly by Attila the Hun, but more likely by the German Odoacer. It was rebuilt in the early 7th Century by Saint Rupert of Salzburg (Figure 8) and became the concession of the Bishop of Salzburg, who derived a great deal of power and money from the salt trade. So mother Church promoted the “salt bishops” to Archbishops. About 1190, however, a competing salt works had opened at nearby Berchtesgaden, without the Archbishop's approval, and a major quarrel between Church and State erupted, with the Archbishop and the Emperor in conflict. The Church lost, and in 1198 the Bavarian saltworks passed into the control of the Duke of Bavaria. Reichenhall's production peaked at about this time, and it later lost out in competition with a new salt works opened to the south by the persistent Archbishops of Salzburg. During that time it remained an important salt centre for several hundred more years and, even today, derives income from geotourism and from the therapeutic salt baths of Heilbaden.

 

Thwarted in Bavaria, the Archbishop of Salzburg turned to salt springs closer by, and so a new salt industry sprang up at Hallein, first mentioned in documents in 1232. By 1300 its production had outstripped that of Reichenhall, and as it was situated closer to the Danube, it was able to ship salt as far as Bohemia, as well as into Austria and Bavaria. The Archbishop gradually bought up shares in Hallein, and by the early 16th century he held them all. However, the crown of Bohemia passed into the Habsburg family, and from the early 1600s, the great market of Bohemia was closed to the Archbishop. The other Austrian salt works were small at first. In the Salzkammergut, salt springs emerged from horizontal tunnels in the valley sides, which, although the locals did not know it, were the ancient galleries into the old flooded salt mines that had been worked in prehistoric times. The salt works at Hall, in the Tyrol, provided a power base for its owners, who were the local Hapsburg Dukes from 1363. The Dukes would sell salt to the Swiss, then use the profits to pay for the Hapsburg campaigns against the Swiss!

Salt production was always limited in Austria by shortages of fuel needed to extract salt by boiling brines. As the boiling houses consumed the local timber, they had to be moved, and fuel was a problem in salt manufacture in this region until modern times and the advent of highly mechanised mining operations. In 1770 there were purpose-built flumes running down the mountain sides, used not for water supply but to float down billets of timber for the boiling houses. Since fuel ran out at Hallstatt very early, the Emperor built a wooden pipeline to take the brine from the ancient mines down the valley to Ischland, on the way it crossed the Gosau Valley via a purpose-built bridge. Salt continued to play a significant role in the politics of the region after 1600,     when it was produced by three major players, Austria, Bavaria, and the Archbishop of Salzburg. The Austrian Empire grew to include Bohemia and Moravia, and this salt-less region became a captive market for the Austrian salt producers, with substantial tax revenue accruing to the Habsburg Emperor. Salt production was considered a state monopoly and Salzmonopol was considered "the brightest jewel in the possession of the Hofkammer.” By1700 it provided some 10% of the total revenue of the state.

In times of military emergency the Habsburgs would regularly use the salt income as collateral for raising money quickly. They did it first when Bohemia revolted in 1618 in the Defenestration of Prague, and Protestant forces besieged Vienna. Emperor Ferdinand II mortgaged his salt revenues to pay for the Catholic army that saved Vienna and won the decisive battle of the White Mountain in 1620. Salt revenue from the Wieliczka salt mine paid the Polish army under King John Sobieski when it rescued Vienna from the Turkish siege of 1683. Interestingly, the Wieliczka salt revenue had earlier passed to the Habsburgs in return for their assistance to the Poles in the Swedish invasion of 1657. Salt was also a state monopoly in Bavaria. Both Austria and Bavaria sought to promote their own salt exports and protect their domestic markets from salt imports, hence there was a flourishing trade in contraband salt.

In 1611 the Archbishop of Salzburg was forced to market his salt through Bavaria, so the rivalry now had only two players. Given that Austria and Bavaria between them controlled all the major salt sources in Central Europe, it is difficult to understand why they did not cooperate to form a cartel. A brief agreement, the Rosenheimer Salt Trade Agreement, was set up in 1649, but lasted for only 40 years. The centrepiece of Bavarian foreign policy became a campaign to sell salt effectively to her western neighbours, given that Austria could sell hers throughout the growing Austro-Hungarian Empire. It is not a coincidence that Bavaria consistently fought on the French side against the Austrians in the War of the Spanish Succession in the early 1700s and during the Napoleonic Wars in the early 1800s.

On the British mainland, Mary Queen of Scots was perhaps the first head of state to have the idea of making salt a taxable source of governmental revenue. She granted a patent to an Italian to make salt in Scotland and then placed a heavy tax on it, which she appropriated to herself. Elizabeth, Queen of England, and Mary's lifelong “dear sister” and eventual executioner, thought this an excellent idea and likewise taxed English salt making. Salt tax was a source of great resentment to everyone, English and Scots alike, and smuggling grew to alarming proportions. In 1785, the Earl of Dundonald wrote that every year in England, 10,000 people were arrested for salt smuggling. During Queen Anne’s reign, the salt tax rose to £30 a ton, an enormous amount of money in those days. The whole of England arose in rioting protest, with the result that the salt tax was finally abolished.

In Burgundy in the 1700s, salt was taxed at more than 100% as it came from the salt-works. This tax was extended to the whole of France when Burgundy was absorbed and the notorious salt tax “la gabelle” became a necessary input to the government's finances. Cardinal Richelieu said that salt was as vital to France as American silver was to Spain. The repeal of the salt tax was a major goal of the revolutionaries of 1789. A few years later, as soon as he became Emperor, Napoleon restored the salt tax to pay for his foreign wars. The salt tax continued until 1945 to feed French government coffers.

It is said that income from a salt pan in southern Spain largely financed Columbus’ voyages. The Erie Canal, an engineering marvel that connected the Great Lakes to New York's Hudson River in 1825, was called “the ditch that salt built” as salt taxes paid for half of the cost of construction. The “Great Hedge of India,” the mid-18th century colonial equivalent of the Great Wall of China, stretched 3,700 km from the western border of Punjab down to the Bay of Bengal. It was manned by 12,000 men and planted by the British to minimise salt smuggling into Bengal and so enforce the collection of the Indian salt tax. As late as the 1940s the people of India under the leadership of Mahatma Gandhi protested British taxes on salt supply. In 1930 Gandhi led a 200-mile march to the Arabian Ocean to symbolically collect untaxed salt for India's poor.

Artisanal salt and culinary expectation

Today, halite is a cheaply-produced commodity extracted from the subsurface in mines, or salt solution plants, or produced at the surface in saltpans. In the production of table salt, processing, packaging and marketing are the major costs for most salt manufacturers. An interesting exception to the low sale price of modern table salt is the artisinal “Fleur de Sel de Geurande” a delicate gourmet form of white seasalt that is still hand-produced on fens along the coast of Brittany (Figure 9). It costs ≈US$40/kg and is produced by “paludiers” only on suitable summer days when halite rafts can be raked from the brine surface of specially maintained coastal salt pans, which are floored with grey clay. According to the local legend, salt flowers only form on hot days when the wind blows from the east (from the sea). It and the cheaper grey salt (sel gris), which is scraped from the pan floors and also prized by gourmands, has been produced this way in French coastal fens since Pre-Roman times. The flowers of salt is marketed as a “natural” product that contains all 84 trace elements and micronutrients found in the sea, and as being a natural source of potassium, calcium, copper, zinc and magnesium.

 

This halite product has an intense white color, with rigid crunchy crystalline structure and high moisture content giving it a distinct “feel on the tongue.” This is because “Fleur de Sel“ is composed of clusters of halite rafts. These rafts formed on the brine surface, as a thin layer of floating salt crystals, which are harvested daily via raking and then placed on plastic sheets to dry in the sun, making it a highly labor intensive product (Figure 9). The flower of salt product is packed with no other processing, unlike what happens to industrially-produced sea salt that undergoes a process that typically consists of varying combinations of washing, centrifugation and drying by the heat of combustion, grinding and sieving. While large saltwork companies need several square kilometres for salt pan installations, a flower of salt product can be obtained in ponds with total areas smaller than 0.1 hectare. There is a definite economic upside to the artisanal production of “flowers of salt.” Since it is a handmade product, small salterns can be constructed/operated by family groups, so offering a new or supplemental income source for low-income populations living in or near hypersaline strand line areas.

Impurities like clay are called grey spot or black spot in highly efficient mechanised salt production plants across the world and are considered undesirable in the processed end-product. To the cynical it says something about French marketing skills, and perhaps the gullibility of middle-class gourmands with too much money and time on their hands, that each year the gourmet industry successfully markets un-processed dirt-polluted salt (sel-gris) scraped from pan floors for top prices. We shall look deeper into the geological characterisitics of various gourmet salt styles in a later essay.

The various untreated salt products from France, the Himalayas and elsewhere are typically marketed as a "natural organic" product, "completed untreated" so it retains all its "essential nutrients." Such blanket claims from marketers targeting a moneyed, health-conscious and "new-age" mostly middle-class demographic, should at times be taken with a grain of salt. For example, some types of Himalayan "natural" salt produced from high altitude continental lakes in Tibet is iodine deficient. Its local use has led to high levels of cretinism and other thyroid problems in the local peasant population. Introduction of a "processed" iodised salt by the Chinese authorities is still met with resistance, yet such use of iodised salt in China has reduced goiter to 10% of previous levels in the Chinese population. For the similar reasons, "back to nature" and "organic" foods are increasingly popular in middle-class consumers of Australia. The associated resistance to the use of "iodised" salt and other processed products with iodine additives by urban new-age parents has lead to unhealthy levels of iodine depletion in preschool-age urban children, when tested in Melbourne and Sydney (Li et al., 2001; McDonnell et al., 2003). Likewise the use of natural "untreated" salt from Lake Magadi and Natron as a food additive has led to significant health problems (fluorosis) in the local population due to "naturally" high levels of fluorine in the harvested salt (Vuhahula et al., 2009).

Salt, social standing, and religious superstition

Salt, because of its high value in the ancient world, has maintained both cultural and religious significance over more than three millennia. For example, in Medieval and Renaissance European kingdoms, easy access to salt during meals assigned social status. Intricately carved salt cellars would be placed on select tables within easy reach of those deemed worthy. Accordingly at any noble table, to be seated “below the salt” was to be seen as unworthy of access to such luxury (Figure 10).

 

From its value in its use as a preservative and food additive in the ancient pre-rationalist world, salt became a religious symbol, representing immutability and incorruptible purity. In many religions, salt is still included on the altar to represent purity, and it is mixed into holy water of various sects for the same reason. Ancient Greek worshippers consecrated salt in their rituals, for example the Vestal Virgins sprinkled all sacrificial animals with salt and flour. Salt was a token of permanence to both Jews of the Old Testament and Christians of the New Testament. To the Jews it came to signify the eternal covenant between Jews and Israel. Jewish temple offerings still include salt on the Sabbath and orthodox Jews still dip their bread in salt as a remembrance of those sacrifices. Covenants in both the Old and New Testaments were often sealed with salt, explaining the origin of the word “salvation.” In the Catholic Church, salt is used in a variety of purifying rituals. Jesus called his disciples “the Salt of the Earth”, a statement that was commemorated by the Catholic church until Vatican II, by placing a small taste of salt on a baby's lip at his or her baptism. 

So to the religious, salt is supernatural symbol of the permanent sanctity of Jesus and offers supposed protection to the superstitious. For example, salt is still used to make holy water and also the more powerful exorcised water of the Roman Catholic Church. Salt is also used to make protective circles during exorcisms of demons. In the middle of the last millennia in Europe, salt was believed to provide defence against witches, witchcraft, demons, sprites, and the evil eye. It was a common belief that witches, and the animals they bewitched, could not eat anything salted. Inquisitors were advised by demonologists to protect themselves by wearing an amulet of salt, consecrated on Palm Sunday, along with other blessed herbs, pressed into a disk of blessed wax. Carrying a concealed packet of salt was said to ward off the evil eye as well. Another known talisman to ward off evil spirits was a jar of salt and a knife. Some people put salt and pepper in their left boot for good fortune. To ward off an evil witch, a peasant might throw salt outside the front door and lean a broom next to it. A passing witch would have to count the grains of salt and the blades of straw on the broom before she could do any harm.

Similarly, any waste of salt can be a portent of evil. In Leonardo Da Vinci's famous painting, “The Last Supper,” Judas Escariot has just spilled a bowl of salt - a portent of evil and bad luck. In Buddhist tradition, salt repels evil spirits. It is also why in many Asian cultures it's customary to throw salt over your shoulder before entering your house after a funeral: it scares off any evil spirits that may be clinging to your back. In the Christian tradition you should throw spilt salt over you left shoulder as, according to the Medieval Church, the devil or his demons reside behind or on your left shoulder, with your guardian angels on the right. In Hawaii and Samoa sea salt is used for protection, both by placing salt in each of the four corners of the house and by poring salt on the door threshold to prevent any spirits from crossing into one’s home. Shinto religion also uses salt to purify an area. Before sumo wrestlers enter the ring for a match - which is actually an elaborate Shinto rite - a handful of salt is thrown into the centre to drive off malevolent spirits (Figure 11). In the American Southwest, the Pueblo worship the Salt Mother. Other native American tribes had significant restrictions on who was permitted to eat salt. Hopi legend holds that the angry Warrior Twins punished mankind by placing valuable salt deposits far from civilization, requiring hard work and bravery to harvest the precious mineral.

 

Chinese folklore credits the Phoenix with the discovery of salt. In Norse mythology the gods first came from a salty ice-block over the course of four days as the sacred cow, Auðumbla brought Búri the first god in Norse mythology, and grandfather of Odin, out of the salty ice block. In another creation myth, Tiamat is the symbol of the chaos of primordial creation in Mesopotamian religion (Sumerian, Assyrian, Akkadian and Babylonian). She is a primordial salty goddess of the ocean, mating with Apsu (the god of fresh water) to produce the younger gods. Her husband, Apsu, later makes war upon their children and is killed. When she, too, wars upon her husband's murderers, she is then slain by Enki's son, the storm-god Marduk. and the arch of the heavens and the earth were formed from her divided body. Records from the Middle Euphrates Hittite kingdom of Mari attest to the veneration of Hatta, the god of salt, through the erection of a statue to him by the city’s ruler, Zimri-Lim (Stackert, 2010). Among Hittite rituals, perhaps the best-known use of salt is one that parallels its use in various Mesopotamian curses: the First Hittite Soldier’s Oath employs salt within an analogical curse ritual against that soldier who would commit sedition. Ancient Greek worshippers also consecrated salt in their rituals.

 

Outcrops of diapiric salt masses can also have superstitious significance (Genesis 19:26); Lot’s wife was noted in the journals of Fulcher of Chartres (Chaplain to King Baldwin) who accompanied the crusader Baldwin I across the Dead Sea valley in December 1100 AD. In reality, the apophenic feature described as Lot’s wife is a 12m-high column of diapiric salt lying at the foot of the much larger Mt Sedom (Usdum) on the edge of the Dead Sea (Figure 12). It is one of a number of dissolutional remnants along the gypsum-capped cavernous edge of an outcropping diapir composed of Miocene salt, which makes up to core of Mount Sedom.

References

Aufderheide, A. C., 2011, The Scientific Study of Mummies, Cambridge University Press, 634 p.

Hay, R. L., and T. K. Kyser, 2001, Chemical sedimentology and paleoenvironmental history of Lake Olduvai, a Pliocene lake in northern Tanzania: Geological Society of America Bulletin, v. 113, p. 1510-1521.

Li, M., G. Ma, K. Guttikonda, S. Boyages, and C. Eastman, 2001, The re-emergence of iodine deficiency in Sydney, Australia: Asia Pacific J of Clin. Nutr., v. 10, p. 200-203.

Lonn, E., 1933, Salt as a Factor in the Confederacy: New York, Walter Neale, 324 p.

McDonnell, C., M. Harris, and M. Zacharin, 2003, Iodine Deficiency and goitre in schoolchildren in Melbourne, 2001: Med. J. Aust., v. 178, p. 159-162.

Pollard, A. M., D. R. Brothwell, A. Aali, S. Buckley, H. Fazeli, M. H. Dehkordi, T. Holden, A. K. G. Jones, J. J. Shokouhi, R. Vatandoust, and A. S. Wilson, 2008, Below the salt: A Preliminary study of the dating and biology of five salt-preserved bodies from Zanjun Province, Iran: Iran, v. 46, p. 135-150.

Stackert, J., 2010, The variety of ritual application for salt and the Maqlu salt incantation, in T. Abusch, and D. Schwemer, eds., Corpus of Mesopotamian Anti-witchcraft Rituals: Volume One, Brill, p. 235-252.

Vuhahula, E. A. M., J. R. P. Masalu, L. Mabelya, and W. B. C. Wandwi, 2009, Dental fluorosis in Tanzania Great Rift Valley in relation to fluoride levels in water and in "Magadi" (Trona): Desalination, v. 248, p. 610-615.

Danakil potash: K2SO4 across the Neogene: Implications for Danakhil potash, Part 4 of 4

John Warren - Tuesday, May 12, 2015

How to deal with K2SO4

In this the fourth blog focusing on Danakhil potash, we look at the potash geology of formerly mined Neogene deposits in Sicily and the Ukraine, then compare them and relevant processing techniques used to exploit their K2SO4 ore feeds. This information is then used to help guide a discussion of processing implications for potash extraction in the Danakhil, where kainite is the dominant widespread potash salt. As seen in the previous three blogs there are other potash mineral styles present in the Danakhil, which constitute more restricted ore fairways than the widespread bedded kainaite, these other potash styles (deep meteoiric -blog 2 of 4 and hydrothermal - blog 3 of 4), could be processed to extract MOP, but these other potash styles are also tied to high levels of MgCl2, which must be dealt with in the brine processing stream. The most effective development combination is to understand the three occurence styles , define appropriate specific brine processing strams and then combine the products in an single processing plant and then produce sulphate of potash (SOP), rather the Muriate of Potash (MOP), as SOP has a 30% price premium in current potash markets.

Kainite dominated the bedded potash ore feed in former mines in the Late Miocene (Messinian) sequence in Sicily and the Middle Miocene (Badenian) sequence in the Carpathian foredeep], Ukraine. Kainite also occurs in a number of potash deposits in the Permian of Germany and Russia. In Germany a combination of mined sylvite and kieserite is used to manufacture sulphate of potash (SOP). Interestingly, Neogene and the Permian are times when world ocean waters were enriched in MgSO4 (Lowenstein et al., 2001, 2003). In contrast, much of the Phanerozoic was typified by an ocean where MgSO4 levels were less. It is from such marine brine feeds that most of the world’s larger Phaneorzoic (SOP) potash ore deposits were precipitated (Warren, 2015). SOP is also produced from Quaternary Lake brines in China and Canada (see cryogenic salt blog; 24 Feb. 2015).

SOP in Messinian evaporites, Sicily

A number of potash mines on the island extracted kainitite from the late Miocene Solofifera Series of Sicily (Figure 1). The last of these mines closed in the mid-1990s, but portions of some are maintained and are still accessible (eg Realmonte mine). The halite-hosted potash deposits are isolated ore bodies within two generally parallel troughs, 115 km long and 5- 10 km wide, within the Caltanissetta Basin (Figure 1). They are separated by a thrust-related high 11-25 km wide and capped by the limestones of the “Calcare di Base”. Kainite is the dominant potash mineral in the mined deposits. Across the basin, ore levels constitute six layers of variable thickness, with a grade of 10%-16% K2O (pure kainite contains 18.9% K20), with very little insoluble content (0.4%-2.0%).

At the time the potash was deposited there was considerable tectonic activity in the area (Roveri et al. 2008, Manzi et al., 2011). Host sediments were deposited in piggy-back basins some 5.5 Ma atop a series of regional thrusts, so the ore layers have dips in the mines ranging up to 60° (Figure 2). Little if any of the limestone associated with the deposits was converted to dolomite, nor was the thick Messinian gypsum (upper and lower units), encasing the halite /kainitite units, converted to anhydrite, it remains as gypsum with well preserved depositional textures. However, the elevated salinities, and perhaps temperatures, required for kainite precipitation means anhydrite micronodules, observed in some ore levels, may be primary or syndepositional. A lack of carnallite, along with isotopic data, indicates that when the deposits were formed by the evaporation of the seawater, salinities did not usually proceed far past the kainite crystallization point (in contrast to Ethiopia where carnallite salinities typify the later stages of kainitite deposition)..

 

The largest Sicilian ore body was at Pasquasia, to the west of Calanisseta, covering a 24 km2 area at depths of 300-800 m (Figure 1). There were five ore beds at Pasquasia, all with highly undulating synclinal and anticlinal forms. The Number 2 bed was the thickest, averaging perhaps a 30-m thickness of 10.5% to 13.5% K2O ore. The Pasquasia Mine was last operational from 1952 to 1992.

 

Ore geology remains somewhat more accessible at the former Realmonte mine, near the town of Agrigento. There, four main depositional units (A to D from base to top) typify the evaporite geology. As at Pasquasia, kainitite was the targeted ore within a Messinian evaporite section that has total thickness of 400-600 m. As defined by Decima and Wezel, 1971, 1973; Decima, 1988, Lugli, 1999, the Realmonte mine section is made up of 4 units (Figure 2a):

- Unit A (up to 50 m thick): composed of evenly laminated grey halite with white anhydrite nodules and laminae that pass upward to grey massive halite beds.

- Unit B (total thickness ≈100 m): this potash entraining interval is dominated by massive even layers of grey halite, interbedded with light grey thin kainite laminae and minor grey centimetre-scale polyhalite spherules and laminae, along with anhydrite laminae; the upper part of the unit contains at least six light grey kainite layers up to 18 m-thick that were the targeted ore sequence. Unlike the Danakil, carnallite does not typify the upper part of this marine potash section. The targeted beds are in the low-angle dip portion of a thrust-folded remnant in a structural basin (Figures 2b, 3).

- Unit C (70-80 m thick): is made up of white halite layers 10-20 cm thick, separated by irregular dark grey mud laminae and minor light grey polyhalite and anhydrite laminae (Figure 3).

- Unit D (60 m thick): is composed of a grey anhydritic mudstone (15-20 m thick), passing up into an anhydrite laminite sequence, followed by grey halite millimetre to centimetre layers intercalated with white anhydrite laminae.


According to Lugli, 1999, units A and B are made up of cumulates of well-sorted halite plate crystals, up to a few millimeters in size. Kainite typically forms discrete laminae and sutured crystal mosaic beds, ranging from a thickness of few mm to a maximum of 2 m, intercalated and embedded within unit B (Garcia-Veigas et al., 1995). It may also occur as small isometric crystals scattered within halite mosaics. Kainite textures are dominated by packed equant-granular mosaics, which show possible pressure-dissolution features at some grain boundaries. The associated halite layers are dominantly cumulates, which show no evidence of bottom overgrowth chevrons, implying evaporite precipitation was a “rain from heaven” pelagic style that took place in a stratified permanently subaqueous brine water body, possibly with a significant water depth to the bottom of the permanent lower water mass.

Only the uppermost part of potash bearing portion of unit B shows a progressive appearance of large halite rafts along with localized dissolution pits filled by mud, suggesting an upward shallowing of the basin at that time. In many parts of the Realmonte mine spectacular vertical fissures cut through the topmost part of unit B at the boundary with unit C, suggesting desiccation and subaerial exposure at this level (Lugli et al., 1999).

The overlying unit C is composed of cumulates of halite skeletal hoppers that evolve into halite chevrons illustrating bottom growth after foundering of the initial halite rafts. Halite layers in unit C show numerous dissolution pits filled by mud and irregular truncation of the upper crystal terminations, implying precipitation from a nonstratified, relatively shallow water body. Palaeo-temperatures of the brine that precipitated these halite crystals are highly variable from 22 to 32°C (Lugli and Lowenstein, 1997) and suggest a shallow hydrologically unstable body of water, unlike units A and B.

The bromine content of halite increases from the base of unit A to the horizons containing kainite (layer B) where it obtains values of up to 150 ppm. Upwards, the bromine content decreases once more to where at the top of Unit C it drops below 13 ppm, likely indicating a marked dilution of the mother brine. The dilution is likely a consequence of recycling (dissolution and reprecipitation) of previously deposited halite either by meteoric-continental waters (based on Br content; Decima 1978), or by seawater (based on the high sulphate concentration and significant potassium and magnesium content of fluid inclusions; Garcia-Veigas et al., 1995).

As in the Danakhil succession, evaporite precipitation at Realmonte began as halite-CaSO4 interlayered succession at the bottom of a stratified perennial water body, which shallowed and increased in concentration until reaching potash kainite saturation. In Sicily, this was followed by a period of exposure and desiccation indicated by the presence of giant megapolygonal structures. Finally, seawater flooded the salt pan again, dissolving and truncating part of the previous halite layers, which was then redeposited under shallow-water conditions at the bottom of a nonstratified (holomitic) water body (Lugli et al., 1997, 1999).

Unlike Ethiopia, the Neogene kainite deposits of Sicily were deposited in a thrust “piggy-back” basin setting and not in a rift sump (Figure 2b). Mineralogically similar, very thick, rift-related, now halokinetic, halite deposits of Midddle Miocene age occur under the Red Sea’s coastal plain between Jizan, Saudi Arabia (where they outcrop) to Safaga, Egypt, with limited potash is found in some Red Sea locations at depths suitable for solution mining (Notholt 1983; Garrett, 1995). Potash-enriched marine end-liquor brines characterise Red Sea geothermal springs, implying a more sizeable potash mass may be (or once have been) present in this region. Hite and Wassef (1983) argue gamma ray peaks in two drill hole logs in this area suggest the presence of sylvite, carnallite and possibly langbeinite at depth.

K2SO4 salts in Miocene of Ukraine

Miocene salt deposits occur in the western Ukraine within two structural terranes: 1) Carpathian Foredeep (rock and potash salt) and (II) Transcarpathian trough (rock salt) (Figure 4a). These salt-bearing deposits differ in the thickness and lithology depending on the regional tectonic location (Czapowski et al., 2009). In the Ukrainian part of Carpathian Foredeep, three main tectonic zones were distinguished (Figure 4b): (I) outer zone (Bilche-Volytsya Unit), in which the Miocene molasse deposits overlie discordantly the Mesozoic platform basement at the depth of 10-200 m, and in the foredeep they subsided under the overthrust of the Sambir zone and are at depths of 1.2-2.2 km (Bukowski and Czapowski, 2009); Hryniv et al., 2007); (II) central zone (Sambir Unit), in which the Miocene deposits were overthrust some 8-12 km onto the external part of the Foredeep deposits of the external zone occur at depths of 1.0-2.2 km; (III) internal zone (Boryslav-Pokuttya Unit), where Miocene deposits were overthrust atop the Sambir Nappe zone across a distance of some 25 km (Hryniv et al., 2007).


The Carpathian Foredeep formed during the Early Miocene, located north of emerging the Outer (Flysch) Carpathians. This basin was filled with Miocene siliciclastic deposits (clays, claystones, sandstones and conglomerates) with a maximum thickness of 3 km in Poland and up to 5 km in Ukraine (Oszczypko, 2006). Two main evaporite bearing formations characterise the saline portions of the succession and were precipitated when the hydrographic connection to the Miocene ocean was severely reduced or lost (Figures 4, 5): A) Vorotyshcha Beds, dated as Late Eggenburgian and Ottnangian, some 1.1-2.3 km thick and composed of clays with sandstones, with exploitable rocksalt and potash salt interbeds. This suite is further subdivided into two subsuites: a) A lower unit, some 100-900 m thick with rock salt beds and, b) An upper unit, some 0.7-1.0 km thick, with significant potash beds, now deformed (Hryniv et al., 2007).The Stebnyk potash mine is located in this lower subset in the Boryslav-Pokuttya Nappe region, close to the Carpathian overthrust); B) Tyras Beds of Badenian age reach thicknesses of 300-800 m in the Sambir and Bilche-Volytysa units and are dominated by salt breccias and contain both rock and potash salts. Thicknesses in the Bilche-Volytsya Unit range from 20-70 m and are made up of a combination of claystones, sandstones, carbonates, sulphates and rock salts with little or no potash.


Hence, potash salts of the Carpathian Foredeep are related either to the Vorotyshcha Beds located in the Boryslav–Pokuttya zone, or to the Tyras Beds (Badenian) in the Sambir zone (Figure 5). These associations range across different ages, but have many similar features, such as large number of potash lenses in the section, mostly in folded-thrust setting, and owing to their likely Neogene-marine mother brine contain many sulphate salts, along with a high clay content. Accordingly, the main potash ore salts are kainite, langbeinite and kainite–langbeinite mixtures. Hryniv et al. (2007) note more than 20 salt minerals in the Miocene potash levels and in their weathering products. Bromine contents in halites of the Carpathian Foredeep for deposits without potash salts range from 10 to 100 ppm (on average 56 ppm); in halite from salt breccias with potash salts range from 30 to 230 ppm (average 120 ppm); and in halite from potash beds ranges from 70 to 300 ppm (average 170 ppm). In the ore minerals from the main potash deposits, bromine content ranges are: a) in kainite 800–2300 ppm; b) in sylvite 1410–2660 ppm; and c) in carnallite 1520–2450 ppm. This is consistent with kainite being a somewhat less saline precipitate than carnallite/sylvite (Figure 6).


The brines of Vorotyshcha and Tyras salt-forming basins (based on data from brine inclusions in an investigation of sedimentary halite, listed by Hyrniv et al. (2007), are consistent with mother brines of the Na–K–Mg–Cl–SO4 (MgSO4-rich) chemical type (consistent with a Neogene marine source). Inclusion analysis indicates the temperature of halite formation in the Miocene basin brines in Forecarpathian region was around 25°C. During the potash (Kainite) stages it is likely these solutions became perennially stratified and heliothermal so that the bottom brines could be heated to 40-60°C, more than double the temperature of the brine surface layer (see Warren, 2015 for a discussion of the physical chemistry and the various brine stratification styles). During later burial and catagenesis the temperatures preserved in recrystallised halites are as high as 70°C with a clear regional tectonic distribution (Hryniv et al. (2007).

Maximum potash salt production was achieved under Soviet supervision in the 1960s, when the Stebnyk and Kalush mines delivered 150 x 106 tonnes of K2O and the “New” Stebnyk salt-works some 250 x 106 tonnes as K2SO4 per year.


Stebnyk potash (Figure 7a)

The potash salt deposit in the Stebnyk ore field occurs within the Miocene (Eggenburgian) Vorotyshcha Beds (Figures 4, 5). Salt-bearing deposits in the Stebnyk area were traditionally attributed to two main rock complexes (Lower and Upper Vorotyshcha Beds) separated by terrigenous (sandstones and conglomerates) Zahirsk Beds (Petryczenko et al., 1994). More recent work indicates that the Zahirsk Beds belonged to a olistostrome horizon (a submarine slump, interrupting evaporite deposition) and there are no valid arguments for subdividing the Vorotyshcha Beds into two subunits (Hryniv et al., 2007).

There are multiple salt-bearing series in the Stebnyk deposit (Figure 4b) and their total thickness ranges up to 2,000 m in responses to intensive fold thickening and overthrusting of the Carpathians foredeep. Intervals with more fluid salt mineralogies were compressed and squeezed into the centers of synclinal folds, to form a number of elongate lens-shape ore bodies (Figure 4b). These bodies are often several hundreds meters wide and in mineable zones occur at the depth of 80-650 m, typically at 100-360 m.

The lower part of the Vorotyshcha Suite (Beds) in the Stebnyk Mine area is composed of a salt-bearing breccia, with sylvinite or carnallitite interclayers typically in its upper parts, as well as numerous blocks of folded marly clays (Bukowski and Czapowski, 2009). Above this is the potash-bearing ore series , some 10-125 m thick and, composed of beds of kainite, langbeinite and lagbeinite-kainite with local sylvinite and kieserite (Hryniv et al., 2007). The potash interval is overlain by a rock salt complex some 60 m thick (Koriń, 1994).

The Stebnyk plant is now abandoned and in disrepair. In 1983 there was a major environmental disaster (explosion) at a nearby chemical plant (in the ammonia manufacture section), which was supplied chemical feedstock by the mine. No lives were lost, but damage at the plant, tied to the explosion, released some 4.6 million cubic metres of thick brine from an earthen storage dam into the nearby Dniester River. At the time this river was probably the least environmentally damaged by industrial operations under Soviet administration. The spill disrupted water supplies to millions of people along the river, killed hundreds of tons of fish, destroyed river vegetation and deposited a million tons of mineral salts on the bottom of a 30-mile-long reservoir on the Dniester. Stebnik is located in the Ukrainian province of Lvov. Staff members at the United States Embassy at the time seized on the name to dub the incident ‘’Lvov Canal,’’ after the Love Canal contamination in the United States.

Kalush potash salt geology (Figure 7b)

Thickness of Miocene (Badenian) deposits near the Kalush Mine is around 1 km (Figures 4a). Two local salt units (beds) are distinguished within the Tyras Beds: the Kalush and Holyn suites, which constitute the nucleus of Miocene deposits of Sambir Unit (Figure 5). Beds have been overthrust and folded onto the Mesozoic and Middle to Upper Miocene molasse sediments of the outer (Bilche-Volytsya) tectonic unit (Figure 4b). The Kalush Beds are 50-170 m thick, mostly clays, with sandstone and mudstone intercalations,. In contrast the Holyn beds are more saline and dominated by clayey rock salts (30-60% of clay), salty clays and claystones (Koriń, 1994). Repeated interbeds and concentrations of potash salts up to several meters thick within the Holyn beds define a number of separate potash salt fields in the Kalush area (Figures 4b, 5). Such salt seams are dominated by several MgSO4-enriched mineralogies: kainite, langbeinite-kainite, langbeinite, sylvinite and less much uncommon carnallite and polyhalite. These polymineralogic sulphate ore mineral assemblages are co-associated with anhydrite, kieserite and various carbonates. The potash ore fields typically occur in tectonic troughs within larger synclines, usually at depths of 100-150 m, to a maximum of 800 m.

Conventional processing streams for manufacture of SOP and MOP

To date the main natural sulphate salts that have been successfully processed to manufacture sulphate of potash (SOP) are;

  • Kainite (KCl.MgSO4.3H2O) (as in Sicily - potash mines are no longer active)
  • Kieserite (MgSO4.H2O) (as in Zechstein, Germany - some potash mines active)
  • Langbeinite (K2SO4.2MgSO4) (as in Carlsbad, New Mexico - active potash mine)
  • Polymineralic sulphate ores (as in the Stebnyk and Kalush ores, Ukraine - these potash mines are no longer active)
  • All the processing approaches deal with a mixed sulphate salt or complex sulphate brine feed and involve conversion to form an intermediate doublesalt product, usually schoenite (or leonite at elevated temperatures) or glaserite. This intermediate is then water-leached to obtain SOP.

    For example, with a kainite feed, the process involves the following reactions:

    2KCl.MgSO4.3H2O --> K2SO4.MgSO4.6H2O + MgCl2

    followed by water-leaching of the schoenite intermediate

    K2SO4.MgSO4.6H2O --> K2SO4 + MgSO4 + 6H2O


    In Sicily in the 1960s and 70s, the Italian miners utilized such a solid kainitite ore feed, from conventional underground mining and leaching approaches. The various Italian mines were heavily government subsidized and in terms of a free-standing operation most were never truly profitable. The main kainitite processing technique used in Sicily, is similar in many ways to that used to create SOP from winter-precipitated cryogenic salt slurries in pans that were purpose-constructed in the North Arm area of in Great Salt Lake, Utah (Table 1; see Warren, 2015 for details on Great Salt Lake operations). The Italian extraction method required crushing and flotation to create a fine-sized kainite ore feed with less than 5% NaCl. This product was then leached at temperatures greater than 90°C with an epsomite brine and converted into a langbeinite slurry, a portion which was then reacted with a schoenite brine to precipitate potassium chloride and epsomite solids, which were then separated from each other and from the epsomite brine. A portion of the potassium chloride was then reacted with magnesium sulphate in the presence of a sulphate brine to create schoenite and a schoenite brine. This schoenite brine was recycled and the remaining potassium chloride reacted with the schoenite in the presence of water, to obtain potassium sulphate and a sulphate brine.

    The processing stream in the Ukraine was similar for the various Carpathian ore feeds, which “out-of-mine-face” typically contained around 9% potassium and 15% clay and so were a less pure input to the processing stream, compared to the typical mine face product in Sicily. Like Sicily, schoenite was the main intermediate salt. Ore was leached with a hot synthetic kainite solution in a dissolution chamber. The langbeinite, polyhalite and halite remained undissolved in the chamber. Salts and clay were then moved into a Dorr-Oliver settler where the clays were allowed to settle and were then moved to a washer and discarded. The remaining solution was crystallized at the proper cation and anion proportions to produce crystalline schoenite. To avoid precipitation of potassium chloride and sodium chloride, a saturated solution of potassium and magnesium sulfate was added to the Dorr-Oliver settler. The resulting slurry of schoenite was filtered and crystals were leached with water to produce K2SO4 crystals, which were centrifuged and recycled and a liquor of potassium and magnesium sulfates obtained. The liquid phase from the filter was recycled and added to the schoenite liquor from obtaoned by vacuum crystallization. Part of the schoenite liquor was evaporated to produce crystalline sodium sulfate, while the magnesium chloride liquid end product was discarded. The slurry from the evaporation unit was recycled as “synthetic kainite.” This process stream permitted the use of the relatively low quality Carpathian ore and produced several commercially valuable products including potassium sulfate, potassium-magnesium sulfate, potassium chloride, sodium sulfate and magnesium chloride liquors. Being a Soviet era production site, the economics of the processing was not necessarily the main consideration. Rather, it was the agricultural utility of the product that was paramount to the Soviet state.

    Can Danakhil potash be economically mined?

    For any potash deposit (MOP or SOP) there are three approaches that are used today to economically extract ore (Warren 2015): 1) Conventional underground mining. 2) Processing of lake brines 3) Solution mining and surface processing of brines. Historically, method 1 and 2 have been successfully conducted in the Danakhil Depression, although method 1) was terminated in the Dallol area by a mine flood.

    Conventional mining

    To achieve a successful conventional underground MOP potash mine any where in the world, ideally requires (Warren, 2015): 1) A low dipping, laterally continuous and consistently predictable quality ore target, not subject to substantial changes in bed dip or continuity. 2) An ore grade of 14% K2O or higher, and bed thickness of more than 1.2 m. 3) Around 8-m of impervious salt in the mine back or roof, although some potash mines, such as the Boulby mine in the UK are working with < 2 meters of salt in the back (but there the extraction is automated and the access roads approach the target ore zone from below). 4) An initial access shaft that is vertical and typically dug using ground freezing techniques to prevent unwanted water entry during excavation. 5) A typical ore depth in the range 500-1100 metres. Shallower mines are subject to unpredictable water entry/flooding and catastrophic roof collapse, as in the Cis-Urals region (see Solikamsk blog). Mines deeper than 1000-1100 metres are at the limit of conventional mining and the salt surround is subject to substantial creep and possible explosive pressure release outbursts (as in some potash mines in the former East Germany). 6) At-surface and in-mine conditions not subject to damage by earthquakes, water floods or volcanism.

    During the feasibilty phase of the Parsons Mining Project it became evident that the halite material overlying the Sylvinite Member was porous and that there was no adequate hydrologic protection layer above the Sylvinite Member. In my mind, this is further evidence of the hydrologic access needed to convert carnallite to sylvite along the bajada front (see previous blog). In any event the absence of a hydrologic protection layer above the Sylvinite Member means that conventional underground mining is not feasible for this type of potash. In addition, given the tectonic instability of the Danakhil Depression it is likely that no underground conventional mine is feasible in the hydrologically, seismically and hydrothermally active setting, which is the Danakhil depression, even if planning to exploit the deeper widespread kainitite beds (>350-450m)

    Some explorers in the Danakhil depression, especially on the Eritrean side are proposing to use surface or open-pit mining (quarrying) approaches to reach and extract/processing shallow ore salts. For this approach to be successful requires the shallow potash targets to be above regional groundwater level. Depths to the different ore targets on the Ethiopian side of the depression range between 45m and 600m and almost all lie below the regional water. Also, to access the mineralised material a large volume of variably water-saturated overburden would need to be removed. Even if areas with ore levels above the water table do exist on the Ethiopian side, the whole of the Danakhil sump is subject to periodic runoff and sheetflooding, sourced in the western highlands. Open pit areas would be regularly flooded during the lifetime of the pit, resulting in a need for extensive dewatering. For these reasons, and the possibility of earthquake damage, open pit mining is likely not feasible.

    Can the Danakhil potash be solution mined?

    To achieve this, brines extracted from different mineralogical levels and ore types will need to be individually targeted and kept as separate feeds into dedicated at-surface processing streams. On the Dallol surface, there are numerous sites that are suitable for pan construction, the climate is suitable for natural solar concentration as the region is typically dry, flat and hyperarid. If the potash zones in the Dallol depression are to be economically exploited via solution mining it will likely first require an understanding of the geometries of the 3 different forms of potash, namely; 1) Bedded kainitite-carnallitite (widespread in the depression), 2) Diagenetic sylvite via incongruent dissolution (focused by deep meteoric mixing and the bajada chemical interface along the western margin. 3) Hydrothermal potash (largely found in the vicinity of Dallol mound). Next, in order to have known-chemistry feedstocks into a SOP chemical plant, it will require the appropriate application of extraction/solution mining chemistries for each of these deposit styles. This would involve the construction of dedicated brine fields and the pumping of shallow Dallol brines (mostly from <200-250m below the surface) into a series of mineralogically-separated at-surface solar concentrator pans. 

    There are some subsurface aspects that need to be considered and controlled  in a solution mining approach in the Danakhil. The first is the possibility of uncontrolled solution cavity stoping (for example where a solution cavity blanket layer is lost due to cavity intersection with an unexpected zone of high permeability). If cavity shape is not closely monitored (for example by regular downhole sonar scans) and controlled, this could ultimately lead to the collapse of the land surface atop regions of shallow evaporites (<150-200 below the surface). As we saw in blog 3, doline collapse is a natural process in the Dallol Mound region, as it is any region of shallow soluble evaporites in contact with undersaturated pore waters. Ongoing solution via interaction with hydrothermal waters has created the colorful brine springs that attract tourists to the Dallol Mound region. But a operator does not want new dolines to daylight in their brine field, as environmental advocates would quickly lay blame at the feet of the brinefield operator. For this reason, the region in the vicinity of the Dallol Mount (eg the “Crescent deposit”) should probably be avoided.

    Most modern brinefield operators prefer a slowly-dissolving targeted salt bed that is at least 400-500m below the land surface (Warren, 2015). This broadens and lessens the intensity of the cone of ground collapse above the extraction zone and so lessens the possibility of catastrophic surface collapse. Use of a diesel rather than air blanket during cavity operation is also preferred because of potential porosity intersections at the base of the Upper Rock Salt (URF) contact (see blog 2 in the Danakhil blogs) Appropriate deeper potash beds in the Danakhil are laterally continuous beds of kainitite with lesser carnallitite. Drilling to date has identified little sylvite or bischofite in these widespread layers. This simplifies the mineral input chemistry in terms of a kainite target further out in the saltflat with a sylvite or sylvite bischofite operation closer toward the western margin, but there are no currently active solution mines solely targeting a kainite ore anywhere in the world.

    This leads to another consideration with a solution mining approach in the Danakhil depression, and that is that there are no existing brine technologies that can deal economically with high concurrent levels of magnesium and possibly-elevated sulphate levels in a recovered brine feed. The third consideration is reliably predicting the occurrence of, and avoiding, any metre- to decametre-scale brine-filled cavities that the drilling has shown are not uncommon at the sylvinite-bischofite-carnallite level in the Dallol stratigraphy along the Bajada chemistry zone. Intersecting and slowly dewatering such large brine cavities may not lead to at-surface ground collapse, but if not identified could create unexpected variations in the ionic proportions of brine feeds into the solar concentrators (for example drilling has identified subsurface regions dominated by bischofite, which is one of the most soluble bittern salts in the Danakhil depression - see Ercospan 2010, 2011 for drill result summaries).

    And so?

    So, at this stage, there are encouraging possibilities for economic recovery of both MOP and SOP from solution brines pumped to chemistry-specific solar pans in the Danakhil. Processing chemistry will require further site-specific studies to see which of the current known methods or their modification is economically feasible for SOP and perhaps combined SOP and MOP manufacture in the hyperarid climate of the Danakhil, as is being currently done by Allana Potash. It is also possible that a new processing stream chemistry could to be developed for the Dallol brines, in order to deal with very high concurrent levels of MgCl2 (widespread bischofite beds), or develop new or modify existing processing streams that target kainitite at depth. Similar K2SO4 brine processing chemistries have been applied in pans of the margins of the Great Salt Lake. But there salt pan processing was in part seasonally cryogenic, something that the Dallol climate certainly is not, so it is likely modified or new approaches to year-round pan management will be required.

    Any future potash operation in the Danakil will have to compete in product pricing with well established, high-volume low cost producers in Canada, Belarus and Russia (Figure 8). Today, establishing a new conventional underground potash mine is associated with setup costs well in excess of a billion dollars (US$). The costs are high as the entry shaft to a conventional underground mine must be completed without water entry and is usually done via ground freezing. This is the approach currently underway at BHP’s MOP Jansen Mine in Saskatchewan, Canada. Because of the very high costs involved in underground entry construction, and the well established nature of the competition, the proved amount of ore for a conventional mine should be sufficient for at least 20 years of production (subject to a given mill size, mill recovery rate for a given ore depth and the density and origin of salt “horses”). Kogel et al. (2006) states any potash plant or mill should be at capable of least 300,000 t K2O per annum in order to compete with a number of established plants with nameplate capacity in excess of 1 Mt.

    In contrast, the shallow nature of a Danakhil potash source means cheaper access costs, while a solution well approach makes for much cheaper and shorter approach times for brine/ore extraction, providing suitable economic brine processing streams are available (Figure 8). Potash is a mine product where transport to market is a very considerable cost proportion in terms of an operation's profitability. The location of the Danakhil gives it a low-cost transport advantage as a future supplier to the ever-growing agricultural markets of Africa, India and perhaps China. And finally, a potassium sulphate product has a 30% cost premium over a muriate of potash (KCl) product.

    References

    Bukowski, K., and G. Czapowski, 2009, Salt geology and mining traditions: Kalush and Stebnyk mines (Fore-Carpathian region, Ukraine): Geoturystyka, v. 3, p. 27-34.

    Czapowski, G., K. Bukowski, and K. Poborska-Młynarska, 2009, Miocene rock and potash salts of West Ukraine. y): Field geological-mining seminar of the Polish Salt Mining Society. Geologia (Przegląd Solny 2009), Wyd. AGH, Kraków, 35, 3: 479-490. (In Polish, English summary).

    Decima, A., J. A. McKenzie, and B. C. Schreiber, 1988, The origin of "evaporative" limestones: An Example from the Messinian of Sicily: Journal of Sedimentary Petrology, v. 58, p. 256-272.

    Decima, A., and F. Wezel, 1973, Late Miocene evaporites of the central Sicilian Basin; Italy: Initial reports of the Deep Sea Drilling Project, v. 13, p. 1234-1240.

    Decima, A., and F. C. Wezel, 1971, Osservazioni sulle evaporiti messiniane della Sicilia centromeridionale: Rivista Mineraria Siciliana, v. 130–132, p. 172–187.

    Garcia-Veigas, J., F. Orti, L. Rosell, C. Ayora, R. J. M., and S. Lugli, 1995, The Messinian salt of the Mediterranean: geochemical study of the salt from the central Sicily Basin and comparison with the Lorca Basin (Spain): Bulletin de la Societe Geologique de France, v. 166, p. 699-710.

    Garrett, D. E., 1995, Potash: Deposits, processing, properties and uses: Berlin, Springer, 752 p.

    Hite, R. J., and A. S. Wassef, 1983, Potential Potash Deposits in the Gulf of Suez, Egypt: Ann. Geol. Survey Egypt, v. 13, p. 39-54.

    Hryniv, S. P., B. V. Dolishniy, O. V. Khmelevska, A. V. Poberezhskyy, and S. V. Vovnyuk, 2007, Evaporites of Ukraine: a review: Geological Society, London, Special Publications, v. 285, p. 309-334.

    Koriń, S. S., 1994, Geological outline of Miocene salt-bearing formations of the Ukrainian fore-Carpathian area (In Polish, English summary): Przegląd Geologiczny, v. 42, p. 744-747.

    Lowenstein, T. K., L. A. Hardie, M. N. Timofeeff, and R. V. Demicco, 2003, Secular variation in seawater chemistry and the origin of calcium chloride basinal brines: Geology, v. 31, p. 857-860.

    Lowenstein, T. K., M. N. Timofeeff, S. T. Brennan, H. L. A., and R. V. Demicco, 2001, Oscillations in Phanerozoic seawater chemistry: Evidence from fluid inclusions: Science, v. 294, p. 1086-1088.

    Lugli, S., 1999, Geology of the Realmonte salt deposit, a desiccated Messinian Basin (Agrigento, Sicily): Memorie della Societá Geologica Italiana, v. 54, p. 75-81.

    Lugli, S., and T. K. Lowenstein, 1997, Paleotemperatures preserved in fluid inclusions in Messinian halite, Realmonte Mine (Agrigento, Italy): Neogene Mediterranean Paleoceanography, 28–30 September 1997, Erice. Abstract volume, 44–45.

    Lugli, S., B. C. Schreiber, and B. Triberti, 1999, Giant polygons in the Realmonte mine (Agrigento, Sicily): Evidence for the desiccation of a Messinian halite basin: Journal of Sedimentary Research Section A-Sedimentary Petrology & Processes, v. 69, p. 764-771.

    Manzi, V., S. Lugli, M. Roveri, B. C. Schreiber, and R. Gennari, 2011, The Messinian "Calcare di Base" (Sicily, Italy) revisited: Geological Society of America Bulletin, v. 123, p. 347-370.

    Notholt, A. J. G., 1983, Potash in Developing Countries, in R. M. McKercher, ed., Potash '83; Potash technology; mining, processing, maintenance, transportation, occupational health and safety, environment, p. 29-40.

    Oszczypko, N., P. Krzywiec, I. Popadyuk, and T. Peryt, 2006, Carpathian Foredeep Basin (Poland and Ukraine): Its Sedimentary, Structural, and Geodynamic Evolution, in J. Golonka, and F. J. Picha, eds., The Carpathians and their foreland: Geology and hydrocarbon resources, The American Association of Petroleum Geologists Memoir, v. 84, p. 293-350.

    Petryczenko, O. I., G. M. Panow, T. M. Peryt, B. I. Srebrodolski, A. W. Pobereżski, and K. W.M., 1994, Outline of geology of the Miocene evaporite formations of the Ukrainian part of the Carpathian Foredeep (In Polish, English summary): Przegląd Geologiczny, v. 42, p. 734-737.

    Roveri, M., S. Lugli, V. Manzi, and B. C. Schreiber, 2008, The Messinian Sicilian stratigraphy revisited: new insights for the Messinian salinity crisis: Terra Nova, v. 20, p. 483-488.

    Warren, J. K., 2015, Evaporites: A compendium (ISBN 978-3-319-13511-3) Released August 2015: Berlin, Springer, 1600 p.

    Danakhil Potash; Ethiopia - Modern hydrothermal and deep meteoric KCl, Part 3 of 4

    John Warren - Friday, May 01, 2015

    So far we have discussed the modern salt pan geology of the Danakhil (Part 1 of 4) and the initial subaqueous setting for widespread bedded potash, now in the subsurface, mostly as a kainitite bed (Part 2 of 4). In this blog we will discuss examples of potash in the Danakhil where remobilised salts and brines are related to the circulation of hydrothermal and meteoric fluids have facilitated localised reworking of potash to the surface (part 3 of 4). These fluids are related to the thermal anomalies created by the emplacement of the Dallol mound and the chemical front created by the encroachment of the Bajada along the western margin of the saltflat. Notably, we shall see the Dallol Mound is not a volcanic cone, rather it is an anticlinal dome of uplifted and eroded bedded salt, capped and surrounded by hydrothermal crater features typified by karst pools and brine outflows. Its creation is likely related to emplacement of igneous material at depth but, as yet there, has been no breakout of volcanic rock material in the mound area. This has important economic implications for the nature of remobilised potash and the creation of potential potash ores in the Dallol Mound area, these cosiderations are separate from the regional distribution of primary potash beds (kainitite and carnallitite) that were discussed in the previous blog.


    Thermal brine springs and potash occurrences near Dallol mound

    Today, hot springs supply and maintain a number of hydrothermally-fed brine pools and brine filled karst lakes in various depressions both atop and near the regional anticlinal salt mound or salt dome, sometimes called Dallol Mountain (Figure 1). As it only rises some 60 metres from the surrounding surface (-81 m versus -120 m) the term mountain is a misnomer. The highly dissected and eroded slope of bedded halite that is the southwest margin of Dallol mound shows the various springs are active in a region of uplifted and eroded bedded evaporite that defines the Dallol mound (Figure 2a). For example, brine springs still supply a small carnallite deposit known as the Crescent deposit located near the uplifted black halite beds that define Black Mountain and located 1.5 km southwest of Dallol mound (Figure 2b). This potash ore is the result of hydrothermally-driven groundwater activity, likely related to the emplacement of the Dallol Mound. The uplift-related thermal hydrology has broken up the mineralogical continuity of the nearsurface evaporite beds including the equivalents to the potash-rich Houston Fm.


    The Black Mountain potash deposits caught the attention of the Houston-based  Ralph M. Parsons company in 1954 where, according to Holwerda and Hutchinson, 1968, potash mining had previously already taken place at the Crescent carnallite/sylvite deposit. Earlier extraction had involved, amongst other techniques, flooding of salt pans around a continuously flowing hot spring, followed by harvesting of potash-rich salts, once natural deliquescence had flushed most of the highly soluble MgCl2 from the system. A concession was obtained Parsons linked to obligations to investigate the various potash deposits in the area, some of which were tied to actual outcrops of potash salts. The Parsons Company set up its base on Dallol Mountain at a site previously occupied by the Italian mining community, which had operated in the first few decades of last century (Figure 2a; the modification and reuse of older salt brick buildings is still evident on the ground today). As well, Parsons Co. constructed airstrips on Dallol Mountain and in the Musley area. They drilled more than 300 holes in order to better understand the the distribution of the potash beds. Drilling operations in 1959-1961 led to the delineation of the small localized "Crescent" carnallitite deposit in the vicinity of Black Mountain . This was followed by the discovery of the much larger (>80 million tonnes) "Musley" sylvite deposit near the base of the Ethiopian Highlands, some 5km W of Dallol, and extending at least 10km in a N-S orientation. A 92m vertical shaft and a total of 805m of drives were made in this deposit, but all work was stopped in 1967 after rapid influx of water into the conventional mine killed a number of workers. The political tensions in the area at the time probably also played a part in preventing mining activity in the following years.

    Holwerda and Hutchinson (1968) argue that geographical location of the main "Musley" sylvite strata is directly west of Dallol Mound and at the base of the highlands. This, and the fact that sylvite is an alternation product that consistently overlays the carnallite strata and thickens (although discontinuously) along the western margin (see drill hole intersections published in Ercosplan, 2011), suggests that the potash enrichment was produced by selective leaching of MgCl2 from a carnallite precursor, driven by phreatic run-off waters sourced in the Ethiopian highlands. My own observations and plotting of enrichment fairways (using published Ercosplan 2010, 2011 data) confirms Holwerda and Hutchinson’s inferences. If diagenesis, not primary precipitation, is the prime mechanism of sylvite creation in the Musley region, then the regional sylvite control/distribution for this style of enrichment is related to a subsurface meteoric/groundwater phreatic overprint that parallels the encroaching bajada edge. It is a separate ore fairway to the more regional easterly dipping bedded kainitite/carnallitite trend.

    Waters in some of the active brine-filled hydrothermal craters and dolines can locally have temperatures of more than 100°C and when waters cool they precipitate varying combinations of halite, carnallite and bischofite. The brines are so saturated with salts that if a stick is thrust into a boiling brine pool and removed it is immediately covered by layer of carnallite or bischofite and halite (Figure 2b, c). The same pools are also rich in FeCl2, sulphur and manganese, which explains the spectacular bright green, red-orange and yellow colours of many of the saline mineral assemblages precipitating in and about these active spring-formed pools. Occasional intense storm-driven sheetfloods can drive renewed activity in the various springs in vicinity of the mound, as happened in the recent floods of February 2011, when the intensity of water circulation and the areal extent of the pools greatly increased. After the same storm flood, a natural collapse doline tens of metres across formed on the western depression margin. Clearly, the local hydrothermal/karstic enhancement style of bittern enrichment is a separate process set for potash enrichment compared to the widespread earlier deposition of marine-fed subaqueous kainite. Hence, it contrasts with the much more widespread set of depositional/early diagenetic processes that laid down the bulk of the bedded potash association that is the Houston Fm. in the Danakhil Depression (as discussed in the previous Danakhil blog).

    What is the Dallol Mound and what drives its uplift hydrology?

    Despite the widespread misconception that the Dallol mound is a lava cone, Mount Dallol is not a volcanic-centered feature on the Danakhil landscape. A visit to the area reveals no observable volcanic products (lava, ashfall or scoria) on the surface on or near the Dallol mound. This is so even in the region of the most recent phreatic activity in 1926 where a 30 m-diameter phreatic (explosion? or daylighting hydrothermal karst) crater formed, hosted in salt beds (Figure 2b). All the rocks associated with this cavity and its formative event are not volcanic. This means the mechanism that created the Dallol Mound is unlike the magmatic events that created the world famous Erte Ale volcanic cone, with its distinctive longterm active magma lake and located some 80 km to the south of Dallol and still in the Danakhil depression. Instead, the Dallol mound crest is made up of uplifted and eroded halite and potash beds soaked in a thermal hydrology that breaks out on the lake surface as a number of hot bubbling sulphurous brine pools. This is also true of the off-mound crater that formed in 1926 near Black Mountain and still retains bubbling brines with present temperatures ~65-70 °C. Nearby “Black Mountain” is a small area of dark coloured bedded and recrystallised halite, it is not a primary volcanic feature.

    As a sedimentologist visiting the area, I wondered at why the Dallol mound features had ever been called volcanic cones, hornitos, or maars (as they are widely described in the literature). To use such genetic terms in a geologically correct fashion I would like to put my hand on a piece of volcanic debris (lava, pumice, scoria or ash) in any of the craters before I call the Dallol mound a volcanic cone. And yet, many workers in the published literature dealing with the Dallol area are happy to do this. I am not saying there is no influence of magmatic heating in forming Dallol Mound, only that molten volcanic rock has yet to surface in the immediate Dallol region. Hence it is unlike the many actual volcanic cones, maars and hornitos to the south and north and this is an significant observation as it deals with mechanism of local potash enrichment. I will argue in the next section that this is because Dallol Mound is a salt uplift feature or dome capped by phreatic cone/ hydrothermal karst structures and all related to the migration of molten magma into more deeply buried salt beds, which contain hydrated salts at the level of the Houston Fm and perhaps even deeper buried hydrated salt layers (see blog 2).

    Darrah et al (2013) and Detay (2011) argue that the 30m diameter 1926 crater and other nearby pools on the Dallol saltflat in the vicinity of the Dallol mound are the result of a phreatic explosions, tied to the increasing gas pressure in superficial hydrothermal reservoirs atop a deeper mass of molten rock. The mound is a landscape feature indicative of deep dyke/sill intrusion that did not surface. According to Holwerda and Hutchinson (1968) this yet-to-daylight dyke complex explains the linear orientation of the mound, its pools and other karst/erosion features on the salt flat surface in vicinity of the Dallol mound. That is, the various Dallol hot springs typically consist of 30-40m diameter circular to sub-circular ponds, initially formed by explosive vapor eruptions, to form at-surface circular features, which are widely termed maars, although I would prefer to call them "maar-like." A “maar” is defined in the AGI Glossary of Geology as “a low relief, broad volcanic crater formed by multiple shallow explosive eruptions. It is surrounded by a crater ring, and may be filled by water. Type occurrence is in the Eifel area of Germany.” Given the lack of a volcanic crater rim the Dallol Mound and adjacent brine-filled cavities are not really maars, nor are they hornitos. They will likely evolve into such features, but in their current state better considered brine-filled fumaroles or solfateras or even better, hydrothermal karst cavities that have daylighted. Once the cavities have broken out onto the salt flat surface, these circular (possibly-explosive) features can continue enlarge due to ongoing rise of undersaturated waters and so evolve into expanding hydrothermal karst pools or they can be partially to completely filled with saline precipitates (with no volcanic products derived from molten igneous rock materials).


    So, instead of at-surface volcanic products such as lava and ashfall, most of the superficial precipitates/sediments observed in and around the various on- and off-structure Dallol brine pools are evaporite salts, along with some remnants of older clay-sediments. Brine fluids in various hot spring pools in the Dallol area (in the Dallol “hill” crest and the “Crescent” region near Black mountain, and in the “Boiling Lake” region south of the mound) are typically multi-coloured warm/hot ponds (Figure1, 3; Gebresilassie et al., 2011). The various pools are extremely salty (>500g/L), can be highly acidic (sometimes with a pH approaching 0.5), and gas-rich (as evidenced by steady, vigorous bubbling of gases). According to Darrah et al. (2013) the Dallol “salt dome” fluids and associated hot springs are hypothesized to result from the interaction between hot mantle fluids or basalt dyke injections with evaporite deposits at unknown depths. However, direct observations of the volumes of pool waters and the vigour of the outflow are known to increase after the occasional heavy rain event, as happened in February, 2011. Hence, it is unclear if sulfur-rich gases and the low pH brine fluids provide evidence of the interaction of hot mantle fluids with the evaporites (as inferred by Darrah et al., 2013) or the pool waters are, at least in part, related shallower ongoing hydrothermal/karst interactions with more deeply circulated meteoric waters sourced in the 1000-m high adjacent rift highlands.

    Why hydrated salts are important in some salt-hosted thermal systems: a Permian Zechstein analog

    Most published volcanogenic-related studies of the Dallol Mound have not considered the effects of hydrated salt layers in a situation of rising molten rock, where the country rock contains beds of hydrated evaporites such as kainite or carnallite. This situation is exposed in the dyke-intruded halite-carnallite levels in the mines of the Werra-Fulda mine district of Germany (Schoefield et al., 2013; Warren, 2015). There, the Permian Zechstein salt series contains two important potash salt horizons (2-10m thick), which are mined at a depths ≈ 800 m from within a 400m thick halite host (Figure 4a). In the later Tertiary, basaltic melts intruded these Zechstein evaporites, but only a few dykes reached the Miocene landsurface. Basaltic melt production was related to regional volcanic activity some 10 to 25 Ma. Basalts exposed in the mine walls are typically subvertical dykes, rather than sills. These basaltic intervals can crosscut the salt over zones up to several kilometres wide (Figure 4b). However, correlations of individual dyke swarms, between different mines, or between surface and subsurface outcrops, is difficult.


    The basalts are phonolitic tephrites, limburgites, basanites and olivine nephelinites. Dyke margins in contact with halite are usually vitrified, forming a microlitic limburgite glass along dyke edges (Knipping, 1989). At the contact on the evaporite side of the glassy rim there is a cm-wide carapace of high temperature salts (mostly anhydrite and ferroan carbonates). Further out, the effect of the high temperature envelope is denoted by transitions to clear halite, with higher temperature fluid inclusions (Knipping 1989). All of this centimetre to metre-scale alteration is an anhydrous alteration halo, the salt did not melt (halite’s melting temperature is 804°C), rather than migrating, the fluid driving recrystallisation was largely from local movement of entrained brine inclusions. The dolerite/basalt interior of the basaltic dyke is likewise altered and salt soaked, with clear, largely inclusion-free halite typically filling vesicles in the basalt.

    Worldwide, dykes intersecting salt beds tend to widen to become sills in two zones: 1) along evaporite units within the halite mass that contain hydrated salts, such as carnallite or gypsum and, 2) where rising magma has ponded and so created laccoliths at the upper or lower halite contact with the adjacent nonsalt strata or against a salt wall (Warren, 2015). The first is a response to a pulse of released water as dyke-driven heating forces the dehydration of hydrated salt layers. The second is a response to the mechanical strength contrast at the salt-nonsalt contact. The first is what is observed in the Fulda region and is also likely relevant to the formation of the Dallol Mound and its remobilised potash-precipitating brines.

     

    In such subsurface regions, the heating of hydrated salt layers (such as carnallite or kainite), adjacent to a dyke or sill, drives off the water of crystallisation (chemical or hydration thixotropy) at a much lower temperatures than that at which anhydrous salts, such as halite or anhydrite, thermally melt (Table 1). In the Fulda region the thermally-driven release of water of crystallisation within particular Zechstein salt beds creates thixotropic or subsurface “peperite” textures in carnallitite ore layers, where heated water of crystallisation escaped from the hydrated-salt lattice. Dehydration-driven loss of mechanical strength focuses zones of magma entry into particular horizons in the salt mass, wherever hydrated salt layers were intersected (Figure 4c verses 4d). In contrast, dyke and sill margins are much sharper and narrower in zones of contact with anhydrous salt intervals (Figure 4b; Schofield et al., 2014).

    Accordingly, away from immediate vicinity of the direct thermal aureole, heated and overpressured dehydration waters can enter a former Zechstein carnallite halite bed, and drive the creation of extensive soft sediment deformation and [1]peperite textures in the former hydrated layer (Figure 4d, e). Mineralogically, sylvite and coarse recrystallised halite dominate the salt fraction in the peperite intervals/beds. These are evaporite-related beds formed within a hydrated salt bed and so differ from the common notion of volcanic peperites indicating water-saturated sediment intercations with very shallow dyke or sill emplacements. Sylvite in these altered zone is a form of dehydrated carnallite, not a primary-textured salt. In the Fulda region, such altered zones and deformed units can extend along former carnallite layers to tens or even a hundred or more metres from the dyke feeder. Ultimately, the deformed potash bed passes laterally out into the unaltered bed, which retains abundant inclusion-rich primary chevron halite and carnallite (Figure 4d versus 4e). That is, nearer the basalt dyke, the carnallite is largely transformed into inclusion-poor halite and sylvite, the result of incongruent flushing of warm saline fluids mobilized from the hydrated carnallite crystal lattice as it was heated by dyke emplacement. During Miocene salt alteration/thermal metamorphism in the Fulda region, NaCl-fluids were mixed with fluids originating from thermally-mobilised crystallisation water in the carnallite, as it converted to sylvite.

    In the Dallol depression I think it is highly likely that a similar set of destabilisation processes occurred when rising dyke magma reached the levels of hydrated salts (kainite and carnallite beds) in the Houston Formation of the Danakhil fill, after passing relatively passively through the Lower Rocksalt Formation (see the previous blog). Emplacement of the magma/dyke into  hydrated evaporites in the vicinity of what is now the Dallol mound would have mobilised and deformed the hydrated salt level, converting carnallite to sylvite, kainite to bischofite and lesser kieserite, as well as creating widespread cavities filled with pressured volatiles carried by MgCl and KCl brines. Once these hydrothermal cavities dissolve their way to surface, the feeder brines can cool and precipitate as prograde salts such as halite, sylvite and perhaps bischofite. Such destabilisation would have accommodated the emplacement of a basaltic sill at the level of the potash salts, in turn driving the uplift of the lake beds above this region. Mound-related uplift and hydrothermal activity then drives the formation of natural regions of ground collapse, sulphurous and acidic springs and fumaroles, along with the creation of water-filled chimneys and doline sags, filling with various hydrothermal salts, in the vicinity of the volcanic mound.

    Implications for Potash distribution in the Danakhil Depression

    The discussion of potash mineral-forming processes in this and the previous blog clearly underlines a trichotomy in the way potash has accumulated in halite host-beds across the Danakhil Depression. The most widespread form of potash in the Danakhil Depression is as a primary evaporite bed, composed of primary marine kainitite precipitates with a carnallite cap (Houston Formation). Across the western side of the depression this easterly dipping bed is now buried beneath 30-150 m of overburden salts. It likely precipitated as a marine seepage-fed bittern layer, at a time the Danakhil depression was hydrographically isolated from a direct surface connection with the Red Sea. Its brine hydrology was dominantly subaqueous and not unlike that of modern Lake Asal in Djibouti, although it was more saline than Asal in the subaqueous potash sump areas. Thus, the Danakhil potash bed (Houston Fm) formed sometime ago, its formative hydology is no longer present in the depression and it may be as old as Pliocene or more likely early to mid Pleistocene. There has been sufficient time for this bed to tilt toward the east. The unit is underlain by the subaqueous Lower Rocksalt Formation (LRF) and subsequently overlain by the Upper Rocksalt Formation (URF). Both these halite formations do not entrain primary potash beds. The LRF contains numerous CaSO4 layers, while the URF contains clayey laminite beds and locally hosts regions of remobilised potash salts. The URF evolves upward into the saltflat/ephemeral lake hyperarid hydrology that typifies the modern depression.

    More localised forms of potential potash ore typify occurrences in the Dallol and Musley areas (Figure 2a). There potash in the Dallol Mound region is hydrothermally reworked from the uplifted equivalents of the Houston Formation. Even today this hydrology is precipitating carnallitite (associated with bischofite and minor kieserite) in various hydrothermal brine pools atop and around the Dallol Mound, such as the carnallite-dominant Crescent deposit (Figure 2b). These hydrothermal salts owes their origins to daylighting of pressurised fluid systems and cavities. They were created by the volatile products of hydrated salt layers (Houston Fm) where these salts had come into contact with thermal aureoles or actual lithologies of newly emplaced dykes that had penetrated the underlying halite section. Actual molten volcanic rock has yet to make it to the surface in the Dallol Mound region, although active volcanic mounds and flows do typify the saltflat surface tens of kilometres to the south (Erte Alle ) and north. Based on the analogy exposed within the Zechstein-hosted potash mines of the Fulda region of Germany, it is likely that as well as creating at-surface brine pools, this hydrothermal dyke-related hydrology converts any carnallitite to a sylvinite bed at the level of contact with the Houston Fm. 

    Then there is the deep-meteoric alteration system that is altering the kainitite/carnallitite of Houston Fm into sylvinite, it is active along the deep meteoric alteration front located at the irregular interface between the downdip end of the Musley Fan and the updip portion of the Houston Fm. This diagenetic mechanism formed the Musley potash deposit, defined and exploited by the Parsons Company operations and documented in Holwerda and Hutchison (1968). Variations on this deep-meteoric alteration theme likely extend south and north of the Musley fan, wherever the active phreatic hydrology of the bajada located at the foot of the Ethiopian Highlands interacts and interfingers with the updip edge of the easterly dipping Houston Formation.

    Once again there is no "one-size-fits-all) model for economic potash understanding (Warren, 2010, 2015). Even in what is probably the youngest known marine-fed potash system in the world, the original potash mineralogy and distribution has been altered and locally upgraded via diagenetic interactions with hydrothermal or deep-meteoric fluids. Predicting ore distributions in this, and all potash systems worldwide, requires an understanding of formative process evolution through deep time, and not just the simple application of a layer-cake primary stratigraphic model. 

    References

    Carniel, R., E. M. Jolis, and J. Jones, 2010, A geophysical multi-parametric analysis of hydrothermal activity at Dallol, Ethiopia: Journal of African Earth Sciences, v. 58, p. 812-819.

    Darrah, T. H., D. Tedesco, F. Tassi, O. Vaselli, E. Cuoco, and R. J. Poreda, 2013, Gas chemistry of the Dallol region of the Danakil Depression in the Afar region of the northern-most East African Rift: Chemical Geology, v. 339, p. 16-29.

    Detay, M., 2011, Le DALLOL revisité: entre explosion phréatomagmatique, rifting intra-continental, manifestations hydrothermales et halocinèse: LAVE. Liaison des amateurs de volcanologie européenne, v. 151, p. 7-19.

    ERCOSPLAN, 2010, Techical report and current resource estimate: Danakhil Potash Deposit, Afar State, Ethiopia: Project Reference: EGB 08-024.

    ERCOSPLAN, 2011, Preliminary Resource Assessment Study, Danakhil Potash Deposit, Afar State, Ethiopia: G & B Property: Project Reference: EGB 10-030.

    Gebresilassie, S., H. Tsegab, and K. Kabeto, 2011, Preliminary study on geology, mineral potential, and characteristics of hot springs from Dallol area, Afar rift, northeastern Ethiopia: implications for natural resource exploration: Momona Ethiopian Journal of Science, v. 3, p. 17-30.

    Holwerda, J. G., and R. W. Hutchinson, 1968, Potash-bearing evaporites in the Danakil area, Ethiopia: Economic Geology, v. 63, p. 124-150.

    Knipping, B., 1989, Basalt intrusions in evaporites: Lecture Notes in Earth Sciences (Springer-Verlag), v. 24, p. 132 pp.

    Schofield, N., I. Alsop, J. Warren, J. R. Underhill, R. Lehné, W. Beer, and V. Lukas, 2014, Mobilizing salt: Magma-salt interactions: Geology.

    Warren, J. K., 2010, Evaporites through time: Tectonic, climatic and eustatic controls in marine and nonmarine deposits: Earth-Science Reviews, v. 98, p. 217-268.

    Warren, J. K., 2015, Evaporites: A compendium (ISBN 978-3-319-13511-3) Released August 2015: Berlin, Springer, 1600 p.

    ------------------------- 

    [1] Peperite is a sedimentary rock that contains fragments of igneous material and is formed when magma comes into contact with wet water-saturated sediments. 

    Danakhil Potash, Ethiopia: Beds of Kainite/Carnallite, Part 2 of 4

    John Warren - Wednesday, April 29, 2015

    The modern Dallol saltflat described in the previous blog defines the upper part of more than 970 metres of halite-dominated Quaternary evaporites that have accumulated beneath the present salt pan of the Northern Danakhil. The total sequence is made up of interbeds of halite, gypsum, anhydrite and shale with a potash succession separating two thick sequences of halite (Figure 1; Holwerda and Hutchison, 1968; Augustithis, 1980). At depths of more than 35-40 meters, and deepening to the east, this km-thick subcropping Quaternary halite-dominated fill contains one, and perhaps two or more, potash beds. For a more detailed description of the upper part of the fill the reader is referred to the previous blog and Chapter 11 in Warren, 2015.


    Bedded Pleistocene evaporites may underlie the entire Danakil depression, but younger lava flows of the Aden Volcanic Series and alluvium washed in from the surrounding bajada obscure much of the older Pleistocene sedimentary series across much of the southern part of the depression beyond Lake Assale). Potash exploration drilling and core recovery is concentrated in the accessible parts of the northern Danakhil rift, where the saltflat facilitates vehicle access, compared with the lava-covered regions south of Lake Assale. The most recent volcanic activity affecting the known potash region was the emplacement of the Dallol Mound, which has driven local uplift of the otherwise subsurface potash section to where it approaches the surface in the immediate vicinity of the mound (Figure 2a).

    Away from the Dallol volcanic mound the upper potash bed beneath the saltflat lies at a depth of 38-190 metres. A lower inferred potash bed likely occurs at depth along the eastern end of the saltflat, but this second bed is inferred from high API kicks in gamma logs run in deeper wells, no solid salt was recovered (Holwerda and Hutchison, 1968). The upper proven potash bed is now the target zone for a number of minerals companies currently exploring for potash in the region. Regionally, both potash units dip east, with the deepest indicators of the two units encountered by the drill in a single well on the eastern side of the saltflat at depths of 683 and 930 m, respectively (Figure 2: Holwerda and Hutchison, 1968). The likely Quaternary age of the potash units, the marine brine source, explains the high magnesium content of the potash bittern salts, as modern seawater contains high levels of Mg and SO4.


    My study of core that intersected the potash interval and that is sandwiched between the Lower and Upper Halite units shows both the lower and the upper halite units retain pristine sedimentary textures, with features and vertical successions that indicate distinct hydrologies during their deposition (Figure 3). There is no textural evidence of halokinetic recrystallization in halites any of the studied cores and published seismic also indicates consistent dips in the evaporites . Most of the textures in the cored potash interval indicate a subaqueous density-stratified environment, with brine reworking of the upper part of primary kainitite, carnallitite units. Perennial subaqueous, density-stratified brines also typify the hydrology of the Lower Halite unit, albeit with somewhat lower salinities tan those precipitating the bitterns (Figure 3). The brine that precipitating the Upper Rocksalt Formation was shallower and more ephemeral. The following paragraphs summarise my core-based observations and interpretations that led to this interpretation of the evolving brine hydrology.

    The Lower Rocksalt Formation (LRF) is dominated by bottom-growth-aligned subaqueous halite textures and lack of siliciclastic detritus, unlike the Upper Rocksalt Formation (Figure 3). Halite textures in the LRF lack porosity and dominated by coarsely crystalline beds made up of cm-scale NaCl-CaSO4 couplets dominated by upward-pointing halite chevrons and mantled by thin CaSO4 layers (Figure 3). This meromictic-holomictic textural association passes up into the upper part of the LRF with cm-scale proportions alternating of less-saline to more-saline episodes of evaporite precipitation decreasing, indicating an “on-average” increasingly shallow subaqueous depositional setting as one approaches the base of the kainitite unit. The combination of bottom-nucleated and cumulate textures in the LRF are near identical to those in the halites in the kainitite interval in the Messinian of Sicily (see later). 

    The laminated Kainitite Member is also a subaqueous unit with layered cumulate textures (Figure 3), it was likely deposited on a pelagic bottom beneath a shallow body of marine-fed bittern waters, which never reached carnallite saturation. Above this are the variably present carnallitic Intermediate and Sylvinitite members and the overlying Halite marker beds in turn overlain by the Upper Halite unit. All retain pristine textures indicating mostly subaqueous deposition, soon followed by varying exposure and reaction with shallow phreatic brines moving across the top of Kainitite member. This shallow phreatic brine crossflow drove syndepositional mineral alteration and collapse in the upper part of the kainitite and carnallitite units.


    The potash-entraining interval between the URF and LRF is called the Houston Formation has been drilled and cored extensively by explorers in the basin, showing it is consistently between 15 and 40 metres thick (Figure 1). Stratigraphically, it consists of lower Kainitite Member (4-14m thick) atop and in depositional continuity with the LRF (more than 500m thick) (Figure 3). The Kainitite Member is fine-grained, laminated, locally wavy-bedded, containing up to 50% kainite cumulates in a cumulate (non-chevron) halite background, along with small amounts of a white mineral that is likely epsomite. It is overlain by what older literature describes as the “Carnallitite Intermediate unit” (3-25 m thick). More recent potash exploration drilling has shown all the members that constitute the Intermediate Carnallitite Member is not always present within the Dallol depression. Mineralogically is at best considered as variably developed (Figure 3). Its lower part is a layered to laminated carnallite-halite mixture with some kieserite, anhydrite and epsomite. This can pass up or laterally into kainitite with sylvite. Above the Intermediate Member is the 0-10m thick Sylvinite Member containing 20-30% sylvite, along with polyhalite and anhydrite (up to 10%). Typically the sylvinite member shows primary layering disturbed by varying intensities of slumping and dissolution. Often the upper part of a carnallite unit (where present) also shows similar evidence of dissolution and reprecipitation.

    Cores through the sylvinite member and parts of the upper carnallitite member sample a range of recrystallization/flow/slump textures, rather than primary horizontal-laminar textures. Beneath the sylvinite member, the variably-present upper carnallitite member contains a varied suite of non-commercial potash minerals that in addition to carnallite include, kieserite, kainite (up to 10% by volume) and polyhalite, along with minor amounts of sylvite. Minor anhydrite is common, while rinneite may occur locally, along with rust-red iron staining. Sylvite is more abundant near the top of the carnallitite member and its proportion decreases downward, perhaps reflecting its groundwater origin. Kainite is the reverse and its proportion increases downward. The sylvinite member and the carnallitite member also show an inverse thickness relationship. Bedding in the carnallitite member is commonly contorted with folded and brecciated horizons interpreted as slumps. The base of the carnallitite member is defined as the level where carnallite forms isolated patches in the kainite before disappearing entirely.

    Drilling in the past few years has clearly show that in some parts of the evaporite unit, located nearer to the western side of the basin, the lower and upper carnallite units are separated by thick bischoftite intervals (Figures 2b, 3). The bischofite is layered at a mm-cm scale and with no obvious breaks related to freshening and exposure, implying it too was deposited in a perennially subaqueous or phreatic cavity setting (Pedley et al., in press).

    The potash/bischofite interval passes up into a slumped and disturbed halite-dominated unit that is known as the “Marker Beds” because of the co-associated presence of clay lamina and bedded halite, along with traces of potash minerals (Figure 3). This unit then passes up into the massive Upper Rock Salt unit across an unconformity at the top of the halite “Marker Beds.” Bedded, and at times finely laminated cumulate textures in the various magnesian bittern units, are used by many to argue that the kainitite and the lower carnallitite members are primary or syndepositional precipitates.

    Three types of potash-barren zones can occur within it and are possibly related to the effects of groundwaters and solution cavity cements within the carnallitite unit, perhaps precipitated before the deposition of the overlying marker halites. Barren zones in the Sylvinite member are regions where: a) the entire sylvinite bed is replaced by a relatively pure stratiform halite, along with dispersed nodules of anhydrite, b) zones up to 23 m thick and composed of pure crystalline halite (karst-fill cements?) that occur patchily within the sylvinite bed and, c) potash-depleted zones defined by coarsely crystalline halite instead of sylvinite. Bedding plane spacing and layering and some slumping styles in the halite in styles a and b are similar to that in the sylvite bed. Contact with throughflushing freshened nearsurface and at-surface waters perhaps created most of the barren zones in the sylvinite. Fluid crossflow may also have formed or reprecipitated sylvite of the upper member, via selective surface or nearsurface leaching of MgCl2 from its carnallite precursor (Holwerda and Hutchison, 1968; Warren 2015). Due to the secondary origin of much of the sylvite in the Sylvinite member, the proportion of sylvite decreases as the proportion of carnallite increases, along with secondary kieserite, polyhalite and kainite.

    The kainite member is texturally distinctive and is composed of nearly pure, fine-grained, dense, relatively hard, amber-coloured kainite with ≈ 25% admixed halite (Figure 3). Core study shows the lamina style remained flat-laminar (that is, subaqueous density-stratified with periodic bottom freshening) as the mineralogy passes from the LRF up into the flat-laminated kainitite member (Figure 4: Warren, 2015; Pedley et al., in press). Throughout, the kainitite unit shows a cm-mm scale layering, with no evidence of microkarsting or any exposure of the kainitite depositional surface. That is, the Kainitite Member is a primary depositional unit, like the underlying halite and still retains pristine evidence of its dominantly subaqueous depositional hydrology. The decreased proportion of anhydrite in the Kainitite Member, compared to the underlying LRF, indicates a system that on-average was more saline than the brines that deposited the underlying halite. The preponderance of MgSO4 salts means the Kainitite unit like the underlying LRF formed by the evaporation of seep-supplied seawater.

    This situation differs from the present “closed basin” hydrology of the Danakil Depression which typifies the URF and the overlying Holocene succession (Hardie, 1990; Warren, 2015).

    Units atop the primary laminated textures of the kainitite, lower carnallitite and bischofite members (where present) tend to show various early-diagenetic secondary textures (Figure 4). It seems much of the sylvinite and upper carnallitite member deposition was in shallow subsurface or at-surface brine ponds subject to groundwater crossflows and floor collapse, possibly aided by seismically-induced pulses of brine crossflow. In addition, this perennial density-stratified brine hydrology was at times of holomixis subject to brine reflux and the brine-displacement backreactions that typify all evaporite deposition, past and present (Warren 2015).

    The observation of early ionic mobility in potash zone brines in the Danakil depositional system is also not unusual in any modern or ancient potash deposit. It should not be considered necessarily detrimental to the possibility of an extensive economically exploitable potash zone being present in the Danakil Depression. Interestingly, all the world’s exploited potash deposits, including those in the Devonian of Canada and Belarus, the Perm of the Urals and the potash bed of west Texas, show evidence of syndepositional and shallow burial reworking of potash (Warren, 2015). Early potassium remobilization has created the ore distributions in these and other mined potash depositsTextures and mineralogies in the Upper Rocksalt Formation (URF) define a separate hydrological association to the marine-fed LRF and Houston Formation (Table 4). Compared to the LRF, the URF has much higher levels of depositional porosity, lacks high levels of CaSO4, and has high levels of detrital siliciclastics. This is especially so in its upper part, which shows textural evidence of periodic and ongoing clastic-rich sheetflooding and freshening (Figure 4). It was deposited in a hydrology that evolved up section to become very similar to that active on today’s halite pan surface. The URF contains no evidence of salinities or textures associated with a potash bittern event and is probably not a viable exploration target for solid potash salts.

    Above the URF is a clastic unit with significant amounts of, and sometimes beds dominated by, lenticular gypsum and displacive halite. The unit thickens toward the margins of the depression (Figure 2). The widespread presence of diagenetic salts indicates high pore salinities as, or soon after, the saline beds that stack into the clastic unit were deposited. Some of these early diagenetic evaporite textures are spectacular, as seen in the displacive halite recovered in a core from the lower portion of the clastic overburden, some 45 m below the modern pan surface (Figure 3).

    What is clear from the textures preserved in the potash-rich Houston formation and the immediately underlying and overlying halites is that they first formed in a subaqueous-dominated marine-fed hydrology (Figure 4), which evolves up section into more ephemeral saltpan hydrologies of today (see the previous blog). The potash interval encapsulated in the Houston formation has primary mineralogical associations that are derived by evaporation of Pleistocene seawater (kainitite, carnallitite). In contrast the sylvite section in the Houston tends to form when these primary mineralogies are altered diagenetically perhaps soon after deposition but, especially, when hydrothermal waters circulated through uplifted beds of the Houston Formation, as is still occurring in the vicinity of the Dallol Volcanic Mound. Or where the chemical/meteoric interface associated with the encroachment of the bajada sediment pile drove incongruent dissolution of carnallite along the updip edge of the Houston Fm (as we shall discuss in the next blog). 

    References

    Augustithis, S. S., 1980, On the textures and treatment of the sylvinite ore from the Danakili Depression, Salt Plain (Piano del Sale), Tigre, Ethiopia: Chemie der Erde, v. 39, p. 91-95.

    Hardie, L. A., 1990, The roles of rifting and hydrothermal CaCl2 brines in the origin of potash evaporites: an hypothesis: American Journal of Science, v. 290, p. 43-106.

    Holwerda, J. G., and R. W. Hutchinson, 1968, Potash-bearing evaporites in the Danakil area, Ethiopia: Economic Geology, v. 63, p. 124-150.

    Pedley, H. M., J. Neubert, and J. K. Warren, in press, Potash deposits of Africa: African Mineral Deposits, 35TH International Geological Congress (IGC), Capetown (28 August to 4 September 2016).

    Warren, J. K., 2015, Evaporites: A compendium (ISBN 978-3-319-13511-3) Released August 2015: Berlin, Springer, 1600 p. 

    Danakhil Potash, Ethiopia: Is the present geology the key? Part 1 of 4

    John Warren - Sunday, April 19, 2015

    Geology of potash in the Danakil Depression, Ethiopia: Is the present the key?

    The Danakhil region, especially in the Dallol region of Ethiopia, is world renowned for significant accumulations of potash salts (both muriates and sulphates), and is often cited as a modern example of where potash accumulates today. What is not so well known are the depositional and hydrological dichotomies that control levels of bittern salts in the Pleistocene stratigraphy that is the Danakhil fill. Geological evolution of the potash occurrences in the Dallol saltflat and surrounds highlights the limited significance of Holocene models for potash, when compared to the broader depositional and hydrological spectra preserved in ancient (Pre-Quaternary) evaporite deposits (see Warren, 2010, 2015 for a more complete analysis across a variety of evaporite salts).

    Across the next four blogs, I shall discuss the geological character of the Danakhil fill and the controls on potash in the depression via four time-related discussions; A) Current continental fan - saltflat hydrology that typifies present and immediate past deposition in the depression (Danakhil Blog1). B) A time in the latest Pleistocene when there was a marine hydrographic connection exemplified by a healthy coralgal rim facies (probably ≈ 100,000 years ago, and a subsequent drawndown gypsum rim facies. Both units are discussed in this blog, (Danakhil Blog1), and C) a somewhat older Pleistocene period when widespread potash salts were deposited via a marine seepage fed hydrology (Danakhil Blog2). Then, within this depositional frame, we will consider D) the influence of Holocene volcanism and uplift driving remobilisation of the somewhat older potash-rich evaporite source beds into the Holocene hydrology (Danakhil Blog3) and finally how this relates to models of Neogene marine potash deposition (Danakhil Blog4). These observations and interpretations are based in large part on a two-week visit to the Dallol, sponsored by BHP minerals, and focused on the potash geology of the region. 

    Dallol Physiography

    The Danakhil Depression of Ethiopia and Eritrea is an area of intense volcanic and hydrothermal activity, with potash occurrences related to rift magmatism, marine flooding, and deep brine cycling. The region is part of the broader Afar Triple Junction and located in the axial zone of the Afar rift, near the confluence of the East African, Red Sea and Carlsberg rifts (Figure 1a; Holwerda and Hutchison, 1968; Hutchinson and Engels, 1970; Hardie, 1990). The depression defines the northern part of the Afar depression and runs SSE parallel to the Red Sea coast, but lies some 50 to 80 km inland, and is separated from the Red Sea by the Danakil Mountains. The fault-defined Danakil Depression is 185 km long, up to 70 km wide, with a floor that in the deeper parts of the depression is more than 116 meters below sea level. It widens to the south, beginning with a 10 km width in the north and widening up to 70 km in the south (Figure 1a). In the vicinity of Lake Assele, the northern portion of the Danakil is known as the Dallol Depression and has been the focus for potash exploration for more than a century and is in the deepest region of the depression with elevations ranging between 50m to 120m below sealevel (Figure 1b, c). Shallow volcano-tectonic barriers, behind Mersa Fatma, Hawakil Bay and south of the Gulf of Zula, prevent hydrographic (surface) recharge to the depression. Marine seepage is not occurring at the present time, but likely did so at the time the main potash unit was precipitated. Lake Assele (aka Lake Kurum) with a water surface at -115m msl should not be confused with Lake Asal (-155 msl), located 350 km to the southeast of the Danakil. Asal an active marine-fed hydrographically isolated lacustrine drawdown system, which today is depositing a combination of pan halite and subaqueous gypsum in the deepest part of the Asal-Ghubbat al Kharab rift (Figure 1a; Warren, 2015).

    Today the halite-floored elongate saltpan, known as the Dallol saltflat, occupies the deepest part of the northern Danakil Depression, extending over an area some 40 km long and 10 km wide (Figure 1b, c). The pan’s position is asymmetric within the Danakil Depression; it lies near the depression’s western edge, some 5km from the foot of the escarpment to the Balakia Mountains and the Ethiopian Highlands, but some 50 km from the eastern margin of the depression, which is in Eritrea. The Dallol saltpan and adjacent Lake Assele today constitute the deepest continental drainage sump in the Afar depression (Figure 1b, c). The area, located east and northeast of the main modern Dallol saltpan depression, is mostly an extensive gypsum plain (Bannert et al., 1970). As we shall see, the gypsum pavement, and its narrower equivalents on the western basin flank, defines a somewhat topographically higher (still sub-sealevel) less-saline, lacustrine episode in the Dallol depression history fill. To the south of the Dallol salt pan, bedded Pleistocene evaporites may underlie the entire Danakil depression, but younger lava flows of the Aden Volcanic Series in combination with alluvium washed in from the surrounding bajada obscure much of the older Pleistocene sedimentary series in southern part of the depression beyond Lake Assele (Figures 1a).

    Climate

    In terms of daily and monthly temperatures, the Dallol region currently holds the official record for highest average, year-round, monthly temperatures; in winter the daily temperature on the saltflat is consistently above 34°C and in summer every day tops 40°C, with some days topping 50°C (Figure 2; Oliver, 2005). These high temperatures and a lack of rainfall, typically less than 200 mm each year, place the Dallol at the hyperarid end of the world desert spectrum and so it lies at the more arid end of the BWh Köppen climate zone (Kottek et al., 2006; Warren, 2015).


    History of extraction of salt products and their transport (Table 1)

    Using little-changed extraction and transport methods, salt (halite) has been quarried by local Afar tribesmen for hundreds of years. First, using axes, a crust of pan salt is chopped into large slabs (Figure 3a). Then workers fit a set of sticks into grooves made by the axes. Next, working the stick, workers lever slabs of bedded salt, which is cut into rectangular tiles of standard size and weight, called ganfur (about 4kg) or ghelao (about 8kg). Tiles are stacked, tied and prepared for transport out of the depression on the backs of camels and donkeys (Figure 3b). Around 2,000 camels and 1,000 donkeys come to the salt flat every day to transport salt tiles to Berahile, about 75 km away. Previously, salt tiles were carried via camel train to the city of Mekele, some 100 km from the Danakil. Mekele, located in the Ethiopian highlands is known as the hub of Ethiopia’s former “white gold” salt trade and still today is known as the “old” salt caravan city. Today, the salt caravans walk the extracted salt to Berahile, located some 60 km from Mekele. From there, trucks transport the salt to Mekele. Each truck can transport up to 350 camel salt loads. From the Mekele salt market, Dallol salt blocks are transported and sold to all parts of Ethiopia for use mainly as table salt or as an add-on in animal feed. The lifestyle of the miners and the camel trains is likely to change in the next few decades as sealed roads are now under construction that will link Mekele to Dallol.


    Once potash (sylvite and carnallite) was discovered in the Dallol region in 1906, an Italian company by the name of Compagnia Mineraria Coloniale (CMC) established the first mining operation. In 1918 a railway was completed from the port of Mersa Fatma to a termination some 28 km from Dallol (Table 1). Rail construction took place from 1917-1918, using what was then the British and French “military-standard” 600 mm rail-gauge Decauville system. "Decauville" rail construction used ready-made sections of small-gauge track and so the trackway was rapidly assembled; <2 years to complete more than 50 km of track. Once completed, the railway transported extracted potash salt from the "Iron Point" rail terminal near Dallol, via Kululli to the port. Potash production is said to have reached some 50,000 metric tons in the 1920s, extracted from an area centred on the Crescent Deposit, which is located near the foot of uplifted lake beds on the southern flanks of Mt Dallol. However, significant salt production had ceased by the end of the 1920s, as large-scale mines in Germany, the USA, and the USSR began to supply the world market with cheaper product. Unsuccessful attempts to reopen potash production were made in the period 1920-1941. Between 1925-29 some 25,000 tons of sylvite were shipped by rail from the Dallol, with a product that averaged 70% KCl. After World War II, the British administration dismantled the railway and removed all traces of it. In 1951-1953, the Dallol Co. of Asmara sold a few tons of product from the Dallol.


    The potash concession title was transferred to the American “Ralph M. Parsons Company” (Parsons) at the end of the 1950s. Parsons initiated the first systematic exploration for potash in the Danakil depression and drilled more than 250 exploration holes during their 9-year evaluation campaign. Major potash resources were confirmed a few km west of Mount Dallol, in a mineralized zone that was named the “Musley” Deposit. Following on from positive exploration results, they began an engineering study to investigate potential processing and mining methods for the Musley Deposit and subsequently in October 1965 sank a shaft into the orebody. They installed underground mine facilities and established a pilot processing plant on surface, to investigate recovery from the bulk samples collected from the underground workings. They envisaged developing the Musley Deposit as a conventional room-and-pillar operation and to this end developed six underground drifts totalling some 805 m in length. Unconfirmed reports suggest that an influx of water flooded the mine (possibly triggered by a seismic event) and after failed attempts to solve the water problem, the activities Parsons ceased activities in 1968. As of end 2014, some salt block buildings built by the Italian and other companies still partially stand as ruins, along with rusting equipment.

    Based on the previous work conducted by Parsons, a German potash producer, Salzdetfurth AG (SAG), began a new exploration campaign in the Danakil Depression in 1968 and 1969. In addition to their work in the Dallol depression, SAG drilled a number of wells in a concession south of Lake Assale, and conducted a geological mapping campaign as far north as Lake Badda, on the border with Eritrea. SAG’s exploration work away from the known Dallol deposits did not prove fruitful as they drilled only one drill hole that reached the potash level. This drill hole, located approximately 25 km to the southeast of Mount Dallol, intersected a kainitite bed, with no sylvinite intersection. The SAG concession was returned to state authorities of Ethiopia. Subsequent drilling by other explorationists in this region has confirmed the deepening of the kainitite level to the southeast of Dallol and the lack of sylvinite at greater depths.

    Since the dismantling of the railway, there has been no high-volume transport system to carry potash product the Red Sea coast. Currently, the Ethiopian Government is constructing all-weather roads from Dallol to Mekele and Afdera When complete this road system will facilitate transport of future potential potash product from the Dallol to Afdera, from where existing roads provide access to Serdo and from there to the seaport of Tadjoura in Djibouti (Figure 1a). This section requires an addition 30 km of all-weather road to be completed to the coast and will facilitate cost-effective transport of potash product to the large agricultural markets of India and China. The transport distance to the Eritrean coast from Dallol is much shorter, but political considerations mean such a route is not a viable option at the present time.

    EVAPORITE DEPOSITIONAL PATTERNS IN OUTCROP

    Surficial sediment distributions outline classic drawdown facies belts in the Dallol region, with a wadi-fed alluvial fan fringe passing down dip into sandflats (local dune fields), dry mudflats (with springs), saline mudflats and ephemeral to perennial brine pans of Lake Assele (Figure 1b). The fans, especially along the western margin of the depression are indented or locally covered by a mostly younger succession of constant-elevation marine, biochemical and evaporitic sediments fringes or “bathtub rim” facies (Figure 4).

     

    Alluvial Fan fringe (Bajada) 

    Pleistocene alluvial/fluvial beds, exposed by local uplift, deflation and ongoing watertable lowering, outcrop about updip edges to the salt-crusted parts of the northern Danakil, and form low flat-topped plateaus or mesas on the plain. These mesas define the tops of alluvial fans aprons, which are heavily dissected and eroded by occasional storm runoff and rainfall. This fan fringe contains relatively fresh water lenses in a desert setting that is one of the world’s harshest (Kebede, 2012). Most of the depositionally active fans line the western margin of the basin and many of the downdip fan edges occur slightly up dip a still-exposed gypsum pavement (Figure 5a), showing depositional equilibration largely with an earlier higher lake stage, while others, such as the Musley fan, have flowed across cut into the gypsum pavement level and now feed water and sediment directly into the edges of the saltflat that defines the lower parts of the depression (Figure 4). Watercourses of the fans that have dissected earlier wadi (bajada) deposits as well as the earlier lacustrine gypsum and limestone pavements so create excellent windows into the stratigraphy of these units. Fan avulsion is indicated by palaeosol layers exposed by downcutting of younger streams (Figure 5b, c).

     

    The Musley fan characteristics are well documented by current and previous potash explorers in the basin as these permeable gravels and sands store a reliable water source for potential solution mining/ore processing in the Musley area and so has been cored by a number of proposed water wells. Internally, the fan is composed of interfingering layers and lenses of sand, gravel and clay (paleosols), with highly porous intervals in the sand and gravels (Figure 5b, c). Depth to the watertable varies from >2m to 60m, and salinities from 760 ppm to more than 23,400 ppm. The principal source of recharge is flash flooding, originating in Musley Canyon, which drains the Western Escarpment, along with minor inflows from the adjacent uplifted volcanic block and local highly intermittent rains (Figure 1a). Of six potential water wells drilled in the fan by the Ralph M. Parsons Company in the 1960s, four returned water of good quality (<2000 mg/l), while the other two had waters with salinities in excess of 20 g/l. Pumping test data indicate average transmissivity of the water-bearing beds around 870m2/day, with salinities in the fan increasing from west to east, approaching the saltflat.

    Chemical sediments outcropping in the depression

    Overall, surface sediment patterns in the Danakil depression define a depositional framework of brine drawdown, related to basin isolation from an earlier hydrographic (at surface) marine connection to the Red Sea, followed by stepped evaporative drawdown. This is indicated by fringing topographically-distinct belts or rims of now inactive coralgal carbonates and gypsum evaporites (aka “bathtub ring” patterns) that cover earlier Pleistocene and Neogene clastics (Figures 3a, e, Figure 4). These “rims” of marine limestone and subsequent gypsum were followed by today’s drawdown saline-pan halite-dominant hydrology (Figures 4, 7a-c). The current hydrological package of sediments encompassing the current drawdown episode lies atop and postdates the Pleistocene potash-hosting Lower Halite Formation in the depression and is probably equivalent to the uppermost part of the clastic overburden facies, as illustrated in the drilled and cored portions of the depression stratigraphy. As we shall discuss in the next blog, only the uppermost portion of the recovered core stratigraphy has equivalents in current depression hydrology (Figure 6). 


    In earlier work, some authors interpreted the fringing belts, especially the exposed coralgal reef belt, as being possibly of Pliocene or even Miocene age. However, when one looks at the stratiform nature of the outcrop trace of both the reef belt and the gypsum belt, and the carapace nature of its depositional boundaries in the field, it is immediately apparent they must be younger (Figure 5a, c; Figure 7). Both the reefal and gypsum belts track horizontal hydrological intersections with the landscape, in what is an ongoing volcanogenic and tectonically active depression. When the reefal belt image is overlain by a DEM it shows the reef belt is consistently at sea level (Figure 1c). If the outcrops of the reef belt and the gypsum pavement were older than late Pleistocene or Holocene, then ongoing episodes of tectonism and volcanism would have modified the elevations of the two outcrop belts in the landscape, as is seen in Miocene redbed outcrops. These underlying and centripetal Miocene sections clearly show the influence of ongoing tilting and tectonism and hence why the flat-lying tops to the reef and gypsum belts imply a late Pleistocene-Holocene (Figure 5d).

    That is, the topographic distribution of the top of the reef facies, which lies within a metre or two of current sea level, implies that the Danakil depression had a relatively recent connection to the Red Sea. The pristine preservation of aragonitic corals and sand dollars in the adjacent marls suggest the connection was either related to the penultimate interglacial (around 104, 000 years ago) or to an early Holocene transgression into the depression. Bannert et al. (1970) assign a C14 age of 25.4-34.5 ka to this formation. However, we consider this is unlikely as DEM overlay levelling shows the reef rim, wherever it outcrops, lies within a meter of current sealevel. World eustacy clearly shows that sealevel was more than 50-60m below its present level some 25,000-30,000 years ago. A 25-35 ka determination of the reef rim would require the whole basin was subject to a single basinwide wide vertical uplift event that did not fragment or disturb the lateral elevation of the rim.

    The coralgal reef terrace indicates normal marine water were once present in the Dallol depression, while the occurrences of the stratiform gypsum pavement are consistent with a former arid lake hydrology at a somewhat lower elevation than the reef rim (Figures 1c, 5a). Like the reef rim, the gypsum pavement fringe defines a consistent elevation level or surface, most clearly visible along the western margin of the present salt flat. It is the result of gypsum deposition during a period of drawdown associated with brine level stability, subsequent to the isolation of the depression from its former “at-surface” marine connection. During this time gypsum accumulated as a stack of subaqueous aligned gypsum beds, along with a series of gypsiferous tufas and rhizoconcretions in zones about the more marginward spring-fed parts of the gypsiferous lake margin (Figure 7d-f). The evolution from marine waters that deposited the reefs and adjacent echinoid-rich lagoonal marls at a higher level in the depression (the lit zone) into a more saline seepage-fed system, with no ongoing marine surface connection to the Red Sea is indicated by the diagenetic growth of large lenticular (“bird’s-beak”) gypsum crystals within the marine marls and the dominant subqueous bottom-nucleated textures in the gypsum beds. In a similar way, the now-outcropping subaqueous-gypsum drawdown rim deposits, located at higher elevations than current saline pan levels typify other drawdown saline lakes in the Afar region, such as Lake Asal in Djibouti, all such occurrences indicate an earlier, somewhat less saline, hydrological equilibrium level (Warren, 2015).


    Active today is the lowest parts of the Dallol saltflat is an ephemeral saltpan hydrology indicated by bedded salt crusts dominated by megapolygonal crusts made up of aligned-chevron halite stacks separated by mm-cm thick mud layers . This current pan hydrology is associated with even greater drawdown levels compared to the former gypsum-dominant hydrology (Figure 8). Current deposits, made up a series of stacked brine-pan salt sheets,  are still quarried as a renewable resource by the local tribesmen (Figure 3). These modern brine flats accumulate pan halite whenever the Lake Assele brine edge (strandline) is periodically blown back and forth over the modern brineflat. It driven by southerly winds, which are frequent in the annual weather cycle, and can move thin sheets of brine kilometres across the pan in a few hours (Figure 1a, Figure 8). Superimposed on this southerly supply of brine is an occasional land-derived sheetflood event, driven by rare rainstorms and the deposition of silt-mud layers from water sheets sourced from the adjacent wadi belt. This ephemeral brineflat hydrology is stable with respect to the current climate (groundwater inflow ≈ outflow). It means the current brineflat of the northern Danakil low is  accumulating bedded pan salt at an even lower topographic level in the basin than the surrounding gypsum pavement, so implying today’s halite-dominant pan beds form under more arid conditions (less humid, more drawndown) than that of the gypsum pavement.


    This stepped (reef to gypsum to halite) late-Pleistocene-early Holocene hydrology, captured in the modern surficial geology of the Dallol Depression, likely postdates a somewhat wetter (humid) climatic period indicated by the widespread deposition of a clastic overburden unit, atop the Upper Halite Formation (UHF; Figure 6). That is, the modern hydrology in present-day Lake Assale, and the adjacent saline mud flats of the Dallol pan, is not the same hydrology as that which precipitated the massive salt of the Upper Halite Formation (UHF). A potash-free halite unit extensively cored beneath the present clastic-dominated saline pan (to be discussed in the next blog). Texturally and hydrologically the depositional system of stacked salt crusts, which dominates the upper part of the UHF in the cored wells, is similar to today's halite-dominated passage from the salt flat into the subaqueous Lake Assale. However, as we shall see, a wetter moister period, dominated by sheetfloods and higher amounts of clastics, separates the two hydrological events in all the cored wells. Today's outcrop geology of alternating saltpan and clastic beds are a different to marine-fed seepage hydrology formed the Lower Halite Formation (LHF), with its potash bittern cap (Houston Formation). 

    Most importantly there is no evidence of primary potash deposition in the modern lake/ pan hydrology of the Dallol saltflat. It is clear that the world-famous bedded potash (mostly kainitite) units of the Danakhil accumulated in a bittern hydrology that is not present in today's Dallol depositional hydrology (Blog 2). As we shall see, Holocene potash only occurs in the vicinity of the Dallol Volcanic Mound, where uplift has moved older, formerly buried, potash beds into a more active hydrothermal hydrology (Blog 3).

    References

    Bannert, D., J. Brinckmann, K. C. Käding, G. Knetsch, M. KÜrsten, and H. Mayrhofer, 1970, Zur Geologie der Danakil-Senke: Geologische Rundschau, v. 59, p. 409-443.

    Ercosplan, 2011, Resource Report for the Danakhil Potash Deposit, Afar State Ethiopia, comissioned by Allana Potash. Document EGB 11-008.

    Hardie, L. A., 1990, The roles of rifting and hydrothermal CaCl2 brines in the origin of potash evaporites: an hypothesis: American Journal of Science, v. 290, p. 43-106.

    Holwerda, J. G., and R. W. Hutchinson, 1968, Potash-bearing evaporites in the Danakil area, Ethiopia: Economic Geology, v. 63, p. 124-150.

    Hutchinson, R. W., and G. G. Engels, 1970, Tectonic significance of regional geology and evaporite lithofacies in northeastern Ethiopia: Philosophical Transactions of the Royal Society, v. A 267, p. 313-329.

    Kebede, S., 2012, Groundwater in Ethiopia: Features, numbers and opportunities, Springer.

    Kottek, M., J. Grieser, C. Beck, B. Rudolf, and F. Rubel, 2006, World Map of the Köppen-Geiger climate classification updated: Meteorologische Zeitschrift, v. 15, p. 259-263.

    Warren, J. K., 2010, Evaporites through time: Tectonic, climatic and eustatic controls in marine and nonmarine deposits: Earth-Science Reviews, v. 98, p. 217-268.

    Warren, J. K., 2015, Evaporites: A compendium (ISBN 978-3-319-13511-3) Released August 2015: Berlin, Springer, 1600 p.

    Salt ablation indicators; how flowing salt dissolves at the surface

    John Warren - Tuesday, March 10, 2015

    When a salt sequence is in contact with undersaturated waters it dissolves to leave behind layers of salt dissolution residues and breccias. The process of salt dissolution is most obvious and most visually stunning is where solution collapse dolines or stoping karst chimneys daylight, as was discussed in an earlier blog dealing with salt collapse features forming today in the vicinity of Uralkali’s Solikamsk 2 mine. Although spectacularly catastrophic, such features are not typical of an area where a halokinetic salt mass is flowing and ebbing across the landsurface. The continual resupply of salt prevents such features forming. Instead, stratiform or stratabound dissolution residue layers and perhaps breccia horizons, along with salt welds, are the more typical indicators of the former presence of a flowing and dissolving salt mass. In this essay I illustrate this by focusing on the geological development of dissolution features within and about three current or former halokinetic salt masses (namakiers), currently outcropping at the surface in central Iran. The three salt features are Kuh-e-Namak (Qom), Kuh-e-Namak (Shurab) and Kuh-e-Gach (Saveh). Listed in this order they illustrate a sequence of dissolution from an initial salt fountain or namakier, through supra-salt overburden foundering, to a time when all at-surface halite has dissolved and only a carapace of gypsum and clay residues remain. Similar features are seen in the rock record, wherever salt ablation breccias occur.

    Namakier is a term suggested by Martin Jackson utilising the comparison to an ice glacier, where namak is the Persian (Farsi) word for salt. Gach is Farsi for gypsum. while kuh means mountain. So, the term kuh-e-namak when translated means “mountain of salt.” There are many such mountains of salt (namakiers) composed of Oligo-Miocene salt in Central Iran and mountains of Neoproterozoic/Cambrian salt in the Zagros fold belt, located further to the south and southeast of Qom. This means it is unwise to describe any salt glacier as kuh-e-namak, without a geographic suffix to better define its location. Likewise there are a number of mountains of gypsum, which are typically the remnants of a namakier, at the stage where all the halite has dissolved to leave behind a carapace of gypsum residues, which outline the former extent of the salt tongue. Gypsum residues too will ultimately dissolve, as is occurring about the current edges of Kuh-e-Gach (Saveh) in central Iran (Figure 1).

     

    Qom Kuh is one of a number of salt diapirs currently emergent along releasing bends in the Qom-Saveh basin, which are regions that have been pulled apart along major dextral transpressive faults, crossing and defining the margin of the Qom-Saveh Basin in the central plateau of Iran. The other two features discussed in this essay, Kuh-e-Namak (Shurab) and Kuh-e-Gach Saveh are located on the current thrust margin of the basin and so have been cut-off from the mother salt supply, rather than directly above the current salt area (Figure 1). Interestingly, Kuh-e-Namak (Qom) is only ten or so kilometres away from the site of the famous 1956 blowout site at the Alborz 5 well atop the Alborz anticline (Figure 1). With an estimated volume of 5 million barrels of oil escaping at Alborz 5 over 82 days (until the blowout self-bridged), the salt-sealed and salt-cored fractured Qom carbonate (Oligo-Miocene) reservoir remains the feeder to one of the world’s largest blowouts. For comparison, Macondo in the Gulf of Mexico (another salt-related blowout) had an estimated total volume of escaped fluids of around 4.9 million barrels. We shall discuss salt-related blowouts and salt seal integrity in a future blog. For now, we shall focus on what typifies the geology where a namakier flows over the landsurface and then starts to dissolve.

     

    Kuh-Qom, is still located in a supplying salt mass and so is actively flowing or fountaining salt to the surface, while the other two salt features are in varying stages of collapse and dissolution (Figure 1). Qom-Kuh is located on a releasing bend of an offset in the Qom Fault that crosscuts the plunging nose of the Alborz Anticline (Figure 2). The Qom Fault intersects a number of normal faults in the vicinity of Kuh Qom which are exposed to the east of Qom Kuh, where it extends beneath alluvium to the northwest of the diapir (Figures 1, 3). Splays of oblique-reverse and oblique-normal faults offset the nearby asymmetric salt-cored thrust anticline, known as the Alborz Anticline (where the Alborz 5 blowout occurred). These faults are well exposed in the 60m-high ephemeral-stream eroded cliff that defines the eastern flank of Qom Kuh. The adjacent ephemeral-river bench has bevelled into the Upper Red Formation marls at the 920m above sea level (Figure 3b). Faults are clearly seen in the zone separating the diapir from the hill created by the northern limb outcrop of the anticline. The largest of these faults has an irregular trace that was superimposed on a ductile dextral strike-slip shear zone before it offset the river bench some 1.5 m down to the east (Figure 3a). Today the crest of the salt fountain of Qom Kuh rises to 1235 m asl and is some 315 m above the surrounding plateau (Figure 3d).

     

    The Qom Kuh namakier can be divided into two main components: (1) a smooth hemispheric dome or summit with a diameter of 2.5 km and surrounded by, (2) a dissected and dissolving apron of salt allochthonous flowing atop Upper Red Formation and Recent gravels (Figure 3b, d). The margin of this rim is defined by large blocks of salt-buoyed exotics, mostly Eocene volcanics, that have slide down the salt mass to accumulate as a pile of blocks about the salt edge (Figure 3a-c). Steps in the salt topography indicate that the extruded salt overflows a collar of Upper Red Formation up-tilted around the inferred diapiric vent. This raised collar is highest beneath the NE shoulder of Qom Kuh, where it is exposed in windows up to 1100 m asl and lowest in its SW corner (Figure 3; Talbot and Atabi, 2004; Cosgrove et al., 2009). At the Qom-Kuh summit the salt mass is characterised by diffusion karst and variably covered by up to a metre of clay-dominant residuals, along with rare cm-sized clasts of anhydrite and limestone.

    The river cliff truncates the eastern end of a sedimentary bedrock collar (Figure 3a, b). The bench and cliff exposes the apron of allochthonous salt, which is up to 200 m thick and about 0.75 km long to the west of the summit, and some 2 km to the south (Figure 3d). The underlying rhombic vent, through which Qum Kuh extrudes, is interpreted to be a pull-apart between normal faults that hard-link a releasing offset along a regional transpressive strike-slip fault, which trends west–east (Figure 2b). In profiles though the summit, Qum Kuh can be considered as a pile of recumbent fold nappes of gneiss-like and mylonite-like Oligocene and Miocene salts that have extruded from depth and the gravity spread over recent alluvial gravels (Figure 3c).

    An ephemeral stream still flows anticlockwise around most of Qom Kuh, and is still eroding the cliff on the east side of Qom Kuh. Broken pottery lying on the river-bevelled bench to the east, and exposed by wind-deflation of the soils overridden by the southern namakier, suggest that the river eroded the salt of Qom Kuh in historical times (≈ 10,000 years ago according to Talbot and Atabi, 2004). Salts dissolved from the Qom Kuh namakier are now being re-precipitated as Holocene evaporite beds in the adjacent playa low, which is possibly situated atop an active salt withdrawal basin (Figures 2, 3b).

     

    The rhombic transtensional nature of the salt neck to this pull-apart graben structure that allows the ongoing salt supply to the Kuh-Qom salt fountain is seen more clearly in a now inactive rhomboid neck , which now outcrops at Kuh-e-Gach, Saveh, a highly dissolved namakier remnant (Figures 1, 4). Only the nodular gypsum carapace and the underlying reducing brine halo remain to define the position of the former namakier (Figure 4a-d). The leaching of the namakier salt created dense plumes of reducing brines that coloured the underlying redbeds grey, as it converted iron in the redbeds from its ferrous to ferric state (Warren, 2008). Clearly, at the other end of the dissolution spectrum from an active salt fountain, the actual evaporite mass (both halite and gypsum residues) will disappear, as at Kuh-e-Gach (Saveh). Once the reducing brine source is gone, it is also likely the greybeds will transition back into redbeds in this semiarid climate. All in all, not much evaporite evidence of a former salt fountain will remain as mineral salts in the namakier-influenced stratigraphy.  


    As a general rule, the upper portion of an actively flowing salt mound is covered by residual soils, along with a variety of gypsum textures growing in the insoluble components (suffusion karst). Small ephemeral features, including shallow diffusion caves, are typical of the salt glacier surface, but in an active namakier the salt is flowing too fast to preserve them for more than a few hundred to a few thousand years depending on the lateral extent and rate of extrusion of the diapir. The karstic carapace atop the moving and dissolving namakier is dominated by soils, typically composed of a combination of dissolution residues, water-carried deposits and wind transported dust (loess). As the moving salt dissolves, dolines develop in the surface of these soils, they enlarge with time as the soils thicken. These non-halite features remain in place and can even accrete as they are carried downslope by underlying flowing salt evan as it dissolves. Dolines below the surface of these soils enlarge and the soils thicken as the moving salt dissolves in its passage away from the vent. Ultimately, once the supply of mother salt is depleted the at-surface namakier salt dissolves to leave behind a gypsum-encrusted surficial layer, no more than a metre or two thick (Figure 5; Warren, 2008).

     

    Once the supply from the mother salt layer ceases, an at-surface namakier stops its outward expansion and begins to shrink to ultimately subside and disappear. This occurs over time scales of hundreds to thousands of years post-flow, to leave behind a salt-ablation carapace. As the salt dissolves, soils that previously had accumulated about to namakier edge, now touch down over the whole extent of the former salt glacier. Ultimately, karstification and salt core collapse becomes so pervasive all the way across the former extent of the namakier core so that only the edges of the relict plug still show any positive relief (Figure 6). Salt plug ruins and overburden blocks once buoyed by the salt mass are  broken and disturbed, while a halo of salt-buoyed exotic blocks defines the former extent of the salt tongue. Overall, in the local Iranian landscape underlain by salt, inactive and shrunken plugs (as at Shurab - Figure 7) tend to be negative landscape features (depressions). 

     

    Although not leaving behind much in the way of salt mineral remnants, the salt-buoyed transport of exotic blocks, which can be hundreds of metres in diameter, and the remnants of lithologies that were plucked and buoyed by the salt as it made its way to the surface are the best evidence for the former presences of a namakier mass. Such a horizon or layer of exotic blocks can also be accompanied by, and typically is immediately underlain by, a disturbed layer of scattered and rotated overburden blocks. This style of mega-breccia creation is seen in the foundering and collapse blocks of Qom Limestone in the salt-cored breached anticline that defines the Shurab diapir (Figure 7).


    Today, to the south of the Qom Basin, in the Persian (Arabian) Gulf there are numerous salt-cored islands in various stages of namakier dissolution, such as Hormuz, Das and Yas islands (Figure 8). These once fountaining extrusive salt masses are now largely inactive, with only minor at-surface evaporite residues present, mostly as gypcretes and caprock remnants. Highest parts of the diapir-cored islands are typically covered by dissolution breccias that cap the still dissolving salt core below. Halokinetic Precambrian (Hormuz) salt beneath the breccia carapace is pervasively karstified and where relatively shallow, as at Hormuz Island, is covered by suffusion karst and crosscut by tube caves. Active flow of the squeezed salt in the various island cores seems to have ceased sometime in the Miocene, so that terrains of exotic blocks of meta-igneous and dolomite litologies now outline much of the surface expression of former outcropping salt masses, Across the Gulf region this stratiform complex of residues and large salt-buoyed blocks is termed the "Hormuz complex." But what it actually is a series of salt ablation breccias. Today these regions of outcropping Hormuz Complex are heavily eroded and partially covered by Neogene shoal water carbonates and marine-cut platforms, as in Das and Hormuz islands (Figure 8a, b)


    Texturally near-identical Neoproterozoic megabreccias, breccias and breccia trains crop out in the Flinders Ranges of South Australia (Figure 9; Dalgarno and Johnson, 1968; Lemon, 1985). The salt no longer remains in these features, but they are clearly remnants of what were once diapiric structures (Hearon IV et al, 2015). Many breccias in anticlinal cores of the this halokinetic terrane are still located at or near the level of the former mother salt bed (Callanna Beds), with current outcrop patterns largely indicative of the positions of deeper basement faults (Backé et al., 2010). Today, the breccias define the polytectonic remnants of former autochthonous salt pillows or glide planes and inversion structures, rather than true salt-cored diapirs. Some structures, such as the Oratunga and Wirrealpa diapirs, still preserve evidence of their earlier formative extension phase and show subcircular patterns of outcrop and breccia wings at allochthonous stratigraphic positions well above the level of the original mother salt bed. Almost all the transtratal (halokinetic) breccias line up along major regional shears and faults (welds) surrounded by jostling depopods or minibasins, indicating that the mother salt bed was flowing as the basin was under extension. The salt was then remobilised during the inversion phase and the transition from haselgebirge to rauhwacke textures. Lemon (1985, 1988, 2000), Dyson (2004) and Rowan and Vendeville (2006) clearly illustrate synkinematic controls on sedimentary facies and thicknesses adjacent to many of these breccia masses in the Flinders Ranges (Warren, 2015).

    In summary, when a namakier shrinks it does so via at-surface dissolution of the salt mass, much in the same way a glacier retreats as it melts). Likewise, the outer edge of a salt glacier expands and contracts like the outer edge of an ice glacier. When the salt mass retreats it leaves behind a jumbled mass of material it once buoyed, which can include megabreccia blocks tens of metres across, as well as the insolubles of its carapace and insoluble intrasalt layers. These chaotic breccias are similar to a diapiric collapse breccia, but contain a much more polymict assemblage of clasts than an evaporite collapse breccia (after bedded salt) and often show evidence of mechanical reworking of portions of the breccia material by waves or currents. A separate term is probably needed to distinguish them from the more general term diapiric breccia; I call them salt ablation or salt-retreat breccias (Table 1; Warren, 2015).

     

    As we saw in the successive stages of diapir dissolution and retreat exposed in the Qom-Saveh Basin of Central Iran, wherever an active but shrinking salt tongue subcrops or lies beneath a dissolution-derived gypcrete carapace, adjacent larger clasts accumulate by breakup of the soft overburden and by stacking of fragments and rafts formerly held within the now dissolved salt matrix. The resulting namakier breccia is composed of a coarse, unsorted, heterolithic (polymict) rubble supported by a fine-grained calcareous matrix dominated by solution flour. This material can be mixed with fluvial and alluvial material (note the fans fed by alluvial streams that cross the ablation zone in Figure 2a) or marine carbonate debris from times when the sea encroached on the salt tongue (as at Hormuz Island; Figure 8).

    Ongoing dissolution of the namakier leaves behind a rubble moraine at the level of the former salt allochthon. Thus, salt ablation breccias associated with salt cores are jumbles of exotic material carried to the surface by upwelling salt, which is then combined with the disrupted remnants of the any brittle overburden that once covered an earlier now-dissolved salt tongues. The mixing of contemporary sediment (alluvial or marine) with salt-buoyed megabreccia blocks and insolubles in the salt, and the stratiform circum-stem or circum-weld nature of its occurrence, is what distinguishes a salt ablation breccia from a diapiric breccia formed by subsurface salt dissolution and from breccias that result from the dissolution of a salt bed (Table 1; Warren 2015). 

     

    References

    Backé, G., G. Baines, D. Giles, W. Preiss, and A. Alesci, 2010, Basin geometry and salt diapirs in the Flinders Ranges, South Australia: Insights gained from geologically-constrained modelling of potential field data: Marine and Petroleum Geology, v. 27, p. 650-665.

    Cosgrove, J. W., C. J. Talbot, and P. Aftabi, 2009, A train of kink folds in the surficial salt of Qom Kuh, Central Iran: Journal of Structural Geology, v. 31, p. 1212-1222.

    Dalgarno, C. R., and J. E. Johnson, 1968, Diapiric structures and late Precambrian-early Cambrian sedimentation in Flinders ranges, South Australia: American Association Petroleum Geologists, Memoir, v. 8, p. 301 -314.

    Dyson, I. A., 2004, Christmas tree diapirs and the development of hydrocarbon reservoirs; A model from the Adelaide Geosyncline, South Australia: Salt-sediment interactions and hydrocarbon prospectivity: concepts, applications and case studies for the 21st Century. Papers presented at the 24th Annual Gulf Coast Section SEPM Foundation Bob F. Perkins Research Conference, Houston Tx, December 5-8, 2004 (CD publication), p. 133-165.

    Hearon IV, T. E., M. G. Rowan, K. A. Giles, R. A. Kernen, C. E. Gannaway, T. F. Lawton, and J. C. Fiduk, 2015, Allochthonous salt initiation and advance in the northern Flinders and eastern Willouran ranges, South Australia: Using outcrops to test subsurface-based models from the northern Gulf of Mexico: Bulletin American Association Petroleum Geologists, v. 99, p. 293-331.

    Hurford, A. J., H. R. Grunau, and J. Stöcklin, 1984, Fission track dating of an apatite crystal from Hormoz Island, Iran: Journal of Petroleum Geology, v. 7, p. 365-380.

    Kent, P. E., 1987, Island salt plugs in the Middle East and their tectonic implications, in I. Lerche, and J. J. O'Brien, eds., Dynamical geology of salt and related structures, v. 3-37, Academic Press, New York.

    Lemon, N. M., 1985, Physical Modelling of Sedimentation Adjacent to Diapirs and Comparison with Late Precambrian Oratunga Breccia Body in Central Flinders Ranges, South Australia: American Association Petroleum Geologists Bulletin, v. 69, p. 1327 - 1328.

    Lemon, N. M., 1988, Diapir recognition and modelling with examples from the Late Proterozoic Adelaide Geosyncline, Central Flinders Ranges, South Australia: Doctoral thesis, Univesity of Adelaide.

    Lemon, N. M., 2000, A Neoproterozoic fringing stromatolite reef complex, Flinders Ranges, South Australia: Precambrian Research, v. 100, p. 109-120.

    Morley, C. K., B. Kongwung, A. A. Julapour, M. Abdolghafourian, M. Hajian, D. Waples, J. Warren, H. Otterdoom, K. Srisuriyon, and H. Kazemi, 2009, Structural development of a major late Cenozoic basin and transpressional belt in central Iran: The Central Basin in the Qom-Saveh area: Geosphere, v. 5, p. 325-362.

    Rowan, M. G., and B. C. Vendeville, 2006, Foldbelts with early salt withdrawal and diapirism: Physical model and examples from the northern Gulf of Mexico and the Flinders Ranges, Australia: Marine and Petroleum Geology, v. 23, p. 871-891.

    Talbot, C. J., and P. Aftabi, 2004, Geology and models of Qum Kuh central Iran: Journal of Geological Society of London, v. 161, p. 1-14.

    Warren, J. K., 2008, Salt as sediment in the Central European Basin system as seen from a deep time perspective (Chapter 5.1), in R. Littke, ed., Dynamics of complex intracontinental basins: The Central European Basin System, Springer-Verlag, Berlin-Heidelberg, p. 249-276.

    Warren, J. K., 2015, Evaporites: A compendium (ISBN 978-3-319-13511-3) Released August 2015: Berlin, Springer, 1600 p.

     


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    Dead Sea caves carbon cycle Enceladus water on Mars Kara bogaz gol well blowout marine brine DHAB hectorite McMurdo Sound phreatic evaporite natural geohazard deep meteoric potash Density log lapis lazuli Karabogazgol waste storage in salt cavity evaporite-metal association astrakanite cryogenic salt lot's wife intrasalt Patience Lake member gem evaporite dissolution circum-Atlantic Salt Basins 13C enrichment Musley potash gassy salt sodium silicate Muriate of potash Jefferson Island salt mine gas in salt seal capacity MOP supercontinent African rift valley lakes nuclear waste storage carnallitite jadarite CaCl2 brine hydrothermal anhydrite knistersalz Thiotrphic symbionts meta-evaporite organic matter flowing salt lunette Precambrian evaporites brine pan halokinetic sulfur anthropogenic potash GR log stable isotope salting-out anomalous salt zones deep seafloor hypersaline anoxic basin halotolerant Zaragoza High Magadi beds phreatomagmatic explosion capillary zone Badenian Stolz diapir mass die-back ancient climate HYC Pb-Zn saline giant CO2: albedo authigenic silica 18O potash ore Lake Peigneur Catalayud perchlorate gas outburst tachyhydrite Large Igneous Magmatic Province halophile Koeppen Climate salt tectonics Sulphate of potash Mulhouse Basin dissolution collapse doline dihedral angle Calyptogena ponderosa methanotrophic symbionts salt trade endosymbiosis Messinian Europe mirabilite solikamsk 2 rockburst bedded potash zeolite Stebnyk potash Kalush Potash chert ozone depletion recurring slope lines (RSL) climate control on salt Hadley cell: hydrothermal halite halite-hosted cave palygorskite Platform evaporite Seepiophila jonesi subsidence basin retrograde salt Salar de Atacama SO2 phreatic explosion supercritical halite well log interpretation Red Sea antarcticite Sumo Schoenite End-Triassic source rock Weeks Island salt mine brine evolution SedEx allo-suture silicified anhydrite nodules wireline log interpretation Ure Terrace Deep halite geohazard K2O from Gamma Log alkaline lake water in modern-day Mars Warrawoona Group Danakhil Depression, Afar Paleoproterozoic Oxygenation Event anthropogenically enhanced salt dissolution Mega-monsoon epsomite Great Salt Lake NPHI log Koppen climate Lomagundi Event evaporite oil gusher Phaneozoic climate Proterozoic namakier eolian transport sinkhole nacholite Dallol saltpan Neoproterozoic Oxygenation Event salt leakage, dihedral angle, halite, halokinesis, salt flow, MgSO4 depleted causes of glaciation auto-suture hydrothermal karst hydrogen non solar heating salts End-Cretaceous Dead Sea karst collapse trona Crescent potash Belle Isle salt mine nitrogen vadose zone extraterrestrial salt potash Lamellibrachia luymesi cauliflower chert kainitite Neutron Log Ethiopia vanished evaporite dark salt halocarbon snake-skin chert MgSO4 enriched cryogenic spring salts doline sinjarite white smokers salt periphery Archean Ganymede hydrohalite saline clay Mesoproterozoic base metal sedimentary copper Deep seafloor hypersaline anoxic lake methanogenesis collapse doline Lop Nur supercritical phase salt mine Clayton Valley playa: basinwide evaporite hydrological indicator carbon oxygen isotope cross plots silica solubility Noril'sk Nickel blowout sulphur Hadley Cell intersalt Ingebright Lake lithium battery stevensite causes of major extinction events mummifiction methane Atlantis II Deep brine lake edge salt suture Neoproterozoic Mars Lop Nor End-Permian Evaporite-source rock association SOP Stebnik Potash evaporite karst 13C Zabuye Lake venice lithium brine Bathymodiolus childressi potash ore price well logs in evaporites lazurite NaSO4 salts gypsum dune salt ablation breccia bischofite evaporite-hydrocarbon association Dead Sea saltworks freefight lake Boulby Mine vestimentiferan siboglinids hydrothermal potash mine stability extrasalt Turkmenistan Prairie Evaporite Pangaea sepiolite Realmonte potash Five Island salt dome trend crocodile skin chert 18O enrichment Magdalen's Road lithium carbonate seawater evolution Mixing zone H2S LIP DHAL salt karst Hyperarid halogenated hydrocarbon magadiite McArthur River Pb-Zn Prograde salt North Pole Gamma log salt seal Quaternary climate CO2 RHOB MVT deposit sulfate Pilbara black salt Corocoro copper Belle Plain Member Beebe hydrothermal field Hell Kettle Ripon York (Whitehall) Mine Lake Magadi solar concentrator pans Ceres sulphate

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