Salty Matters

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Silica mobility and replaced evaporites: 2 - replaced CaSO4

John Warren - Sunday, July 31, 2016

Postdepositional silicification of sulphate evaporites, that is the precipitation of authigenic silica as a replacement of a CaSO4 host, is the focus of this article, but can be considered a subtopic of the broader styles of silica deposition and silicification that have occurred throughout the geological record from the Precambrian to the Quaternary (Knauth and Epstein 1976; Bustillo 2010). The next article will extend the silica -precipitate discussion back in time across the Proterozoic and into the Archean and consider the influences of atmospheric evolution and seawater chemistry on the styles of silica rocks across deep time. At this point in our discussion, a few relevant geological and mineralogical definitions are needed (see Bustillo, 2010 and Marin-Carbonne et al for more detail). Silica rock is a general term used to define any rock composed mainly of SiO2. In the strict sense, “chert” is used to define a silica rock made primarily of quartz, plus small amounts of opaline minerals, whereas the term “opal” is used to indicate both a mineral and rock. Cherts are sedimentary rocks formed either by direct precipitation from hydrothermal fluids or seawater (known as C-cherts) or by silicification of precursor material (S-cherts). That is, C-cherts are the result of from orthochemical precipitation from seawater (or any Si-rich fluid) and S-cherts are the result of the replacement of a precursor lithology (van den Boorn et al. 2010). This precursor can be evaporitic or volcanogenic sediment (Marin-Carbonne et al., 2014) This article emphasises S-chert examples from the Phanerozoic saline settings, where silica is a secondary phase replacing a pre-exisiting evaporite nodule or crystal. This style of authigenic silica is a common diagenetic constituent in evaporitic carbonates, and occurs in a variety of crystal forms and morphologies (Folk and Pittman 1971; Chowns and Elkins 1974; Knauth 1979; Milliken 1979; Geeslin and Chafetz 1982; Chafetz and Zhang 1998; Scholle and Ulmer-Scholle 2003).

Authigenic silica (S-cherts) can form by: (1) Diagenetic recrystallization of an amorphous silica precursor (Hesse 1989; Knauth 1994); (2) Direct precipitation from aqueous solutions (Mackenzie and Gees 1971; Guidry and Chafetz 2002; Marin et al. 2010); and (3) Direct replacement of pre-existing olcanogenic, carbonate or evaporite host (Hesse 1989; Knauth 1994). Several possible chemical explanations have been suggested to drive the replacement. These include silica precipitation induced by a local decrease in pH that is caused by either biological production of CO2 (Siever 1962), oxidation of sulfide into sulphate (Clayton 1986; Chafetz and Zhang 1998), and mixing of marine and meteoric waters (Knauth 1979).

Types and traits of authigenic silica and cherts

The previous article in this series on silica mobility in evaporitic settings focused on the most mobile (soluble) form of silica known as opal-A or amorphous silica which is defined by a broad peak in XRD determinations (Figure 1). Based on that discussion, it seems there are three main ways modern amorphous silica precipitates; 1) Inorganic precipitate (as in the crusts of the Coorong ephemeral lakes, 2) As a replacement of sodium silicates, such as magadiite (as in alkaline lakes in the African Rift Valley, and 3) Biogenically as in diatom and radiolarian tests in various lakes and the oceans. In all three Opal-A (amorphous opal) is the dominant form of SiO2, but there are other more crystalline forms of sedimentary silica Quaternary sedimentary settings with additional opaline and more quartzose forms. Knauth (1994) classified authigenic silica into, a) 3 types of amorphous opal (opal-A, opal-CT, and opal-C) and, b) 5 types of quartz (granular microcrystalline quartz, megaquartz, length-fast chalcedony, length-slow chalcedony, and zebraic chalcedony). 

Unlike quartz, the opaline minerals are metastable and show different degrees of crystallinity, crystal structure and proportions of water. Jones and Segnit (1971) classified opal minerals into three groups, according to their X-ray diffraction (XRD) patterns (Figure 1a): Opal A (with an XRD pattern that resembles that of amorphous silica), Opal C (which shows four moderately broad peaks that coincide closely with the position of the four most intense peaks of α-cristobalite, plus minor evidence of α-tridymite), and Opal CT (with patterns that show signs of both α-cristobalite and α-tridymite). Opal A can be inorganic, but worldwide is frequently found as siliceous microfossils (diatom frustules, sponge spicules, phytoliths, etc.). Opal C is very rare in sediments. Opal-CT is the most common phase, but its structure can differ owing to its variable water content, the ratio of interlayered cristobalite/tridymite to the amorphous background, and the degree of stacking disorder within the silica framework (Guthrie et al., 1995).


So, amorphous silica is composed of relatively pure SiO2 but with only very local crystallographic order. Amorphous silica includes various kinds of hydrated and dehydrated silica gels, silica glass, siliceous sinter formed in hot springs, and the skeletal materials of silica-secreting organisms. Opal or opaline silica is a solid form of amorphous silica with some included water (Figure 1b). It’s abundant in young cherts, extending back into the Mesozoic. Its geological occurrence is varied it can be by alteration of volcanic ash, precipitation from hot springs, and, volumetrically most significant in the Phanerozoic via precipitation as skeletal material by certain silica-secreting organisms. Opal starts out as what is called opal-A, which shows only a very weak x-ray diffraction pattern, indicating that any crystallographic order is very local. With burial, during the initial stage of diagenesis, opal-A is transformed into opal-CT, which shows a weak x-ray diffraction pattern characteristic of cristobalite). Upon further diagenesis, opal-CT is transformed into crystalline quartz, resulting in chert that consists of an equant mosaic of microquartz crystals. Chalcedony is made up of needles or fibers, often spherulitic, composed of quartz. There’s probably amorphous silica in among the needles, and a variable water content. It is metastable with respect to ordinarily crystalline quartz, but it persists across long time frames; it’s found even in some Paleozoic cherts. Porcellanite is the porous form of chert while silicilyte is a related form that typifies evaporite-associated bacterially-mediated sediment forming a producing reservoir in the South Oman Salt Basin (later blog).

During burial diagenesis, opaline phases age by undergoing successive dissolution-precipitation-recrystallization reactions including the well-known opal A→opal CT→quartz transition (Williams and Crerar, 1985; Williams et al., 1985). These transformations depend mainly on time and temperature, but accelerate in meteoric diagenetic settings, where quartz crystals can form directly, and bypass the opaline silica polymorph phase (Arakel et al., 1989; Bustillo and Alonso-Zarza, 2007). The existence of opal-CT in very young and at-surface rocks (Jones and Renaut, 2007; Jones et al., 1996) shows that time is not necessarily “a cause” in silica diagenesis. According to Bustillo (2010) in continental environments, very rapid silica alteration appears to be related to efficient fluid delivery (i.e., hydrogeology), as much as to time.


When opal-A or opal-CT occur in a sedimentary host, their ageing sets silica free in a dissolved form and so influences the diagenetic evolution of the adjacent carbonates, generally producing silica/carbonate replacements, silica cement, or neoformed silicate clay. Quartz is the last stage of the recrystallization of opals, but can also form directly via replacement or the cementation of voids. Such quartz shows many textures under polarising light. Common quartz can have different crystal sizes and forms crypto-, micro-, meso- or macrocrystalline mosaics. Maliva and Siever (1988) indicated that meso- and macrocrystalline quartz are not produced by ageing but only by direct precipitation during replacement or cementation. Chalcedony is a fibrous-texture quartz made up of several different varieties classified by the orientation of the fibres with respect to the crystal’s c-axis, namely (Figure 2a): Calcedonite (length-fast chalcedony, in which the elongation of the fibres is perpendicular to the crystallographic c-axis), quartzine (length-slow chalcedony, in which the elongation is parallel), lutecite (another type of length-slow chalcedony, in which the fibre axis is inclined by approximately 30°), and helicoidal calcedonite or zebraic chalcedony (which shows a systematic helical twisting of the fibre axes around the crystallographic c-axis). These varieties of chalcedony allow the identification of the environment reigning during the replacement or cementation as acid or non-sulphate (length-fast), or basic or sulphate/magnesium-rich (length-slow) (Figure 2; Folk and Pittman, 1971). The host material, therefore, has geochemical control over the textures of quartz precipitated. Unfortunately, there are exceptions to these rules, and the strict application of these criteria can lead to errors of interpretation.

Moganite is a metastable monoclinic silica polymorph that is structurally similar to quartz (Miehe and Graetsch, 1992). The identification of moganite in the presence of quartz is difficult. It can be detected, however, by detailed XRD analyses with Rietveld refinements, and by other techniques such as Raman and NMR analysis. This mineral is found mixed with quartz in many cherts, preferentially in those that developed in evaporitic environments. However, it can also be produced by the replacement of biogenic carbonates during the interaction of the latter with groundwater (Heaney, 1995). Moganite transforms into quartz, as do the opaline phases, and it probably does so quite readily (Rodgers and Cressey, 2001).

In addition to its replacement style, a number of studies have investigated oxygen isotopic compositions (d18O) in chert to infer climate-driven temperature change through time (Degens and Epstein 1962; Knauth and Epstein 1976; Knauth and Lowe 2003).

 

Silicification of calcium sulphate nodules and isolated crystals

Silicified anhydrite nodules and CaSO4 crystals are widely reported and reliably documented in sediments as old as Paleoproterozoic and as young as Holocene (Table 1). Quartzine and lutecite (aka length-slow chalcedony) typically infill or replace nodules that preserve characteristic cauliflower shapes of the antecedent anhydrite/gypsum nodule (Figure 3; Arbey, 1980; Hesse, 1989). According to Folk and Pittman (1971), rates of nucleation and crystallisation are the primary controls on crystal size and variety of silica precipitating in a void in a dissolving nodule. Rates, in turn, depend on the level of silica saturation or its concentration in the mother brine (Figure 2b, c). According to Keene (1983), precipitation of length-slow quartz is favoured in waters with high SO4 and Mg levels.


High pH levels (alkaline conditions) in the mother solution tend to ionise dissolved silica. Neutral or low pH levels favour silica crystallites made up of combined Si(OH)4 groups. These tend to polymerise into spiral chains at lower pH and higher concentrations. At high concentrations and high pH, the silica precipitates possess a fibrous chalcedonic form reflecting their rapid rates of precipitation. High pH at the precipitation site means silica crystallites also tends to be present in solution as single ionised tetrahedra that attach themselves one by one to the growing surface, so creating fibres of quartz with the c axes oriented parallel to the long axis of the growing fibres (length-slow). Under low pH or in non-sulphate settings the silica is polymerised into spiral silica chains that attach tangentially to the growth surface of the silica gel, with their c-axes parallel to the growing crystal surface and perpendicular to the future direction of the fibres (Figure 2c; length-fast; Folk and Pittman, 1971).

Milliken (1979) summarised the typical petrographic and hand specimen scale features of silica that replaced CaSO4 nodules in Mississippian sediments of southern Kentucky and northern Tennessee (Figure 4). Such nodules typically have knobbly irregular cauliflower-like surfaces, while internal diagnostic textures include: 1) length-slow chalcedony after lathlike evaporites, especially anhydrite; 2) quartzine; and 3) small amounts of lutecite associated either with megaquartz that shows strong undulose extinction, or with euhedral megaquartz (Chowns and Elkins, 1974). The megaquartz often encloses small blebs of residual anhydrite.


Many buried calcium sulphate nodules are silicified in a multistage process that involves both replacement and void filling (West, 1964; Chowns and Elkins, 1974). The process commences about the margins of a nodule (stage 1) with a volume for volume replacement of anhydrite by microcrystalline quartz. It generally ends with the growth of euhedral drusy quartz crystals into a central vug (stage 2 and 3). This mode of replacement exemplifies textural changes as seen from the edge toward the centre of the geode in texture style A in Figure 4. However, as noted by Milliken (1979) this edge inward evolution of the geode or nodule fill is typified by a variety of textural styles, which she denoted a styles A through D.

Stage 1 chalcedony or quartzine mimics or pseudomorphs the felted lath textures of the precursor anhydrite in the outer portion of the nodules in all styles. Anhydrite pseudomorphs occur as radiating or decussate aggregates with a distinctive flow-like pattern indicating a felted anhydrite precursor. Identical decussate and flow textures occur in laths that make up sabkha anhydrite nodules and defines their explosive mode of growth, as well as the typical coalesced nodule texture that, when replaced, ultimately controls the broad-scale “cauliflower” outline of the whole replaced nodule (Figures 3 and 5). And so, as well as silicified lath microtextures seen in thin section, outlines of larger crystals that predated anhydritisation and silicification may be preserved by the nodule margin, these crystal outlines vary from prismatic to bladed. Many silicified nodules still retain the knobbly cauliflower surface morphology of its precursor anhydrite; other nodule edges preserve crystal pseudomorphs with the interfacial outlines of gypsum or anhydrite precursors.


Stage 2 microquartz and quartz fill can assume euhedral faces as they grow into voids created by the dissolution of the nodule. At the same time the quartz may continue to engulf and pseudomorph small areas of residual anhydrite or other less common evaporite salts (e.g. styles A, C, D). Quartz crystals precipitated at this stage are commonly zoned, with more anhydrite inclusions found within the inner region of the pseudomorph. Some quartz crystals are doubly terminated and probably grew via the support of a dissolving meshwork of anhydrite. With the final dissolution of the supporting mesh, these quartz crystals sometimes dropped to the floor of the void to create a geopetal indicator. For example, such highly birefringent anhydrite spots define cauliflower nodules in 2.2 Ga sediments in the Yerrida Basin, Australia (El Tabakh et al., 1999).

Stage 3, the final stage of the void fill is typified by the precipitation of coarse drusy euhedral quartz with no included anhydrite. This coarse quartz resembles coarse vein quartz and often has 18O values indicating temperatures of the mesogenetic or burial realm.

Sometimes the processes of void fill may be arrested to leave a hollow core in the silica-lined geode (Styles A, B, C). The void may be filled later by a different burial stage cement such as baryte, sparry carbonate (e.g. ferroan dolomite or calcite), or even metal sulphides. This is the case with the large (up to 1 m diameter) silicified cauliflower-shaped anhydrite nodules of Proterozoic Malapunyah Formation of the McArthur Basin in Northern Australia where baryte, then metal sulphides and then sparry calcite typify the latter stages of void fill (pers. obs.). Similar fracture-filling baryte characterises the later diagenetic stages of silicified and calcitised anhydrite nodules in the Triassic Bundsandstein redbeds of the Iberian Range of central Spain (Figure 6; Alonso-Zarza et al., 2002). Such geodes are typically excellent indicators of burial cement stratigraphy in a mudstone matrix that otherwise preserves few signs of the evolving pore fluid chemistry. Thus textures and isotopic signatures in a replaced nodule can indicate ongoing diagenesis of the anhydrite nodule that preserves aspects of the shallow active phreatic (eogenetic), the mesogenetic zone with basinal brines and then uplift-related telogenetic fluids.


Internally, cauliflower chert may retain no evidence of former anhydrite lathes mimicked in chalcedony, but can be filled with various styles of coarser-grained megaquartz. The resulting nodules still retain the outline of the precursor evaporite nodule (Figure 3). Work on diagenetic timing of numerous silicified CaSO4 nodules (e.g. Milliken, 1979; Geeslin and Chafetz, 1982; Gao and Land, 1991; Ulmer-Scholle and Scholle, 1994) shows that most silica replacement begins with shallow burial, either in the zone of active phreatic flow or in the upper portion of the zone of compactional flow (probably at depths of less than 500-1000 m). Early silica replacement in the zone of active phreatic flow is indicated by a lack of compressional flattening of the nodule, by the preservation of delicate surface ornamentation and the preservation of compactional drapes around replaced nodules. If replacement of an anhydrite nodule occurs later in the burial cycle, the anhydrite nodule has by then become flattened or sluggy and no longer retain a rugose surface. The result can be a series of “cucumbers” rather than “cauliflowers.”

Milliken’s (1979) isotopic evidence implies much silica replacement in the nodules she studied was relatively early in the burial cycle at temperatures that were < 40°C. Silica was supplied by through flushing pore fluids with compositions ranging from seawater to mixed meteoric-seawater. Of course, nodule replacement by silica or calcite does not have to happen on the way down in the burial cycle; it may also happen during uplift back into the telogenetic realm, where the strata have once again entered the zone of active phreatic flow (Figure 7).

 

Until the turn of the century, there were no documented examples of the process of evaporite replacement by quartz in Quaternary sediments. Now, autochthonous, doubly-terminated, euhedral megaquartz crystals have been observed infilling voids in a gypsum- and anhydrite-bearing Pleistocene sabkha dolomite sequence in the Arabian Gulf, as well as forming overgrowths on detrital quartz grains (Chafetz and Zhang, 1998). These siliceous sabkha precipitates are forming within metres of the present sediment surface with a silica source that is probably recycled biogenic material. Individual quartz crystals attain lengths of 1 mm. Many quartz crystals faces preserve impressions of dolomite rhombs or they partly, or entirely, engulf dolomite rhombohedra. This process of replacement is a response to changing fluid chemistry tied early phreatic burial, to see the full suite of silica replacement textures and the variations in the timing of the replacement means one must study ancient evaporite sequences (Table 1).

  

Overall, the texture of silica infill or replacement in a CaSO4 nodule is dependent on the rate of sulphate dissolution, the timing of silica precipitation and the rate of silica supply. Some nodules are dominated by the early lit-par-lit replacement textures (styles A and C in Figure 4), others have textures indicating silica cement (aligned megaquartz)growing into an open phreatic void left after the complete dissolution of the CaSO4. Such nodules may still retain a hollow centre where the anhydrite once resided (Figure 8). When a silica-filled geode did not start to accumulate silica until after all the CaSO4 dissolved, the primary evidence for an evaporite precursor comes from the shape of the replaced nodule and its stratigraphic position within the evaporitic depositional sequence, e.g. beneath an erosional surface that defines the top of the capillary zone.

  

Not all the anhydrite nodules, now replaced by silica, were syndepositional. Maliva (1987) showed that nodular anhydrite parent, now indicated by quartz geodes in the Sanders Group of Indiana, first precipitated in the subsurface, while its surrounding matrix of normal-marine Sanders Group sediment was still unlithified (Figure 9). Anhydrite nodules formed in the subsurface during early burial as hypersaline reflux brines sank into the normal-marine limestones of the Ramp Creek and Harrodsburg Formations. Silica subsequently replaced the anhydrite nodules. These geodes are almost invariably associated with the development of reflux dolomite.

Similarly, not all silica-replacing anhydrite in a particular region need come from the same source or be emplaced by the same set of processes. Silicified nodules within middle-upper Campanian (Cretaceous) carbonate sediments from the Lafio and Tubilla del Agua sections of the Basque-Cantabrian Basin, northern Spain preserve cauliflower morphologies, together with anhydrite laths enclosed in megaquartz crystals and spherulitic fibrous quartz (quartzine-lutecite). All this shows that they formed by ongoing silica replacement of nodular anhydrite (Figure 10; Gómez-Alday et al., 2002). Anhydrite nodules at Lafio were produced by the percolation of saline marine brines, during a period corresponding to a depositional hiatus. They have d34S and d18O mean values of +18.8‰ and +13.6‰ respectively, consistent with Upper Cretaceous seawater sulphate values. Higher d34S and d18O (mean values of + 21.2‰ and 21.8‰, respectively) characterise nodules in the Tubilla del Agua section and are interpreted as indicating a partial bacterial sulphate reduction process in a more restricted marine environment (Figure 10a). Later calcite replacement and precipitation of geode-filling calcite in the siliceous nodules occurred in both sections, with d13C and d18O values indicating the participation of meteoric waters in both regions (Figure 10b). The synsedimentary activity of the Penacerrada diapir (Kueper salt - Triassic), which lies close to the Lafio section, played a significant role in driving the local shallowing of the basin and in the formation of the silica in the anhydrite nodules. In contrast, eustatic shallowing of the inner marine series in the Tubilla del Agua section led to the generation of morphologically similar quartz geodes, but from waters not influenced by brines derived from the groundwater halo of a diapir.


So far the various papers we have discussed relate the onset of silicification to active phreatic hydrologies (brine reflux or meteoric) typically in evaporites in host rocks that are shallow, either in the early stages of burial or later in the uplift realm. In contrast in a paper discussing silicification of sulphate nodules in Permian (Guadalupian) back-reef carbonates of the Delaware Basin, Ulmer-Scholle et al., 1993, conclude these nodules were silicified in the Mesogenetic realm. Replacement occurred at temperatures of 60-90°C at the same time as hydrocarbons were moving with basinal brines through the adjacent porous matrix (Figure 11). Silicification of these evaporite nodules proceeded from the exterior to the interior of the nodules. The fluid inclusions in the replacive megaquartz are primary, and many contain both hydrocarbons and water. In this setting it seems evaporite silicification was coeval with or slightly postdated hydrocarbon migration and the silica was likely sourced by dissolution of siliciclastics in nearby back-reef units.
 
 

Birnbaum and Wireman (1985) argued that bacterial degradation of organic matter must be important in forming silica precipitates in most evaporites. They demonstrated, through experiment, the strong influence of bacterial sulphate reduction on silica solubility. The ability of sulphate-reducing bacteria to remove silica from solution is related to local changes in pH and hydrogen bonding within amorphous silica, followed by polymerization to higher weight molecules. During silica replacement of sulphate evaporites at relatively shallow burial depths, the pore fluid becomes depleted in dissolved sulphate as it is reduced to H2S by the action of anaerobic sulphate-reducing bacteria, which metabolise sulphate from an anhydrite or gypsum substrate. Where this selective dissolution of the sulphate occurs in the presence of amorphous silica, the reaction is accompanied by the precipitation of silica. Hence the microscale mimicry of the lath outlines in the outer parts of many replaced nodules. According to Birnbaum and Wireman, it reflects bacterially-mediated silica replacement of nodules in relatively shallow burial settings where bacteria flourish.

In summary, in terms of processes and diagenetic settings associated with Phanerozoic evaporite silicification it seems abiological processes, including thermochemical sulphate reduction and hydrocarbon migration, are more important at greater burial depths where bacteria no longer survive. Providing matrix permeability is retained, silica replacement can continue into the thermobaric stage and if the sulphate nodule survives mesogenetic replacement can even persist into exhumation. Replacement under a thermobaric regime is frequently indicated by the preservation of hydrocarbon inclusions in the infilling silica cement. Both BSR and TSR will be discussed further in the next blog article, dealing with silicification associated with ancient evaporites, but with more emphasis on possible hydrochemical contrasts between the Precambrian and Phanerozoic subsurface waters.

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Maliva, R. G., 1987, Quartz geodes; early diagenetic silicified anhydrite nodules related to dolomitization: Journal of Sedimentary Petrology, v. 57, p. 1054-1059.

Marin, J., M. Chaussidon, and F. Robert, 2010, Microscale oxygen isotope variations in 1.9 Ga Gunflint cherts: assessments of diagenesis effects and implications for oceanic paleotemperature reconstructions: Geochimica et Cosmochimica Acta, v. 74, p. 1161030.

 

Marin-Carbonne, J., F. Robert, and M. Chaussidon, 2014, The silicon and oxygen isotope compositions of Precambrian cherts: A record of oceanic paleo-temperatures?: Precambrian Research, v. 247, p. 223-234.

 

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Nagy, Z. R., I. D. Somerville, J. M. Gregg, S. P. Becker, and K. L. Shelton, 2005, Lower Carboniferous peritidal carbonates and associated evaporites adjacent to the Leinster Massif, southeast Irish Midlands: Geological Journal, v. 40, p. 173-192.

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Silica mobility and replaced evaporites: 1 - Alkaline lakes

John Warren - Saturday, July 02, 2016

 

In this series of blog articles, I plan to look at silica mobility, along with characteristic marine and nonmarine hydrogeochemistries over time, and how these parameters control chert and quartz precipitates and replacements in hypersaline settings. First, I will do this in modern surface and nearsurface settings in the marine and nonmarine realms, ending this first article with a focus on silica distribution and its precipitates in sulphate-depleted saline alkaline lacustrine sediments. Next we shall attempt a synthesis of controls on silica mobility and precipitation in the sulphate-enriched hypersaline marine surface, subsurface and burial realms, as well as defining relevant atmospheric and seawater chemistry changes across deep geologic time. And finally, we shall look at silica mobilisation in subsurface hydrothermal saline settings. The context for this discussion initially comes from the utility of recognising various silicified structures (including chert nodules) as indicative of typical marine (smooth-walled nodules), sulphate evaporite-enriched (cauliflower chert nodules) or sulphate-depleted alkaline lake (Magadi or crocodile skin chert) deposits (Figure 1).


 

Silica geochemistry

Modern river water typically contains less than 15 ppm dissolved silica, shallow meteoric groundwater typically has 10-50 ppm dissolved silica, while the modern ocean has between 0.5 and 10 ppm dissolved silica (Bridge and Demicco, 2008). Silica concentration is lowest in the ocean’s surface layer (less than 1 mg/l), and relatively constant at around 10 ppm below the thermocline. Deep saline sodium chloride and calcium chloride groundwaters contain comparatively little dissolved silica (30-80 ppm), relative to their total ionic content, although some saline subsurface waters in shale pore waters that also entrain dissolved organic acids can contain up to 330 ppm dissolved silica. Some of the highest values present in significant volumes of surface water are found in saline alkaline lakes with elevated pH levels. These waters can contain more than 1000 ppm silica in solution (Figure 3). One of the highest known natural water values, some 3970 mg/l of dissolved, is from a cold water spring known as Aqua de Rey, near the town of Mt. Shasta, California. The spring has a temperature around 54°C ana pH around 11.6.


At normal environmental pH, the dissolution-precipitation reaction of quartz,

Si02(s)quartz + H20(1) <-> H4SiO40(aq)

produces non-ionized silicic acid (H4SiO40). Because quartz is not very soluble at 25 °C, this reaction puts only ≈6 ppm silicic acid into solution. Therefore, most of the silicic acid in river water and groundwater is considered to come from the incongruent dissolution of silicate minerals, such as feldspars, during weathering. Non-crystalline amorphous silica gels are considerably more soluble, putting up to 120 ppm silicic acid into solution across the normal pH range (Figure 2a). The solubility of quartz is significantly affected by an increase in temperature, and at 300°C, approximately 600 ppm silicic acid is dissolved in groundwater with normal pH (Figure 2b; Verma, 2000; Fleming and Crerar, 1982). Siliceous sinter precipitates where such hot waters rise to the surface in hot springs and then cools, as at Mt Shasta. In the subsurface, cooling hydrothermal waters drive considerable silica mineral replacement and other cements associated with some types of epithermal and halokinetic ore-deposits (later blog). Rising pH also significantly affects the solubility of quartz, and this mechanism helps explain the elevated silica levels in the waters of many alkaline lakes (Figures 2a, 3a). At pH > 9 (at 25 °C) silicic acid dissociates:

H4Si04(aq) <-> H+(aq) + H3Si04(aq)

With the elevated pH of alkaline water, this reaction is driven to the right, and the solubility of both quartz and amorphous silica is greatly enhanced (Figure 1a). Saline alkaline lake waters with elevated pH, as in Lake Magadi in the African Rift Valley and the Alkali Valley playa brines of the south-west USA, consistently contain more than 1000 ppm dissolved silica as H3Si04(aq) (Figure 3a).


 

Modern siliceous sediments

Modern siliceous sediments accumulate as biogenic marine oozes, biogenic freshwater-lake deposits, chemical precipitates in alkaline lakes, chemical precipitates in soils (silcrete), and chemical precipitates around subaqueous and subaerial hot springs. Significant volumes of dissolved silica occur in the waters of saline alkaline lakes in the African Rift Valley (e.g. Lake Magadi) and a number of Basin and Range playa lakes in Oregon and California (e.g. Lake Abert, Oregon and Alkali Valley playa in California). Inflows in both regions are leaching highly labile volcanics (Figure 3).

Volumetrically, the most significant accumulations of modern siliceous sediments worldwide are constituted by seafloor deposits dominated by opal-A skeletons of planktonic diatoms. Today, diatoms (bacillariophytal algae with siliceous tests) scavenge virtually all of the silica in fresh to somewhat saline surface waters of the continental lakes and most significantly in open ocean waters of the marine realm (Figure 4). Diatoms arose in the Mesozoic, but have become particularly abundant over the past 30 million years, and their remains now dominate silica deposits of the ocean floor and the biogenic siliceous component of the sediment in many less-saline lakes and playa inflows (Knauth, 2003; Katz et al., 2004). Today, the oceans are everywhere undersaturated with amorphous silica, so diatoms build their shells despite thermodynamics of the dissolved content, so that some 90% of dead diatoms’ tests dissolve before they finally settle to the sea floor (Bridge and Demicco, 2008). However, appreciable thicknesses of diatom oozes can accumulate on the sea floor beneath regions with high productivity of diatoms, namely polar areas and areas of oceanic upwelling where there is a high flux of sinking diatoms, such as off the California coast. Other silica-secreting, single-celled eukaryotic plankton include the heterotrophic radiolarians and silicoflagellates. Radiolarian oozes are common beneath equatorial zones of oceanic upwelling.


 

Pore waters of marine siliceous oozes remain undersaturated with respect to amorphous silica for some depth below the sediment surface, and diagenetic reorganisation of silica is common in deep-sea siliceous sediments. This involves the slow conversion of opal-A to opal-CT, and eventually to microcrystalline quartz, via a complicated series of pathways involving quartz cementing and replacing the oozes. In turn, this leads to the diagenetic precipitation of significant volumes of smooth-walled chert nodules (Figure 1; Hesse, 1989). Worldwide the sampling of oceanic sediment by the Deep Sea Drilling Project (DSDP) has recovered well-developed chert in cemented layers within otherwise still unconsolidated Eocene siliceous oozes. The oldest sedimentary opal-A preserved in DSDP cores is Cretaceous in age, beyond that, cherts are made up of recrystallized quartz.

Dissolved silica and biogenic silica cycling in mesohaline lakes

Diatoms are a significant photosynthetic group driving the accumulation of amorphous silica in the bottom sediments of many freshwater temperate and saline lakes. When salinity and nutrient levels are appropriate, the diatom population in these lakes characteristically undergo explosive population growth. In the lit water column of temperate relatively-freshwater lakes, diatom blooms typically occur in the Spring and Fall of the year. This is when the water column in temperate lakes turns over, allowing nutrient-rich bottom waters to reach the near-surface lit zone at times when the surface water temperature is high enough to support diatom growth. In more saline lakes, diatom blooms occur whenever meromictic lake waters freshen to salinities appropriate for diatom growth. This leads to pulses of biogenic silica accumulating in the laminar distal (profundal) bottom sediments of these saline stratified lakes. As in the oceans, most lake columns in systems with periodic diatom blooms are undersaturated with amorphous silica, so most lacustrine diatom tests dissolve as they sink, and continue to dissolve in the bottom sediment. Closer to the lake shore, the greater volume of silica in the lake bottom sediments tends to come from detrital components washed into the lake as siliciclastic sand and mud.

Chert deposits through time preserve a record of secular change in the oceanic silica cycle. The evolutionary radiation of silica-secreting organisms resulted in a transition from abiological silica deposition, characteristic of the Archean and Proterozoic aeons, to the predominantly biologically-controlled silica deposition of the Phanerozoic (Maliva et al., 2005).

Biogenic lacustrine silica

Diatoms also flourish in the fresher water inflow areas of many salt lakes and playas. They are commonplace primary producers in mesohaline and lower penesaline environments with populations expanding across the lake when salinities are suitable. At less favourable time, healthy diatoms are restricted to refugia areas of fresher water springs and ponds. Their buoyant cells, often augmented by chitin threads or colonial adaptions, enable them to keep within the photic zone better than many other halotolerant algae. Some varieties of benthic diatoms live in lake brines with long-term salinities around 120 ‰, while the upper limit for diatom growth is around 180 ‰ (Clavero et al. 2000; Cook and Coleman 2007; Warren, 2016). The most halotolerant diatom taxa in the saltern ponds of Guerrero Negro are; Amphora subacutiuscula, Nitzschia fusiformis (both Amphora taxa), and Entomoneis sp.; all grow well in salinities ranging from 5 to 150 ‰. Three strains of the diatom Pleurosigma strigosum were unable to grow in salinities of less than 50 ‰ and so are true halophilic alga. A similar assemblage of Amphora sp., along with Cocconeis sp. and Nitzschia sp. dominate the high salinity (150 ‰) saltern ponds near Dry Creek, Adelaide, Australia (Cook and Coleman 2007).


Most mesohaline diatom species flourish at times of freshened surface lake waters or in and about perennial seepage and dissolution ponds (posas) about the edges of some salars, where they can be a major component in some lacustrine stromatolites (Figure 5). Several benthic diatom species are conspicuous in building diatomaceous stromatolites in these freshened (refugia) regions of the saline playa system, for example, Mastogloia sp., Nitzschia sp., Amphora sp., Diploneis sp. They function in a manner analogous to that of cyanobacteria in that they produce extracellular gel, are motile, phototropic, can trap and bind sediment, and create surface irregularities in the biolaminate mat (Winsborough et al., 1994).

Many Quaternary saline lakes have experienced significant fluctuations in water level and salinity across their millennial-scale sedimentary histories. For example, some 10,000 years ago Lake Magadi water depths were hundreds of metres above the present-day water levels and the diatomaceous High Magadi Beds (mostly laminites) were deposited. At that time most of the silica in the mesohaline stratified lake resided in sodium silicates deposited as laminites over most of the profundal lake floor. Diatoms flourished in the fresher waters inflow areas tied to deltaic sediments. Similar diatom rich-zones typify the fresher-water inflow areas of the Jordan River where it flows into the Northern end of the Dead Sea (Garber, 1980).

Silica in alkaline surface waters

It seems diatoms are efficient biogenic vectors for dissolved silica removal, not just in the oceans but also in many mesohaline lake settings. Significant diatom populations occur also in many saline lakes, even some hypersaline alkaline ones. They are important sediment contributors in Lake Manyara (Stoffers and Holdship, 1975), Lake Kivu (Degens et al., 1972), Abert Lake (Phillips and Van Denburgh, 1971), and perhaps also in the Great Salt Lake, Utah (Baxter et al., 2005). Silica levels in the water column of most of the lakesare typically very low (Hahl and Handy, 1969). So, in most mesohaline saline lakes the main biogenic form of silica is as amorphous in diatom skeletons, which tend to periodically accumulate in the bottom sediments with pore waters at the lower end of the hypersaline lacustrine salinity range (Warren, 2016, Chapter 9).

Coorong cherts

However, in some mesohaline to penesaline alkaline lakes, such as those of the Coorong region in South Australia, the ability of a diatom test to survive early burial is likely low. The siliceous frustules tend to dissolve and re-precipitate as amorphous inorganic silica. Abiogenic opal-CT precipitates are commonplace in evaporitic carbonate crusts of a number of Coorong Lakes about the edges of alkaline marginal-marine lagoons and lakes of the Coorong District of South Australia, especially those alkaline lakes containing magnesite or hydromagnesite (Peterson and von der Borch, 1965; Warren, 1988, 1990). Silica precipitation happens within the sediment, or just at the sediment surface, as opal gels (colloids) that can contribute up to 6% by weight of the high-magnesium carbonate sediment in the surface crusts of ephemeral saline lakes such as Milne Lake. Subsequent desiccation of the gel (locally known as yoghurt-textured mud) and consequent cracking creates hardened discs and plates of silica-impregnated mud, about 1 cm thick and 10 cm in diameter. About lake edges, these sediment discs and plates tend to erode into intraclast breccias that coat the uppermost massive unit as crust zones. In some lakes these crust breccia are located immediately edgeward of well-developed hemispherical stromatolites (Figure 6a and b; Warren, 1990).

Under seasonal high-pH conditions, the silica source in a Coorong ephemeral lake and its surrounds dissolves, it then re-precipitates as opal-CT as fresher subsurface groundwater with a lower pH seeps into the lake edges, and mixes with the surface lake brines. Measured pH in Milne Lake surface brines is ≈ 9.5 to 10.2. Silica precipitation tends to occur mostly in the periodically subaerial lake edges during times of incipient lake drying and shrinkage, prior to complete lake desiccation (that is abiogenic silica tends to precipitate in the late spring to early summer of the Southern Hemisphere). The initial silica phase impregnating the lake edge carbonate mud is opal-A. Peterson and von der Borch (1965) argued the likely source of the inorganically precipitated amorphous silica was the dissolution of detrital quartz (sand and silt), which is a common detrital component in the early stages of Holocene lake fill. These older Holocene units are now exposed about some lake edges (Warren, 1998, 1990).

However, diatoms do still thrive in many ephemeral Coorong lakes when surface waters have the appropriate levels of salinity and nutrients. They retreat to refugia about fresh water springs and seeps as the lake dries. Even tough common in plankton populations. intact diatom microfossils (siliceous frustules) have not been recognised in most cores from the same lakes. For example, diatom remains are not present in sapropelic muds in North Stromatolite Lake, a modern hydromagnesite-aragonite lake (Warren, 1990; McKirdy et al., 2010). But, within the organic geochemical constituents of the same cores, there are unusual T-shaped, C20 and C25 highly-branched isoprenoids, which are prominent among the aliphatic hydrocarbons in the extracted organic matter (Figure 7; Hayball et al., 1991). These unusual organic components were later recognised as bacillariophycean algal biomarkers (molecular fossils: McKirdy et al., 1995).


Diatoms are not organic-walled, and the silica of their frustules is highly susceptible to dissolution in present-day alkaline pore waters of this and other Coorong lakes. Hence, soon after burial, their cellular organic matter is destined to become part of the amorphous component of the kerogen (Barker, 1992). It is likely that the direct physical evidence for diatoms (viz. their siliceous frustules) is largely dissolved as waters become alkaline in many mesohaline Coorong ephemeral lakes, so that only biomarkers for a diatom source of the inorganically precipitated silica may remain.

Hence, the ultimate source of the inorganic siliceous carbonate breccia that defines the periodically subaerial edges of many ephemeral hypersaline Coorong lakes is likely from readily dissolved amorphous silica of diatom tests, not the much less soluble quartz, which was postulated as the likely silica source by Peterson and von der Borch (1995). Siliceous breccia zones in the edges of the ephemeral Coorong lakes are intimately tied to characteristic tepee expansion features known as extrusion tepees (Figure 6b). These expansion structures in cemented carbonate crusts are related to the desiccation/cementation of precursor gels washed into fractures beneath mobile sheets of colloid  muds ("yoghurts") that wash about the seasonally shrinking lake edge during the late spring to early summer (Kendall and Warren, 1987). No chert nodules are known to occur in the various Coorong Lakes, only siliceous carbonate mud layers and clasts associated with "extrusion" tepees.


Lake Magadi chert

Hypersaline chert is present as nodules, as well as siliceous breccia layers hosted in Pleistocene sediment that crops out landward of the current lake strandline. Precursors to these modern cherts are thought to initially deposit as late Pleistocene sodium silicates across significant portions of the Lake Magadi basin (Figure 8; Eugster, 1969). Today, these cherty precipitates comprise compact, well-indurated layers and nodules in what are otherwise unconsolidated lake sediments, rich in volcanic debris, sodium silicates and diatoms. The chert shows characteristic reticulate shrinkage cracks on the nodule surface giving it the name crocodile-skin or snake-skin chert or Magadi-style chert (Figure 1). The conversion to chert from its sodium silicate precursor is accompanied by many other enigmatic  textural and structural features such as large desiccation polygons, buckling, reticulation, extrusion, casts of mud-cracks and calcite cements.


Trenching in the regions landward of the current Lake Magadi strandline shows chert-rich zones laterally grade into sediments containing subsurface, still unconsolidated layers of sodium silicate gels dominated by the hydrous sodium silicate magadiite (NaSi7013(0H)3-3H20), with lesser amounts of kenyaite NaSi11O20.5(OH)3.H20 and makatite - NaSi2O3(OH)3.H2O (Figure 9a). Magadiite is a highly siliceous phase, running ≈78% SiO2 by weight (Table 1). When fresh, magadiite is white, soft, putty-like and readily deformable, but it dehydrates rapidly on exposure to air to harden irreversibly into fine-grained cherty aggregate. To date, magadiite has been found only in Quaternary alkaline lacustrine settings. In addition to Lake Magadi, Quaternary magadiite occurs in Lake Bogoria and Lake Kitagata, in Lake Chad in western central Africa, in the flats of Alkali Valley playa in Wyoming, and Trinity County in California (Sebag et al., 2001; Ma et al., 2011).


 

The solubility-equilibrium trends for silica and amorphous silica are similar, with a marked increase in solubility occurring in more alkaline conditions (pH>9; Figures 1a, 9b; Dietzel and Leftofsky-Papst, 2002). In contrast, SiO2 contents at equilibrium with magadiite show a minimum value at a pH around 8.5 and follow a different dissolution pattern to silica. At low pH the concentration of silica in solution increases, as it also does in alkaline solutions at the other end of the pH spectrum. Thus, the concentration of silica in a solution saturated with respect to magadiite, at constant Na content, is lowest in neutral to slightly alkaline solutions. Below pH 5.9, which is the intersection point of the magadiite and amorphous silica curves, magadiite exhibits a higher solubility than amorphous silica. Thus, at pH < 5.9 magadiite will dissolve, while amorphous silica precipitates (Figure 9; Dietzel and Leftofsky-Papst, 2002).

Conditions associated with the precipitation of magadiite from lake brines in Lake Chad, and probably most other soda lake occurrences, including Lake Magadi, require fluctuations in alkalinity or mixing interfaces between alkaline and less alkaline groundwaters (Figure 9b-d). Sebag et al. (2001) list the following conditions as typical of most modern magadiite occurrences; 1) Elevated alkalinity, typically in the lake dry season (pH >9) allow dissolution of silica, followed by lowering of alkalinity in the wet season driving precipitation of silica (Figure 9b, c), 2) High concentrations of dissolved silica (up to 2700 ppm), 3) Incorporation of sodium ions into the silica lattice that precipitates at the time of silica supersaturation (Figure 9d). Depending on the concentration of Na and Si in the brine at the time of precipitation, various sodium silicate minerals will precipitate (Figure 9).

Two general pathways have been proposed to explain the formation of magadiite in silica-rich sodium carbonate brines: a decrease in pH and evaporative concentration. Magadiite can precipitates when dilute inflow waters flow across a dense, sodium carbonate brine layer rich in dissolved silica interface mixing lowers the pH at the chemocline/halocline. In Lake Chad, and in some American examples in Califonia and Oregon, magadiite may have also precipitated by evaporative concentration or by capillary evaporation of saline, alkaline brines at a shallow subsurface water table. Other inferred mechanisms for sodium silicate precipitation include: 1) subsurface mixing of dilute and saline, alkaline groundwaters, 2) a reduction in pH of an alkaline brine resulting from an influx of biogenic or geothermally sourced CO2, and 3) precipitation from interstitial brines. Different sodium silicate minerals may form according to the concentrations of Na and SiO2 in the brine.

Magadiite (sodium silicate) layers and nodules in Lake Magadi weather into cherts and cherty breccia layers and so define Magadi-style cherts, with a characteristic reticulate, cracked or “crocodile-skin” surface created by shrinkage during the transformation from sodium silicate gel to chert nodule (Figure 1; Schubel and Simonson, 1990). In places, the Magadi chert layers preserve laminae of the original sodium silicate precursor. Conversion of magadiite to bedded and nodular chert is thought to take place close to the sediment surface and be related either to 1) the mobilisation and flushing of sodium by dilute waters in these shallow environments or, 2) to spontaneous conversion to chert in slightly deeper brine-saturated zones Intermediate diagenetic products, including kenyaite, amorphous silica and moganite, may form during the transformation to chert (see inset; Icole and Perinet, 1984; Sheppard and Gude, 1986). Both magadiite and the associated cherts have a distinctive trace element signature, unlike most other cherts (Kerrich et al., 2002).


 

Eugster et al., (1967) proposed that magadiite of the Lake Magadi High Beds was precipitated in the Late Pleistocene water column by diluting silica-rich, sodium carbonate lake brines with fresher waters at the lake chemocline or mid column interface. Mixing lowered the pH, and although the pH change may have been as little as 0.5, a decrease in pH from 10.3 to 9.8 lowers amorphous silica saturation by more than 500 ppm (Figure 2a). Silica solubility changes very little when pH varies below a maximum of 8. Highly alkaline sodium carbonate waters containing abundant SiO2 readily form in the Magadi rift valley via weathering and rapid subsurface hydrolysis of labile volcanic materials. Biogenically produced CO2 can also reduce the pH of the brine and drive magadiite precipitation (Eugster, 1969). Hay (1968) also suggested that simple evaporative concentration of the brine would lead to magadiite precipitation. Independent of freshwater flushing and pH changes, magadiite decomposes thermally in the lab into quartz and calcite at temperatures of 500-700°C (Lagaly et al., 1975).

In the perimeter sediments of Lake Magadi, Eugster (1969) described what he considered to be an impressive syndepositional result of this transformation of sodium silicates to chert, namely shrinkage megapolygons up to 50 m across in a bedded chert host, with bounding upturned chert ridges up to 2 m high. Historically, the megapolygons, extrusion tepees, convolute folds and other syntransformation features were interpreted as recording the shrinkage-induced flow and collapse of the sediments hosting the magadiite gels, as they lost sodium, dehydrated and shrank.

Subsequent work on the same megapolygonal structures by Behr and Röhricht (2000) concluded the megapolygons are not a response to mineralogical transformation, rather they are part of a suite of prelithification seismite structures in soft, siliceous lake sediment of the precursor Lake Magadi. That is, the chert megapolygons are a soft sediment response to intense deformation and local-scale diapirism, as are the numerous pillow-chert mounds, chert extrusives along dykes and fault ramps, horizontal liquefaction slides with breccias, slumps, petees, flows and shear-structures in the magadiite beds (now all preserved in chert at outcrop).

Collapse, liquefaction and extrusion of the pre-lithified siliceous matrix were caused by seismotectonic rift activity in the lake basin, and it activated fault scarplets and large-scale dyke systems. Seismic activity led to liquefaction and other earthquake-induced intrasediment deformation, especially along fault ramps and on tilted blocks. The textures all indicate the chert megapolygons are a form of seismite and do not mean volume changes in the transition from magadiite to chert. After liquefaction and extrusion, the exposed magadiite material solidified via spontaneous crystallisation to chert in an environment that was characterised by highly variable pH and salinity.

So, since the pioneering work of Eugster and others in the 1960s, three sets of diagenetic processes are now thought to be responsible for driving the conversion of magadiite to Magadi-style (crocodile-skin) chert in Lake Magadi and other soda lakes:

(1) Leaching of sodium by dilute surface runoff during weathering of the High Magadi beds, as evidenced by tracing unweathered beds into outcrop and summarised in the chemical transformation (Eugster, 1969; see inset);

NaSi7O13(OH)3.H2O + H+ —> 7SiO2 + Na+ + 5H2O

(2) Spontaneous release of sodium driving the conversion of magadiite to chert, whatever the nature of through-flushing solutions and environments (see inset). In this process sodium is expelled even in the presence of brines; it does not require the fresh water needed for process 1, and perhaps better explains the occurrence of calcite-filled trona casts in cherts and the presence of chert nodules in unweathered magadiite horizons in Lake Magadi (Hay, 1968, 1970).

(3) To this inorganic perspective on the transformation to form chert, Behr (2002) and Behr and Röhricht (2000) added a biological one. Based on field observations and microbiological studies of the cherts in Lake Magadi region, they argue that inorganic cherts are rare at the type locality of Magadi-style cherts. Rather, as inferred in many modern bacterial dolomites, the cherts at Lake Magadi may have been precipitated as amorphous silica via microbial processes and may not have had a sodium silicate precursor. They go on to note that most of the cherts in the Magadi depression are older than the High Magadi Beds and perhaps developed from flat-topped calcareous bioherms of Pleurocapsa, Gloecocapsa, and other coccoid cyanobacteria, along with thinly bedded filamentous microbial mats, stromatolites, bacterial slimes, diatoms, Dascladiacea colonies and other organic matter accumulations. Silicification occurred from a microbially mediated silicasol, via opal-A to opal-C, with final recrystallisation to a chert of quartzine composition. They conclude that metabolic processes of cyanobacteria controlled the pH of the brine and strongly influenced dissolution-precipitation mechanism that created the chert (Figure 1a). Today the debate as to inorganic versus organic origin of cherts in Lake Magadi continues, and is yet to be resolved.

Surdam et al. (1972) listed the following textures as indicators of Magadi-style cherts that have likely evolved from a sodium silicate gel: 1) Preservation of the soft-sediment deformation features of the putty-like magadiite, such as enterolithic folding, lobate nodular protrusions, casts of mudcracks and trona crystals, and extrusion forms; 2) Contraction features, especially the reticulate cracks and polygonal ridges on the surface of the chert, reflecting the loss of volume in the transition (crocodile-skin chert). If the arguments of Behr and Röhricht (2000) are accepted, then only criteria 2) the characteristic shrinkage-related reticulated surface texture (crocodile-skin) should be used to interpret ancient alkaline lake cherts, along with a lack of any indications of calcium sulphate in penecontemporaneous lake sediments.

Given the type-1 hydrogeochemistry needed for highly alkaline continental brines, evaporite minerals likely to be found in association with Magadi-style cherts are the sodium carbonate salts (trona, gaylussite or pirssonite; searlesite) or their pseudomorphs; gypsum and other forms of calcium sulphate are never present in type 1 (trona -precipitating) brines (Figure 10; Hardie and Eugster, 1970). This contrast with the silica replacement mechanisms that occur when gypsum or anhydrite nodules are silicified.

Across the Phanerozoic rock record, ancient examples of crocodile-skin cherts are not common, compared to documented examples of silicified anhydrite nodules (cauliflower chert). Documented examples include: Cambrian alkaline lacustrine sediments in South Australia (White and Youngs, 1980); Eocene Green River sediments in the USA (Eugster and Surdam, 1973), the Middle Devonian in the Orcadian Basin of Scotland (Parnell, 1988) and fluviolacustrine sediments of the Permian Balzano volcanic complex in Italy (Krainer and Spötl, 1998). For all such ancient occurrences of ancient crocodile-skin chert, it should be remembered that the precipitation mechanism in its type area of Lake Magadi is still contentious. That is, Magadi-type chert, is historically interpreted as being diagenetically derived from magadiite, a hydrous sodium silicate precursor deposited from strongly alkaline lake waters. More recent work in Lake Magadi concludes that the same chert, hosted in the same High Magadi Beds is due to chemical decomposition of pyroclastic deposits by alkaline groundwater, and that chert precipitation is strongly influenced by fluctuating levels of biogenic CO2. The numerous deformation features in the High Magadi Beds in this more recent interpretation are unrelated to the mineralogical transformation of magadiite to chert (Behr, 2002).


 

Across longer time frames than that preserved in the Pleistocene sediments of Lake Magadi, the chemical proportions of various ionic components in seawater are not constant (Figure 10). The varying proportions are intimately related to the evolution of the world’s atmosphere and rates of seafloor spreading (Warren, 2016; Chapter 2). Sulphate (rather than sulphide) only became a significant component in the world’s ocean around 2Ga. Before that, the world’s oceans and its atmosphere lacked significant oxygen, and entrained much higher proportions of CO2 and methane. In the Archean, the world’s oceans were Na-Cl-Ca-HCO3 waters, not the Na-Cl-Mg-SO4 systems of today and trona, along with halite were primary precipitates in marine hypersaline settings. Under that scenario, it is likely that some marine-associated hypersaline cherts were formed via replacement of sodium silicate precursors. In Phanerozoic strata, an ability to separate cauliflower cherts (after gypsum/anhydrite nodules) from crocodile-skin cherts (associated with silicate gels in trona/natron soda lakes) is considered significant in defining marine-fed versus continental saline hydrologies. Hydrochemistry and textures associated with silicification in CaSO4-rich environments is the topic of the next blog article.

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Lake Nakuru flamingos– Life's response to feast and famine in schizohaline lacustrine hydrologies

John Warren - Monday, May 23, 2016


The Flamingo Connection

An aviator once described Lake Nakuru as “a crucible of pink and crimson fire,” with a million flamingos painting an astonishing band of colour that burst into pieces as the birds took flight (Figure 1).

Flamingo population levels in Lake Nakuru and any mass “die-offs” are popularly considered as indicators of the environmental health of Nakuru and other lakes in the African rift valley with significant flamingo populations. In 2006, more than 30,000 of the birds were found dead at Nakuru, leaving enough pink carcases to spur an international newspaper to describe the lake as a “flamingo death camp.” Two years prior, 43,800 of the birds had perished at Tanzania’s Lake Manyara, the first major die-off documented at that alkaline, soda-rich lake. Previous mass die-offs occurred at Lake Nakuru and two other Kenyan lakes in 1993, 1995, and 1997, as well as at two lakes in Tanzania in 2002. At the same time, birds were gathering in places they have never been documented before. Since 2006 there have been additional population crashes at Nakuru and Elmentia (Table 1).

      

Flamingos numbers in Lake Nakuru are perhaps one of the most visually impressive responses to episodic but very high levels of organic productivity, driven by a well-adapted species feeding in a layered saline water body subject to periodic salinity stress (Warren, 2011). The fluctuating richness of the lake’s flamingo population was dubbed the “flamingo connection” in a benchmark paper by Kirkland and Evans (1981) that considered mesohaline evaporitic carbonates as hydrocarbon source rocks.

Flamingo biology

Flamingos (Aves, Phoenicopteridae) are an ancient lineage of long-legged, microphagous, colonial wading birds. Although popularly misperceived as tropical species, flamingo distribution is more closely tied to the great deserts of the world and to hypersaline lake sites, than it is to equatorial regions (Bildstein, 1993). Flamingos are filter feeders that thrive on halotolerant cyanobacterial blooms in mesohaline shallows of saline lakes around the world. This creates the context between flamingos, mesohaline planktonic blooms, and saline lakes, well documented in Lake Nakuru by Vareschi (1982) and first noted in the geological literature in that benchmark Kirkland and Evans paper. Sedimentary textures and structures associated with flamingo lifestyles, where these birds dominate the macrofauna in some modern saline lakes, are described by Scott et al. (2009, 2012). Ancient avian counterparts can leave a characteristic set of trackways and trace fossils (including nest mounds) that can be used to refine mid-late Tertiary lacustrine depositional models (Melchor et al., 2012). Flamingo-like ancestors, which would give rise to modern ducks, even left traces in the shallow saline mudflats of Eocene soda lake sediments that define the saline portions of the trona/nahcolite-bearing sediments of the Green River formation in Utah (Figure 2).


Two species of flamingo gather in huge numbers in Lake Nakuru and a few other East African rift lakes, namely, the greater and the lesser flamingo (Phoenicoptus ruber roseus and P. minor respectively), with the lesser flamingo having characteristic and spectacular pink-red colouration in their feathers. These bright pink waders feed and breed in mesohaline rift lake waters where cyanobacterial blooms can be so dense that a Secchi disc disappears within a few centimetres of the lake’s water surface (Warren, 1986, 2011). Lake Natron (Koeppen climate Aw), where trona is the dominant evaporite, is a major breeding ground for flamingos in East Africa and is the only regular breeding site for the lesser flamingo in Africa (Simmons 1995). Lesser flamingos build nesting platforms on the trona pavement in the more central parts of the lake. These spectacular birds not only feed in saline waters, they can choose to build nest mounds on evaporite pavements!

Worldwide, only six sites are used for breeding by the lesser flamingos: Lake Natron (Tanzania), Etosha Pan (Namibia), Makgadikgadi-Pan (Botswana), Kamfers Dam (South Africa), as well as two pans in the “Little Rann of Kachchh” (India). Recent estimates of lesser flamingos at the main distribution areas are as follows: 1.5–2.5 million in eastern Africa; 390,000 in northwestern India; 55,000-65,000 in southwestern Africa; and 15,000–25,000 in western Africa. The highest population densities occur in Kenya (1.5 million) and Tanzania (600,000) (Childress et al. 2008). Lesser flamingos are well adapted to the harsh conditions associated with living and breeding in hypersaline alkaline conditions. Worldwide, they follow an itinerant lifestyle, ranging across their distribution areas in search of saline water bodies with appropriate cyanobacterial blooms. In the east African rift the flocks can travel up to 200 km a day between feeding and breeding sites, which are generally at geographically separate locations (Figure 3).


Lake Natron has the highest concentration of breeding flamingos of any lake in East Africa. Both the greater and the lesser flamingo are found there, with the lesser flamingo outnumbering the greater by a hundred to one. Lesser flamingos bred at Lake Natron in 9 out of 14 years from 1954 to 1967. But while the trona-rich nearby Lake Natron is an essential breeding site, it is not a focal feeding site for flamingos. Major feeding sites in the Africa rift valley are Lakes Nakuru and Bogoria (formerly Lake Hannington) in Kenya and entrain waters that are mesohaline with an abundance of halotolerant cyanobacteria, dominated by Arthrospira (Schagerl et al., 2015). Lake Nakuru is a mesohaline soda lake with a pH ≈ 10.5 and a typical annual salinity range of 15-45‰, nearby lake Bogoria is somewhat larger, also an alkaline soda lake, with somewhat higher salinities. The Lake Nakuru depression measures some 6.5 km by 10 km, with a water covered area of some 5-45 km,2 experiencing an annual pan evaporation rate of 1500 mm beneath a Cfb Köppen climate (Figure 3b; Vareschi,1982; Krienitz and Kotut, 2010). It contains a eutrophic bottom water mass in the lake centre, with thermally stratified water column that can be up to 4.5 metres deep. The Lake Bogoria’ water mass is up to 4 km wide, some 17 km long with thermally stratified eutrophic waters up to 10 m deep. It lies beneath a Cfb climate with a pan evaporation of 2600 mm. It is fed by a combination of rainfall (≈760mm/year) and numerous (>200) hydrothermal springs about the lake edge (Figure 3c).

Lake Bogoria hydrochemistry, it seems, is more stable compared to other endorheic lakes in Kenya, because of its greater depth (10 m), steep shores and larger water volume preventing it from drying up. In contrast, Lake Nakuru resembles more a flat pan; it is much shallower (1 m) and is more subject to changes in size related to changes in water levels, at times Lake Nakuru can dry up completely. Water depths in both lakes vary from year to year, in large part depending on the vagaries of annual runoff/ inflow. The higher level of hydrothermal inflow in Bogoria’s hydrology, along with its somewhat steeper hydrographic profile, means the lake area salinity and water depth vary less from year to year compared to Nakuru (Figure 4).


Lakes Nakuru and Bogoria, which at peak breeding times can support lesser flamingo populations in excess of 1.5 million birds, have surface areas that are less than half those of lakes Natron and Magadi, which in turn have small areas compared to most ancient lacustrine evaporites. Yet Nakuru and Bogoria are two of the most organically productive ecosystems in the world (Warren, 2011). What makes both lakes so productive are dense populations of halotolerant cyanobacteria, especially Arthrospira sp, which flourish and periodically reach peak growth in their waters yet at other times organic productivity can crash. The morphological and hydrochemical contrasts likely accounts for the more frequent Arthrospira crashes and biotal community changes in Nakuru when compared to Bogoria and other lakes in the region (Table 1).

Flamingos as filter feeders

Flamingos pass water through their bill filters in two ways (as documented by Penelope M. Jenkin in her classic article of 1957): either by swinging their heads back and forth just below the water surface, so permitting the water to flow passively through the filters of their beak, or by more efficient and more usual system of an active beak pumping. The latter is maintained by a large and powerful tongue that fills a large channel in the lower beak. As it moves rapidly back and forth, up to four times a second, it drawing water and plankton through the beak filters on the backwards pull and expelling it on the forward drive. The tongue’s surface also sports numerous denticles that scrape the collected food from the filters.


Flamingo beaks have evolved into highly efficient plankton-extraction apparati that exploit the dense cyanobacterial populations periodically found in mesohaline lakes worldwide (Figure 5; Gould, 1987; Bildstein, 1993). The beak is unlike that possessed by any other bird group on earth; the affinity is more to the baleen of whales used to filter planktonic krill from the lit upper waters of world's oceans. A flamingo beak houses a high volume water-filtering system made up of a piston-like tongue and hair-like structures called lamellae made up of rows of fringed platelets that line the inside of the mandible. In the lesser flamingo, the lamellae fibres have the appropriated spacing for capturing coiled filaments of Arthrospira. Lamellar spacings are wider in the beaks of the greater flamingo than those in the beaks of the lesser flamingo so these larger-sized birds are more generalist feeders of lake zooplankton. Thus, in any mesohaline lake where the two bird species feed and co-exist, they do not compete for the same food source. By swinging their upside-down heads from side to side just below the water surface and using the piston-like tongue to swish water through their lamella-lined beaks, flamingos can syphon the lake plankton into their gullets at phenomenal rates. Lesser flamingos can pump and filter as many as 4 beakfuls of plankton-rich water a second. This means some individuals filter upwards of 20,000 litres of water per day.

How the birds manage to cycle so much brackish to mesohaline waters, while maintaining their osmotic integrity, remains a mystery. When the rift lakes are typified by dense cyanobacterial blooms, each adult bird ingests around 72 g dry weight of Arthrospira per day (Vareschi 1978). This means the Nakuru lesser flamingo flock is able to ingest 50–94% of the daily primary production in the lake (about 60-80 tons). Rates of planktonic renewal in these rift lakes is obviously extremely high and the required rates of biomass production by Arthrospira can be spectacular. Vonshak (1997) reported doubling times of 11–20 hours of Arthrospira in a culture growing under mesohaline conditions at 35°C.


Flamingos (flamingoes) are mostly nocturnal feeders and will feed for up to 12-13 hours in a 24 hour period. The preferred planktonic food of the lesser flamingo is the cyanobacterium Arthrospira platensis (formerly known as Spirulina platensis), which for much of an average year is the widespread phytoplankton component in Lake Nakuru and Lake Bogoria waters (Figure 6). Unfortunately, dense populations of Arthrospira function best at a preferred range of temperature and salinity, meaning acme populations tend to collapse irregularly and unpredictably in both Lake Nakuru and Bogoria, leading to highly-stressed malnourished flocks of flamingos subject to mass dieback (Figure 6; Vareschi 1978; Kreinitz and Kotut, 2010).

Arthrospira has high levels of the red pigment phycoerythrin and so when ingested in large volumes it accumulates in flamingo feathers to give the birds their world famous colouration, hence the “flamingo connection.” Once digested, the carotenoid pigment dissolves in fats, which are then deposited in the growing feathers. The same effect is seen when shrimp change colour during cooking due to carotenoid alteration. The amount of pigment laid down in the feathers depends on the quantity of pigment in the flamingo’s diet. Lesser flamingos, with beak design maximised to feed on Arthrospira have a more intense pink colour in their feathers than greater flamingos. The latter species sits higher in the Lake Nakuru food chain and so gets the slight pink tinge in its feather colour, mostly second-hand from the lake zooplankton, which also feed on Arthrospira.

As well as possessing very high levels of phycoerythrin in its cytoplasm, Arthrospira is also unusual among the cyanobacteria in its unusually high protein content (some ten times that of soya). This, and the high growth rate of this species, explains why Nakuru and Bogoria acme populations can support such spectacular population levels of flamingos. A lack of cellulose in the Arthrospira cell wall means it is a source of plant protein readily absorbed by the gut, making it a potentially harvestable human food source in saline water bodies in regions of desertification. In Lake Chad, and in some saline lakes in Mexico, Arthrospira accumulates as a lake edge scum that has been harvested for millennia by the local people (including the ancient Aztecs) and used to make nutritious biscuits.

When Arthrospira stocks are low in the rift lakes, the lesser flamingo will consume benthic diatoms. However, net primary productivity of benthic diatoms in East African soda lakes is one to two orders of magnitude less than that of Arthrospira, and the carrying capacity of the habitat with diatoms is lower by the same order (Tuite 2000). This lower productivity is seen in the peak 25,000 bird population, which are diatom feeders, in Laguna de Pozuelos in Argentina (area ≈100-130 km2, more than 3 times that of Nakuru, yet the peak flamingo numbers are two order of magnitude lower). In Lake Nakuru, Arthrospira and other lake plankton are also consumed by one species of introduced tilapid fish and one species of copepod and a crustacean. Rotifers, waterboatmen, and midge larvae also flourish in the mesohaline waters of Lake Nakuru. The mouth-breeding tilapid Sarotherodon alcalicum grahami was introduced to the lake in the 1950s to control the mosquito problem and fish-feeding birds (such as pelicans) have flourished ever since (Vareschi, 1978). During times of non-optimum water conditions, when either freshening or somewhat elevated hypersalinity lessens the number of Arthrospira, the tilapids can displace the flamingos as the primary consumers of planktonic algae.


In 1972 Lake Nakuru waters held a surface biomass of 270 g/m3 and an average biomass of 194g/m3 but, as in most hypersaline ecosystems, Nakuru’s organic production rate varies drastically from year to year as water conditions fluctuate (Figure 7; Vareschi, 1978). Arthrospira was in a long-lasting bloom in 1971-1973, and accounted for 80-100% of the copious phytoplankton biomass in those years. In 1974, however, Arthrospira almost disappeared from the lake and was replaced by much smaller-diameter planktonic species, such as coccoid cyanobacteria that dealt better with elevated salinities. This transfer in primary producer make-up in the lake waters also made the lesser flamingos less efficient feeders. When the relatively large filaments of Arthrospira dominate the lake plankton, the flamingo’s beak filters between 64 and 86% of the plankton held in each mouthful of lake water (Vareschi, 1978). When the much smaller coccoids come to dominate, the filters are a much less efficient feeding mechanism. The change in plankton species was also tied to a severe reduction in algal biomass (and protein availability), which in 1974 was down to 71 g/m3 in surface waters and averaged 137 g/m3 in the total water mass. As a result, the flamingo numbers feeding in the lake declined from 1 million to several thousand, driving a significant die-back as the salinity-stressed flamingo population moved to other lakes, like Bogoria, where Arthrospira were flourishing (Vareschi, 1978).

The lower salinity limit for an Arthrospira bloom is ≈ 5‰, but it does better when salinity is more than 20‰. The species dominance of Arthrospira and its higher biomass in somewhat more saline lake waters in the African Rift lakes is clearly seen in the near unispecific year-round biomass of nearby Lake Bogoria (salinity 40-50‰). Its surface salinity is higher year round than Lake Nakuru (salinity ≈ 30‰) and the somewhat fresher waters of Lake Elmenteita (salinity ≈ 20‰; Figure 3a). Because of this, Lake Bogoria is a more reliable food source for feeding flamingos compared to either Nakuru or Elmenteita. Birds tend to migrate there to feed when conditions for Arthrospira growth are not ideal in other nearby alkaline lakes (too fresh or too saline). This was the case in 1999 when high rainfall and dilution of lake waters caused the Arthrospira levels to fall in both Nakuru and Elmenteita. It was also true in late 2012 when Nakuru water levels were at near-historic highs and the waters too fresh to support a healthy Arthrospira population.

Flamingo biomass controls

A driving mechanism for the abrupt change in biomass in Lake Nakuru in the period 1972 -1974 was not clearly defined. It was thought to be related to increased salinity and lowering of lake levels, driving the growth of coccoid species other than Arthrospira sp. that are better adapted to higher salinity, but offering less protein to the feeding birds (Figure 7a; Vareschi, 1978). There is also the simple fact that in a lake with no surface outflow, ever more saline waters cover ever-smaller areas on the lake floor. There have been times in the last 70 years when most of Lake Nakuru has dried up and pools of saline water only a few tens of centimetres deep remained. This was the case in 1962 and most recently the case in 2008 (Figure 4).

After the lake level lows of the 1960s, during the mid to late 1970s and in the 1980s the Nakuru hydrology returned to more typical inherent oscillations in water level and salinity (schizohalinity). The flamingo populations in Nakuru returned to impressive numbers but followed the vagaries of Arthrospira blooms. Since the early 1990s more reliable long-term datasets on physical lake condition and flamingo numbers have been compiled (Figures 5, 7b). In that time frame, in 1993, 1995 1998, 2008 and 2012, the flamingo populations feeding in Lake Nakuru were once again at very low levels and the remaining bird populations were stressed and subject to mass dieback.

In 1998, unlike 1974, the stress on the flamingo population was related to lake freshening and rising water levels driving the decrease in Arthrospira biomass, not increased salinity and desiccation. In the preceding bountiful years, the Arthrospira-dominated biomass had bloomed at times when salinities were favourable and died back at times of elevated salinities and lake desiccation, as in 1974. By 2000 formerly low salinities had once again increased making surface waters suitable for another widespread Arthrospira bloom and the associated return of high numbers of feeding flamingos, which continued until 2007 (Figure 7c). In 2013 there was another freshening event, with associated rising water levels and the lakes flamingo population moved to Lake Bogonia to feed.

Freshening favours a cyanobacterial assemblage dominated by picoplanktic chlorophytes (Picocyctis salinarum) and the nostocalean Anabaenopsis; the latter creates slimy masses that clog the flamingo’s feeding apparatus. The combination can drive much of the flamingo population to starvation or migration to other lakes with suitable salinities (Krienitz and Kotut, 2010). With freshening, comes also the possibility of the growth of strains that produce toxins (such as Anabaenopsis or Microcystis), possibly not in the feeding areas, which tend to remain too saline for Microcyctis, but in the spring waters where the flamingos fly in order to bathe and cleanse their feathers after a night spent feeding (Kotut and Krienitz, 2011).

It seems that breeding flamingos come to Lake Nakuru to feed in large numbers when there is water in the lake with appropriate salinity and nutrient levels to facilitate an Arthrospira bloom. In some years when heavy rains occur, lake levels rise significantly and the lake waters, although perennial, stay in the lower salinity tolerance range for Arthrospira platensis, keeping cyanobacterial numbers and protein levels at the lower end of the spectrum, as in the El Niño period between October 1997 to April 1998 and again in 2013. Once lake levels start to fall, salinities and rates of salinity change return to higher levels, then water conditions once again become appropriate for an Arthrospira bloom. But the environmental stress on the flamingos also comes with further-elevated salinities and desiccation moving lake hydrochemistry into salinities at the upper end of Arthrospira tolerance.

One of the reasons Lake Nakuru is suitable for phenomenal cyanobacterial and algal growth at times of Arthrospira bloom is the maintenance of suitable temperatures and oxygenation in the upper water mass, where the photosynthetic Arthrospira thrive. Nakuru develops a daily thermocline in the top 1.5 metres of the water column that dissipates each day in the late afternoon via wind mixing. Overturn recycles nutrients (derived from the decomposition of bird and other droppings, including those of resident hippopotami) back to the oxygenated lit surface waters to facilitate an ongoing bloom the next day (Figure 7c).

Numbers of flamingos feeding in Lake Nakuru and Lake Bogoria are used in the popular press as indicators of the environmental health of the lakes. Thousands of birds died in Lake Nakuru in 1995 and more than 30,000 birds may have died in Lake Bogoria in the first half of 1999. The most dramatic die-offs in the last two decades were at Lake Nakuru in August 2006, when some 30,000 died, and Lake Bogoria in July 2008, when 30,000 birds died. Some environmentalists have argued in the popular press that mass die-offs and their perception of lowered numbers of flamingos in Lake Nakuru and Lake Bogoria across the 1990s and 2000s were indicators of uncontrolled forest clearance, an uncontrolled increase in sewerage encouraging eutrophication, and increase in heavy metals from increasing industrial pollutants in the lake, along with general stress on the bird population from tourists and the drastic increase in local human population centred on the town of Nakuru (third largest in Kenya). Numbers of people in the town, which is the main city in the rift valley, have grown by an average of 10% every decade for the past 30 years.

But like much environmental doomsday argument, it is more based on opinionated prediction than on scientific fact. When numbers of feeding flamingos in Lake Nakuru are plotted across last few decades, it is evident that flamingo numbers oscillate widely, but it is also apparent that the peak numbers in 2000s are equivalent to the peak numbers in 1990s. A notion of longterm fall in numbers rather than wide natural fluctuations in numbers of feeding flamingos in the lake is not based on scientific reality.

Likewise, when studies were done on the cause of the mass die-off in Lake Bogoria in 1999, it was found to have a natural, not an anthropogenic cause (Krienitz at al., 2003). The flamingos had ingested the remains of toxic cyanobacteria that constitute part of the population of the microbial mats that had bloomed to form a floor to the fresh water thermal spring areas about the lake edge. There the mats are dominated by thermally tolerant species; Phormidium terebriformis, Oscillatoria willei, Arthrospira subsalsa and Synechococcus bigranulatus. The influence of cyanotoxins in the deaths of the birds is reflected in autopsies which revealed: (a), the presence of hot spring cyanobacterial cells and cell fragments (especially Oscillatoria willei), and high concentrations of the cyanobacterial hepato- and neurotoxins in flamingo stomach contents and faecal pellets; (b), observations of neurological signs of bird poisoning - birds died with classic indications of neurotoxin poisoning - the ophistotonus behaviour (neck snapped back like a snake) of the flamingos in the dying phase, and the convulsed position of the extremities and neck at the time of death. Cyanobacterial toxins in stomach contents, intestine and faecal pellets were 0.196 g g-1 fresh weight (FW) for the microcystins and 4.34 g g-1 FW for anatoxin-a. Intoxication with cyanobacterial toxins probably occurred via uptake of detached cyanobacterial cells when birds come daily to the springs to drink and wash their feathers after an overnight feeding session in the saline waters of the lake proper.

When heavy metal studies were undertaken in Nakuru lake sediments, the amount of heavy metals (Cd, Cr, Cu, Hg, Ni, Pb, Zn) were found to be in the typical range of metals in sediments in lakes worldwide. The exception is Cd, which is elevated and can perhaps be ascribed to anthropogenic activity (Svengren, 2002). All other metals are present at low levels, especially if one considers that Lake Nakuru lies within a labile catchment where the bedrock is an active volcanogenic-magmatic terrane.

Nearby Lake Magadi is also characterised by seasonal freshening, high productivity levels and bright red waters. In this case, the colour comes from haloalkaliphilic archaea, not cyanobacteria. Archaeal species belonging to the genera Natronococcus, Natronobacterium, Natrialba, Halorubrum, Natronorubrum and Natronomonas, all occur in soda brines of Lake Magadi. Lake centre brines where this biota flourishes is at trona/halite saturation with a pH up to 12. Stratified moat waters around the trona platform edge are less chemically extreme and moat bottom sediments preserve elevated levels of organics (≈6-8%). Lake Magadi also harbours a varied anaerobic bacterial community in the moat waters, including cellulolytic, proteolytic, saccharolytic, and homoacetogenic bacteria (Shiba and Horikoshi, 1988; Zhilina and Zavarzin, 1994; Zhilina et al., 1996). When the homoacetogen Natroniella acetigena was isolated from this environment, its pH growth optimum was found to be 9.8–10.0, and it continued to grow in waters with pH up to 10.7 (Zhilina et al., 1996).

 

Flamingo numbers track feast and famine

Rise and fall of lake levels, drastic changes in salinity, a periodically stressed biota, and a lack of predictability in water character are endemic to life in saline ecosystems (cycles of “feast or famine”). Natural variations in hydrochemistry control the number of feeding flamingos in Nakuru and Bogoria. In general, sufficient base-line scientific data in these schizohaline ecosystems is not yet available and so accurate determinations of the relative import of increased human activity versus natural environmental stresses on longterm bird numbers are not possible.

 

Regionally, salinities in the east African rift valley lakes range from around 30-50‰ total salts (w/v) in the more northerly lakes in the rift (Figure 3a; Bogoria, Nakuru, Elmenteita, and Sonachi) to trona and halite saturation (>200‰) in lakes to the south (Lakes Magadi and Natron). Yet across this salinity range a combination of high ambient temperature, high light intensity and a continuous resupply of CO2, makes some of these soda lakes amongst the highest in the world in terms of their seasonal planktonic biomass (Grant et al., 1999) and also places them among the world’s most productive ecosystems (Figure 8; Melack and Kilham, 1974). Organic production is periodic, and pulses of organic product periodically swamp the ability of the decomposers and so accumulate as laminites in the perennial water-covered areas of some lake centres.


Less-alkaline lakes in the rift valley are dominated by periodic blooms of cyanobacteria, while the hypersaline lakes, such as Magadi, can on occasion support blooms of cyanobacteria, archaea and alkaliphilic phototrophic bacteria (Jones et al., 1998). Halotolerant and halophilic biota living in the variably saline and layered water columns constitute small-scale “feast or famine” ecosystems, which at times of “feast” are far more productive than either tropical seagrass meadows or zones of marine upwelling (Figure 8).

The “flamingo connection” across the African Rift Lakes supports a general observation that short periods of enhanced organic productivity are followed by episodes of lessened productivity in various schizohaline saline lake and seaway waters worldwide. both past and present (Warren, 2011). It reflects the general principle that increased environmental stress favours the survival of a few well-adapted and specialised halotolerant species. This biota is well adapted to the feast and famine life-cycle that exists in most saline depressions and means their numbers are subject to wide fluctuations tied to wide fluctuations in a saline lakes hydrochemistry (schizohaline waters). This same general principle of schizohaline ecosystem adaption is clearly seen in the periodic decrease in invertebrate species (grazers and predators) numbers with increasing salinity in the carbonate lakes of the Coorong of Southern Australia, where the only metazoan to remain alive in waters with salinities more than 200‰ are the southern hemisphere brine shrimp (Paratemia zietziania). It is seen in population fluctuations of the motile alga Dunaliella sp. in the Dead Sea, and in the fluctuating purple bacterial communities of Lake Mahoney in Canada (Warren, 2011). All these examples underline a general principle of “life will expand into the available niche” a paradigm that in the cases we have discussed is driven by fluctuating salinities inherent to saline-tolerant and saline-adapted ecosystems.

References

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Childress, B., S. Nagy, and B. Hughes, 2008, International Single Species Action Plan for Conservation of the Lesser Flamingo (Phoenicopterus minor): CMS Technical Series No. 18, AEWA Technical Series No. 34, Bonn, Germany, 59 pp.

Gould, S. J., 1987, The Flamingo's Smile: Reflections in NAtural History: New York, W. W. Norton and Company.

Grant, S., W. D. Grant, B. E. Jones, C. Kato, and L. Li, 1999, Novel archaeal phylotypes from an East African alkaline saltern: Extremophiles, v. 3, p. 139-145.

Jenkin, P. M., 1957, The Filter-Feeding and Food of Flamingos (Phoenicopteri): Philosophical Transactions of the Royal Society of London. Series B, Biological Sciences, v. 240, p. 401-493.

Jones, B. E., W. D. Grant, A. W. Duckworth, and G. G. Owenson, 1998, Microbial diversity of soda lakes: Extremophiles, v. 2, p. 191-200.

Kirkland, D. W., and R. Evans, 1981, Source-rock potential of evaporitic environment: Bulletin American Association of Petroleum Geologists, v. 65, p. 181-190.

Kotut, K., and L. Krienitz, 2011, Does the potentially toxic cyanobacterium Microcystis exist in the soda lakes of East Africa?: Hydrobiologia, v. 664, p. 219-225.

Krienitz, L., A. Ballot, K. Kotut, C. Wiegand, S. Pütz, J. S. Metcalf, G. A. Codd, and S. Pflugmache, 2003, Contribution of hot spring cyanobacteria to the mysterious deaths of Lesser Flamingos at Lake Bogoria, Kenya: FEMS Microbiology Ecology, v. 43, p. 141-148.

Krienitz, L., and K. Kotut, 2010, Fluctuating algal food populations and the occurrence of Lesser Flamingos (Phoeniconaias minor) in three Kenyan rift valley lakes: Journal of Phycology, v. 46, p. 1088-1096.

Melack, J. M., and P. Kilham, 1974, Photosynthetic rates of phytoplankton in East-African alkaline saline lakes: Limnology and Oceanography, v. 19, p. 743-755.

Melchor, R. N., M. C. Cardonatto, and G. Visconti, 2012, Palaeonvironmental and palaeoecological significance of flamingo-like footprints in shallow-lacustrine rocks: An example from the Oligocene-Miocene Vinchina Formation, Argentina: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 315-316, p. 181-198.

Schagerl, M., A. Burian, M. Gruber-Dorninger, S. O. Oduor, and M. N. Kaggwa, 2015, Algal communities of Kenyan soda lakes with a special focus on Arthrospira fusiformis: Fottea, v. 15, p. 245-257.

Scott, J. J., R. W. Renaut, L. A. Buatois, and R. B. Owen, 2009, Biogenic structures in exhumed surfaces around saline lakes: An example from Lake Bogoria, Kenya Rift Valley: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 272, p. 176-198.

Scott, J. J., R. W. Renaut, and R. B. Owen, 2012, Impacts of flamingos on saline lake margin and shallow lacustrine sediments in the Kenya Rift Valley: Sedimentary Geology, v. 277-278, p. 32-51.

Shiba, H., and K. Horikoshi, 1988, Isolation and characterization of novel anaerobic, halophilic eubacteria from hypersaline environments of western America and Kenya: In: Proceedings of the FEMS symposium—The microbiology of extreme environments and its biotechnological potential, Portugal, p. 371-373.

Simmons, R. E., 1995, Population declines, viable breeding areas and management options for flamingos in southern Africa: Conservation Biology, v. 10, p. 504-514.

Svengren, H., 2002, A study of the environmental conditions in Lake Nakuru, Kenya, using isotope dating and heavy metal analysis of sediments: Masters thesis, Dept. Structural Chemistry, University of Stockholm, Swededn.

Tuite, C. H., 2000, The Distribution and Density of Lesser Flamingos in East Africa in Relation to Food Availability and Productivity: Waterbirds: The International Journal of Waterbird Biology, v. 23, Special Publication 1: Conservation Biology of Flamingos, p. 52-63.

Vareschi, E., 1978, The ecology of Lake Nakuru (Kenya). I. Abundance and feeding of the lesser flamingo: Oecologia, v. 32, p. 11-35.

Vareschi, E., 1982, The ecology of Lake Nakuru (Kenya). III. Abiotic factors and primary production: Oecologia, v. 55, p. 81-101.

Vonshak, A., 1997, Spirulina: Growth, physiology and biochemistry, in A. Vonshak, ed., Spirulina platensis (Arthrospina): Physiology, cell-biology and biotechnology: London, Taylor and Francis, p. 43-65.

Warren, J. K., 1986, Shallow water evaporitic environments and their source rock potential: Journal Sedimentary Petrology, v. 56, p. 442-454.

Warren, J. K., 2011, Evaporitic source rocks: mesohaline responses to cycles of “famine or feast” in layered brines, Doug Shearman Memorial Volume, (Wiley-Blackwell) IAS Special Publication Number 43, p. 315-392.

Zhilina, T. N., and G. A. Zavarzin, 1994, Alkaliphilic anaerobic community at pH 10: Curr. Microbiol., v. 29, p. 109-112.

Zhilina, T. N., G. A. Zavarzin, F. Rainey, V. V. Kevbrin, N. A. Kostrikina, and A. M. Lysenko, 1996, Spirochaeta alkalica sp nov, Spirochaeta africana sp nov, and Spirochaeta asiatica sp nov, alkaliphilic anaerobes from the Continental soda lakes in Central Asia and the East African Rift: International Journal of Systematic Bacteriology, v. 46, p. 305-312.

 

 

Red Sea metals: what is the role of salt in metal enrichment?

John Warren - Friday, April 29, 2016

 

Introduction

Work over the past four decades has shown many sediment-hosted stratiform copper deposits are closely allied with evaporite occurrences or indicators of former evaporites, as are some SedEx (Sedimentary Exhalative) and MVT (Mississippi Valley Type) deposits (Warren, 2016). Some ore deposits, especially those that have evolved beyond greenschist facies, can retain the actual salts responsible for the association, primarily anhydrite relics, in proximity to the ore. Such deposits include the Zambian and Redstone copper belts, Creta, Boleo, Corocoro, Dzhezkazgan, Kupferschiefer (Lubin and Mansfeld regions), Largentière and the Mt Isa copper association. All these accumulations of base metals are associated with the formation of a burial-diagenetic hypersaline redox/mixing front, where either copper or Pb-Zn sulphides tended to accumulate. Mechanisms that concentrate and precipitate base metal ores in this evaporite, typically halokinetic, milieu are the topic of upcoming blogs. Then there are deposits that are the result from hot brine fluids, tied to dissolving evaporites and igneous activity, mixing and cooling with seawater, so precipitating a variety of hydrothermal salts, sometimes in including economic levels of copper, lead and zinc (Warren, 2016)

In this article, I focus on one such hypersaline-brine deposit, the cupriferous hydrothermal laminites of the Atlantis II Deep in the Red Sea and look at the role of evaporites in the enrichment of metals in this deposit. It is a modern example of a metalliferous laminite forming in a brine lake sump on the deep seafloor where the brine lake and the stabilisation of the precipitation interface is a result of the dissolution of adjacent halokinetic salt masses. Most economic geologists classify the metalliferous Red Sea deeps as SedEx deposits, but the low levels of lead and high levels of copper, along with its stratigraphic position atop seafloor basalts, place it outside the usual Pb-Zn dominant system that typifies ancient SedEx deposits. Some economic geologists use the Red Sea deeps as analogues for volcanic massive sulphides, and some argue it even illustrates aspects of some stratiform Cu accumulations. Many such economic geology studies have the propensity to ignore the elephant in the room; that is the Red Sea deeps are the result of brine focusing by a large Tertiary-age halokinetically-plumbed seafloor brine association. This helps explain the large volume of metals compared to Cyprus-style and mid-ocean ridge volcanic massive sulphides (Warren 2016, Chapters 15 and16).

In my mind what is most important about the brine lakes on the deep seafloor of the Red Sea is the fact that they exist with such large lateral extents only because of dissolution of the hosting halokinetic slope and rise salt mass. Seismic surveys conducted in the past decade in the Red Sea show extensive salt flows (submarine salt glaciers) along the whole of the Red Sea Rift (at least from 19–23°N; Augustin et al., 2014; Feldens and Mitchell, 2015)). In places, these salt sheets flow into and completely blanket the axial region of the rift. Where not covered by namakiers, the seafloor comprises volcanic terrain characteristic of a mid-ocean spreading axis. In the salt-covered areas, evidence from bathymetry, volume-balance of the salt flows, and geophysical data all seems to support the conclusion that the sub-salt basement is mostly basaltic in nature and represents oceanic crust (Augustin et al., 2014).

 

The Rift

The Red Sea, located between Egypt and Saudi Arabia, represents a young active rift system that from north to south transitions from continental to oceanic rift (Rasul and Stewart, 2015). It is one of the youngest marine zones on Earth, propelled by an area of relatively slow seafloor spreading (≈1.6 cm/year). Together with the Gulf of Aqaba-Dead Sea transform fault, it forms the western boundary of the Arabian plate, which is moving in a north-easterly direction (Figure 1; Stern and Johnson, 2010). The plate is bounded by the Bitlis Suture and the Zagros fold belt and subduction zone to the north and north-east, and the Gulf of Aden spreading center and Owen Fracture Zone to the south and southeast. The Red Sea first formed about 25 Ma ago in response to crustal extension related to the interface movements of the African Plate, the Sinai Plate, and the Arabian Plate (Schardt, 2016). The present site of Red Sea rifting is controlled, or largely overprinting, on pre-existing structures in the crust, such as the Central African Fault Zone. In the area between 15° and 20° along the rift axis, active seafloor spreading is prominent and is characterized by the formation of oceanic crust with Mid-Ocean Ridge Basalt (MORB) composition for the last 3 Ma (Rasul and Stewart, 2015). In contrast, the northern portion of the Red Sea sits in a magmatic continental rift in which a mid-ocean ridge spreading centre is just beginning to form. That is, the split in the crust that is the Red Sea is unzipping from south to north (Figure 1).

The Salt

The rift basement is covered a thick sequence of middle Miocene evaporites that precipitated in the earlier hydrographically isolated stage of rifting (Badenian – Middle Miocene). The maximum thickness of rift-fill sediments, including halokinetic salt, is around 8,000 m in the Morgan basin in the southern Red Sea (Farhoud, 2009; Ehrhardt et al., 2005). Girdler and Southren (1987) conclude that Miocene evaporites first accumulated on Red Sea transitional crust but must have later flowed downdip to now cover parts of the axial zone (basaltic) of the Plio-Pleistocene oceanic crust. At latitudes of 20° to 23° N, transform fracture zones provide focused passage-ways for salt flow. They also enable the involvement of dissolving salt in axial hydrothermal circulation, so producing pools of dense hot brines and the topographic isolation of spreading segments into evaporite-enclosed deeps (Feldens and Mitchell, 2015). So today, flow-like features cored by Miocene evaporites are situated along the axis of the Red Sea atop younger magnetic seafloor spreading anomalies. However, not all brine seeps occur in or near the deep axis of the Red Sea on the downdip edge of flowing Miocene salt, some occur in much shallower suprasalt positions nearer the coastal margins of the Red Sea, in waters just down dip of actively-growing well-lit coral reefs (Batang et al., 2012).


Six salt flows, most showing rounded fronts in plan-view, with heights of several hundred meters and widths between 3 and 10 km, are seen in high-resolution bathymetry and DSDP core material around Thetis Deep and Atlantis II Deep, and between Atlantis II Deep and Port Sudan Deep (Figure 2; Feldens and Mitchell, 2015; Mitchell et al., 2010). Relief on the underlying volcanic basement surface likely controls the positions of individual salt flow lobes. On the flow surfaces, along-slope and downslope ridge and trough morphologies have developed parallel to the local seafloor gradient, presumably due to the extension of the hemiplegic sediment cover or strike-slip movement within the evaporites.

Some sites with irregular seafloor topography are observed close to the flow fronts, interpreted to be the result of dissolution of Miocene evaporites, which contributes to the formation of brine lakes in several of the endorheic deeps (Feldens and Mitchell, 2015). Based on the vertical relief of the flow lobes, deformation is still taking place in the upper part of the evaporite sequence. Considering the salt flow that creates the Atlantis II Deep in more detail, strain rates due to dislocation creep and pressure solution creep are estimated to be 10−14 sec-1 and 10−10 sec-1, respectively, using given assumptions of grain size and deforming layer thickness (Feldens and Mitchell, 2015). The latter strain rate is comparable to strain rates observed for onshore salt flows in Iran and signifies flow speeds of several mm/year for some offshore salt flows. Thus, salt flow movements can potentially keep up with Arabia–Nubia tectonic half-spreading rates across large parts of the Red Sea (Figure 1)


The Deeps

Beneath waters more than a kilometre deep, along the deep rift axis, there are 26 brine pools and deeps, some of which are underlain by metalliferous sediments (Figure 3; Blanc and Anschutz 1995, Blum and Puchelt, 1991). Because of varying size, age, and formation history between the various deeps, Ehrhardt and Hübscher (2015) discriminate between central and northern Red Sea deeps. The larger central Red Sea deeps are located in the axial trough and are separated by inter-trough zones. They are floored by young basaltic crust and exhibit magnetic anomalies not older than 1.7 Ma. The northern Red Sea deeps are smaller and form only isolated deeps within the axial depression. Some of them are accompanied by volcanic activity. Many of the central Red Sea deeps contain bottom-water brines and metalliferous sediments, pointing to hydrothermal circulation of seawater (Schmidt et al., 2015). The largest and most prominent deep is the Atlantis II Deep, located in the central part of the Red Sea in the vicinity of other large deeps such as the Chain Deep and Discovery Deep. Other prominent deeps are the Tethys and Nereus Deeps further north, but still in the central part of the Red Sea.


Historically, the various deeps along the Red Sea rift axis are deemed to be initial seafloor spreading cells that will accrete sometime in the future into a continuous spreading axis. Northern Red Sea deeps are isolated structures often associated with single volcanic edifices in comparison to the further-developed and larger central Red Sea deeps where small spreading ridges are locally active (Ehrhardt and Hübscher, 2015). But not all deeps are related to initial seafloor spreading cells, and there are two types of ocean deeps: (a) volcanic and tectonically impacted deeps that opened by a lateral tear of the Miocene evaporites (salt) and Plio-Quaternary overburden; (b) non-volcanic deeps built by subsidence of Plio-Quaternary sediments due to evaporite subrosion (dissolution) processes. Type b) deeps develop as evaporite collapse structures (Figure 4: Ehrhardt and Hübscher, 2015). In contrast, the type (a) volcanic deeps can be correlated with their positions in NW–SE-oriented segments of the Red Sea, which are daylighted volcanic segments. The N–S segments, between these volcanically active NW–SE segments, is called a “non-volcanic segment” as no volcanic activity is known, in agreement with the magnetic data that shows no major anomalies. Accordingly, the deeps in the "nonvolcanic segments" are evaporite collapse-related structures creating discontinuities and brine breakout zones in and atop the salt sheets without the need for a seafloor spreading cell.

Such evaporite collapse-type ocean deeps are not limited to the non-volcanic segments, as subrosion processes driven by upwells in hydrothermal circulation are possible at any part of the axial depression, especially along fault damage zones. The combined interpretation of bathymetry and seismic reflection profiles gives further insight into the nature of lateral salt gliding in the Red Sea. Salt rises are typically present where the salt flows above basement faults. The internal reflection characteristic of the salt changes laterally from reflection-free to stratified, which suggests significant salt deformation during the salt deposition. Acoustically-transparent halite accumulated locally and evolving rim synclines were filled by stratified evaporite-related facies. (Figure 5)


Both types of deeps, as defined by Ehrhardt and Hübscher (2015), are surrounded by thick halokinetic masses of Miocene salt with brine chemistry in the bottom brine layer that signposts ongoing halite subrosion and dissolution. Red Sea deeps were discovered in the 1960s at a time when lateral translation of salt (gliding and spreading) and the formation of density stratification that define deepsea hypersaline anoxic lakes (DHALS) were not known (Warren, 2016). Today, with our knowledge of seeps and hypersaline seafloor depressions in halokinetic terranes on the slope and rise in the Gulf of Mexico and accretionary ridges in the parts of the Mediterranean Sea, we now know that the brine-filled deeps on the floor of the Red Sea are just another example of DHALs. What is most interesting in the chemical make-up Red Sea DHALS are the elevated levels of iron, copper and lead that occur in some deeps, especially the deepest and one of the most hypersaline set of linked depressions known as the Atlantis II deep (Figure 6).


Brine Chemistry in Red Sea DHALS

Most Red Sea deeps contain waters with somewhat elevated salinities, compared to normal seawater. Bulk chemistry of major ions in bottom brines from the various Red Sea DHALS are covariant and are derived by dissolution of the adjacent and underlying Miocene halite (Figure 7; replotted from Schmidt et al., 2015).


Mineralization in Red Sea DHALS

Economically, the most important brine pool is the Atlantis II Deep; other smaller deeps, with variable development of metalliferous muds and brine sumps, include; Commission Plain, Hatiba, Thetis, Nereus, Vema, Gypsum, Kebrit and Shaban Deeps (Figure 3; Chapter 15, Warren 2016). Laminites of the Atlantis II Deep are highly metalliferous, while the Kebrit and Shaban deeps are of metalliferous interest in that fragments of massive sulphide from hydrothermal chimney sulphides were recovered in bottom grab samples (Blum and Puchelt, 1991). All Red Sea DHALS are located in sumps along the spreading axis, in the region of the median valley. Most of these axial troughs and deeps are also located where transverse faults, inferred from bathymetric data, seismic, or from continuation of continental fracture lines, cross the median rift valley in regions that are also characterised by halokinetic Miocene salt. Not all Red Sea deeps are DHALS and not all Red Sea DHALS overlie metalliferous laminites.

The variably metalliferous seafloor deeps or deepsea hypersaline anoxic lakes (DHALs) in the deep water axial rift of the Red Sea define the metalliferous end of a spectrum of worldwide DHALs formed in response to sub-seafloor dissolution of shallowly-buried halokinetic salt masses. What makes the Red sea deeps unique is that they can host substantial amounts of metal sulphides, and, as Pierre et al. (2010) show, a Red Sea deep without the seafloor brine lake, is not significantly mineralised.

In my opinion, it is the intersection of the DHAL setting with an active to incipient midocean ridge (ultimate metal source), and a lack of sedimentation in the DHAL, other than hydrothermal precipitates (including widespread hydrothermal anhydrite), that explains the size and extent of the Atlantis II deposit. Its salt-dissolution-related brine hydrology, with a lack of detrital input, changes the typical mid-ocean massive-sulphide ridge deposit (with volumes usually around 300,000 and up to 3 million tonnes; Hannington et al., 2011) into a more stable brine-stratified bottom hydrology, which can fix metals over longer time and stability frames, so that the known sulphide accumulation in the Atlantis II Deep today has a metal reserve that exceeds 90 million tonnes.


The Red Sea DHAL evaporite-metal-volcanic association underlines why vanished evaporites are significant in the formation of giant and supergiant base metal deposits. Most thick subsurface evaporites in any tectonically-active metalliferous basin tend to flow and ultimately dissolve. Through their ongoing flow, dissolution and alteration, chloride- and sulphate-rich evaporites can create stable brine-interface conditions suitable for metal enrichment and entrapment. This takes place in subsurface settings ranging from the burial diagenetic through to the metamorphic and into igneous realms. An overview of a selection of the large-scale ore deposits associated with hypersaline brines tied to dissolving/altered and "vanished" salt masses, plotted on a topographic and salt basin base, shows that the majority of evaporite-associated ore deposits lie outside areas occupied by actual evaporite salts (Figure 8; see Warren Chapters 15 and 16 for detail). Rather, they tend to be located at or near the edges of a salt basin or in areas where most or all of the actual salts have long gone (typically via subsurface dissolution or metamorphic transformation). This widespread metal-evaporite association, and the enhancement in deposit size it creates, is not necessarily recognised as significant by geologists not familiar with the importance of "the salt that was." So evaporites, which across the Phanerozoic constitute less than 2% of the world's sediments, are intimately tied to (Warren, 2016):

 

  • All supergiant sediment-hosted copper deposits (halokinetic brine focus)
  • More than 50% of world’s giant SedEx deposits (halokinetic brine focus)
  • More than 80% of the giant MVT deposits (sulphate-fixer & brine)
  • The world's largest Phanerozoic Ni deposit
  • Many of the larger IOCG deposits (meta-evaporite, brine and hydrothermal)
References

 

Augustin, N., C. W. Devey, F. M. van der Zwan, P. Feldens, M. Tominaga, R. A. Bantan, and T. Kwasnitschka, 2014, The rifting to spreading transition in the Red Sea: Earth and Planetary Science Letters, v. 395, p. 217-230.

Batang, Z. B., E. Papathanassiou, A. Al-Suwailem, C. Smith, M. Salomidi, G. Petihakis, N. M. Alikunhi, L. Smith, F. Mallon, T. Yapici, and N. Fayad, 2012, First discovery of a cold seep on the continental margin of the central Red Sea: Journal of Marine Systems, v. 94, p. 247-253.

Blanc, G., and P. Anschutz, 1995, New stratification in the hydrothermal brine system of the Atlantis II Deep, Red Sea: Geology, v. 23, p. 543-546.

Blum, N., and H. Puchelt, 1991, Sedimentary-hosted polymetallic massive sulphide deposits of the Kebrit and Shaban Deeps, Red Sea.: Mineralium Deposita, v. 26, p. 217-227.

Ehrhardt, A., and C. Hübscher, 2015, The Northern Red Sea in Transition from Rifting to Drifting-Lessons Learned from Ocean Deeps, in N. M. A. Rasul, and I. C. F. Stewart, eds., The Red Sea: Berlin Heidelberg, Springer p. 99-121.

Ehrhardt, A., C. Hübscher, and D. Gajewski, 2005, Conrad Deep, Northern Red Sea: Development of an early stage ocean deep within the axial depression: Tectonophysics, v. 411, p. 19-40.

Farhoud, K., 2009, Accommodation zones and tectono-stratigraphy of the Gulf of Suez, Egypt: a contribution from aeromagnetic analysis: GeoArabia, v. 14, p. 139-162.

Feldens, P., and N. C. Mitchell, 2015, Salt Flows in the Central Red Sea, in N. M. A. Rasul, and I. C. F. Stewart, eds., The Red Sea: Springer Earth System Sciences: Berlin Heidelberg, Springer p. 205-218.

Girdler, R. W., and T. C. Southren, 1987, Structure and evolution of the northern Red Sea: Nature, v. 330, p. 716-721.

Hannington, M., J. Jamieson, T. Monecke, S. Petersen, and S. Beaulieu, 2011, The abundance of seafloor massive sulfide deposits: Geology, v. 39, p. 1155-1158.

Pierret, M. C., N. Clauer, D. Bosch, and G. Blanc, 2010, Formation of Thetis Deep metal-rich sediments in the absence of brines, Red Sea: Journal of Geochemical Exploration, v. 104, p. 12-26.

Rasul, N. M. A., and I. C. F. Stewart, 2015, The Red Sea: Springer Earth System Sciences, Springer, 638 p.

Rowan, M. G., 2014, Passive-margin salt basins: hyperextension, evaporite deposition, and salt tectonics: Basin Research, v. 26, p. 154-182.

Schardt, C., 2016, Hydrothermal fluid migration and brine pool formation in the Red Sea: the Atlantis II Deep: Mineralium Deposita, v. 51, p. 89-111.

Schmidt, M., R. Al-Farawati, and R. Botz, 2015, Geochemical Classification of Brine-Filled Red Sea Deeps, in N. M. A. Rasul, and I. C. F. Stewart, eds., The Red Sea: Berlin Heidelberg, Springer-Verlag, p. 219-233.

Stern, R. J., and P. R. Johnson, 2010, Continental lithosphere of the Arabian Plate: a geologic, petrologic, and geophysical synthesis: Earth Science Reviews, v. 101, p. 29-67.

Warren, J. K., 2016, Evaporites: A compendium (ISBN 978-3-319-13511-3) Released Feb. 2016: Berlin, Springer, 1854 p.


 

 

 

 

 

 

Salt as a Fluid Seal: Article 4 of 4: When and where it leaks - Implications for waste storage

John Warren - Thursday, March 24, 2016

 

In the three preceding articles on salt leakage, we have seen that most subsurface salt in the diagenetic realm is a highly efficient seal that holds back large volumes of hydrocarbons in salt basins worldwide (Article 3). When salt does leak or transmit fluid, it does so in one of two ways: 1) by the entry of undersaturated waters (Article 1 in this series) and; 2) by temperature and pressure-induced changes in its dihedral angle, which in the diagenetic realm is often tied to the development of significant overpressure and hydrocarbon migration (Article 2). The other implication linked to the two dominant modes of salt leakage is the source of the fluid entering the leaking salt. In the first case, the fluid source is external to the salt ("outside the salt"). In the second case, it can be internal to the main salt mass ("inside the salt"). However, due to dihedral angle changes at greater depths and pressures, a significant portion of leaking fluid passing through more deeply-buried altering salt is external. By the onset of greenschist facies metamorphism, this is certainly the case (Chapter 14 in Warren, 2016) 

Diagenetic fluids driving salt leakage are external to the salt mass

Within a framework of fluids breaching a subsurface salt body, the breached salt can be a bed of varying thickness, or it can have flowed into a variety of autochthonous and allochthonous salt masses. Autochthonous salt structures are still firmly rooted in the stratigraphic level of the primary salt bed. Allochthonous salt structurally overlies parts of its (stratigraphically younger) overburden and is often no longer connected to the primary salt bed (mother-salt level).

Breaches in bedded (non-halokinetic) salt

The principal documented mechanism enabling leakage across bedded salt in the diagenetic realm is dissolution, leading to breaks or terminations in salt bed continuity. Less often, leakage across a salt unit can occur where bedded salt has responded in a brittle fashion and fractured or faulted (Davison, 2009). In hydrocarbon-producing basins with widespread evaporite seals, significant fluid leakage tends to occur near the edges of the salt bed. For example, in the Middle East, the laterally continuous Hith Anhydrite (Jurassic) acts as a regional seal to underlying Arab Cycle reservoirs and carbonate-mudstone source rock. The high efficiency of the Hith seal creates many of the regions giant and supergiant fields, including Ghawar in Saudi Arabia, which is the largest single oil-filled structure in the world. Inherent maintenance of the evaporite's seal capacity also prevents vertical migration from mature sub-Hith source rocks into potential reservoirs in the overlying Mesozoic section across much of Saudi Arabia and the western Emirates. However, toward the Hith seal edge are a number of large fields supra-Hith fields, hosted in Cretaceous carbonates, and a significant portion of the hydrocarbons are sourced in Jurassic carbonate muds that lie stratigraphically below the anhydrite level (Figure 1).

 

The modern Hith Anhydrite edge is not the depositional margin of the laterally extensive evaporite bed. Rather, it is a dissolution edge, where rising basinal brines moving up and out of the basin have thinned and altered the past continuity of this effective seal.

The process of ongoing dissolution allowing vertical leakage near the edge of a subsurface evaporite interval, typifies not just the edge of bedded salts but also the basinward edges of salt units that are also halokinetic. The dissolution edge effect of the Ara Salt and its basinward retreat over time are clearly seen along the eastern edge of the South Oman Salt Basin where the time of filling of the Permian-hosted reservoir structures youngs toward the west (Figure 2).


 

Leakages associated with the margins of discrete diapiric structures

Once formed, salt diapirs tend to focused upward escape of basinal fluid flows: as evidenced by: (1) localized development of mud mounds and chemosynthetic seeps at depopod edges above diapirs in the Gulf of Mexico (Figure 3a); (2) shallow gas anomalies clustered around and above salt diapirs in the North Sea and (Figure 3b); (3) localized salinity anomalies around salt diapirs, offshore Louisiana and with large pockmarks above diapir margins in West Africa (Cartwright et al., 2007). Likewise, in the eastern Mediterranean region, gas chimneys in the Tertiary overburden are common above regions of thinned Messinian Salt, as in the vicinity of the Latakia Ridge (Figure 4).



Leakage of sub-salt fluids associated with salt welds and halokinetic touchdowns

Whenever a salt weld or touchdown occurs, fluids can migrate vertically across the level of a now flow-thinned or no-longer-present salt level. Such touchdowns or salt welds can be in basin positions located well away from the diapir edge and are a significant feature in the formation of many larger base-metal and copper traps, as well as many depopod-hosted siliciclastic oil and gas reservoirs (Figure 5: Warren 2016).


Caprocks are leaky

Any caprock indicates leakage and fractional dissolution have occurred along the evaporite boundary (Figure 5). Passage of an undersaturated fluid at or near the edge of a salt mass creates a zone of evaporite dissolution residues, which in the case of diapiric occurrences is called usually called a “caprock,” although such diagenetic units do not only form a “cap” or top to a salt structure.

Historically, in the 1920s and 30s, shallow vuggy and fractured caprocks to salt diapirs were early onshore exploration targets about topographic highs in the Gulf of Mexico (e.g. Spindletop). Even today, the density of drilling and geological data derived from these onshore diapiric features means many models of caprock formation are mostly based on examples in Texas and Louisiana. Onshore in the Gulf of Mexico, caprocks form best in dissolution zones at the outer, upper, edges of salt structures, where active cross-flows of meteoric waters are fractionally dissolving the salt. However, rocks composed of fractional dissolution residues, with many of the same textural and mineralogical association as classic Gulf of Mexico caprocks, are now known to mantle the deep sides of subvertical-diapirs in the North Sea (e. g., lateral caprock in the Epsilon Diapir) and define the basal anhydrite (basal caprock) that defines the underbelly of the Cretaceous Maha Sarakham halite across the Khorat Plateau in NE Thailand (Figure 5; Warren, 2016).


All “caprocks” are fractionally-dissolved accumulations of diapir dissolution products and form in zones of fluid-salt interaction and leakage, wherever a salt mass is in contact with undersaturated pore fluids (Figure 6). First to dissolve is halite, leaving behind anhydrite residues, that cross-flushing pore waters can then convert to gypsum and, in the presence of sulphate-reducing bacteria, to calcite. If the diapir experiences another growth pulse the caprock can be broken and penetrated by the rising salt. This helps explain fragments of caprock caught up in shale sheaths or anomalous dark-salt zones, as exemplified by less-pure salt-edge intersection units described as dark and anomalous salt zones in the Gulf of Mexico diapirs (as documented in Article 1).

2. Fluids that are internal to the salt mass

Fluid entry in relation to changes in the dihedral angle of halite is well documented (Article 2). It was first recorded by Lewis and Holness (1996) who postulated, based on their static-salt laboratory experiments;

"In sedimentary basins with normal geothermal gradients, halite bodies at depths exceeding 3 km will contain a stable interconnected brine-filled porosity, resulting in permeabilities comparable to those of sandstones". Extrapolating from their static halite pressure experiments they inferred that halite, occurring at depths of more than ≈3 km and temperatures above 200 °C, has a uniform intrasalt pore system filled with brine, and therefore relatively high permeabilities.

In the real world of the subsurface, salt seals can hold back significant hydrocarbon columns down to depths of more than 6-7 km (see case studies in Chapter 10 in Warren, 2016 and additional documentation the SaltWork database). Based on a compilation of salt-sealed hydrocarbon reservoirs, trans-salt leakage across 75-100 metres or more of pure salt does not occur at depths less than 7-8 km, or temperatures of less than 150°C. In their work on the Haselbirge Formation in the Alps, Leitner et al. (2001) use a temperature range >100 °C and pressures >70 MPa as defining the onset of the dihedral transition.

It seems that across much of the mesogenetic realm, a flowing and compacting salt mass or bed can maintain seal integrity to much greater depths than postulated by static halite percolation experiments. In the subsurface, there may be local pressured-induced changes in the halite dihedral angle within the salt mass, as seen in the Ara Salt in Oman, but even there, there is no evidence of the total km-scale salt mass transitioning into a leaky aquifer via changes in the halite dihedral angle (Kukla et al., 2011). But certainly, as we move from the diagenetic into the metamorphic realm, even thick pure salt bodies become permeable across the whole salt mass. Deeply buried and pressured salt ultimately dissolves as it transitions into various meta-evaporite indicator minerals and zones (Chapter 14, Warren, 2016).

When increasing pressure and temperature changes the halite dihedral angle in the diagenetic realm, then supersaturated hydrocarbon-bearing brines can enter salt formations to create naturally-hydrofractured "dark-salt". As we discussed in Article 2, pressure-induced changes in dihedral angle in the Ara Salt of Oman create black salt haloes that penetrate, from the overpressured salt-encased carbonate sliver source, up to 50 or more meters into the adjacent halite (Schoenherr et al. 2007). Likewise, Kettanah, 2013 argues Argo Salt of eastern Canada also has leaked, based on the presence of petroleum-fluid inclusions (PFI) and mixed aqueous and fluid inclusions (MFI) in the recrystallised halite (Figure 7 - see also Ara “black salt” core photos in Article 2 of this series).


Both these cases of dark-salt leakage (Ara and Argo salts) occur well within the salt mass, indicating the halokinetic salt has leaked or transmitted fluids within zones well away from the salt edge. In the case of the Argo salt, the study is based on drill cuttings collected across 1500 meters of intersected salt at depths of 3-4 km. Yet, at the three km+ depths in the Argo Salt where salt contains oil and bitumen, the total salt mass still acts a seal, implying it must have regained or retained seal integrity, after it leaked. Not knowing the internal fold geometries in any deeply buried salt mass, but knowing that all flowing salt masses are internally complex (as seen in salt mines and namakiers), means we cannot assume how far the hydrocarbon inclusions have moved within the salt mass, post-leakage. Nor can we know if, or when, any salt contact occurred with a possible externally derived hydrocarbon-bearing fluid source, or whether subsequent salt flow lifted the hydrocarbon-inclusion-rich salt off the contact surface as salt flowed back into the interior of the salt mass.

Thus, with any hydrocarbon-rich occurrence in a halokinetic salt mass, we must ask the question; did the salt mass once hydrofracture (leak) in its entirety, or did the hydrocarbons enter locally and then as the salt continued to flow, that same hydrocarbon-inclusion-rich interval moved into internal drag and drape folds? In the case of the Ara Salt, the thickness of the black salt penetration away from its overpressured source is known as it is a core-based set of observations. In the Ara Salt at current depths of 3500-4000 m, the fluid migration zones extend 50 -70 meters out from the sliver source in salt masses that are hundreds of metres thick (Kukla et al., 2011; Schoenherr et al., 2007).

So how do we characterize leakage extent in a buried salt mass without core?

Dark salt, especially if it contains hydrocarbons, clearly indicates fluid entry into a salt body in the diagenetic realm. Key to considerations of hydrocarbon trapping and long-term waste storage is how pervasive is the fluid entry, where did the fluid come from, and what are the likely transmission zones in the salt body (bedded versus halokinetic)?

In an interesting recent paper documenting and discussing salt leakage, Ghanbarzadeh et al., 2015 conclude:

“The observed hydrocarbon distributions in rock salt require that percolation occurred at porosities considerably below the static threshold due to deformation-assisted percolation. Therefore, the design of nuclear waste repositories in salt should guard against deformation-driven fluid percolation. In general, static percolation thresholds may not always limit fluid flow in deforming environments.”

Their conclusions are based on lab experiments on static salt and extrapolation to a combination of mud log and wireline data collected from a number of wells that intersected salt allochthons in Louann Salt in the Gulf of Mexico. Their lab data on changing dihedral angles inducing leakage or percolation in static salt confirms the experiments of Holness and Lewis (1996 – See Article 2). But they took the implications of dihedral angle change further, using CT imagery to document creation of interconnected polyhedral porosity in static salt at higher temperatures and pressures (Figure 8). They utilise Archies Law and resistivity measures to calculate inferred porosity, although it would be interesting what values they utilise for cementation exponent (depends on pore tortuosity) Sw and saturation exponent. Assuming the standard default values of m = 2 and n =2 when applying Archies Law to back calculate porosity spreads in halite of assumed Sw are likely incorrect.  


They then relate their experimental observations to wireline measurements and infer the occurrence of interconnected pores in Gulf of Mexico salt based on this wireline data. Key to their interpretation is the deepwater well GC8 (Figure 9), where they use a combination of a resistivity, gas chromatograms, and mud log observations to infer that hydrocarbons have entered the lower one km of a 4 km thick salt section, via dihedral-induced percolation.

 

I have a problem in accepting this leap of faith from laboratory experiments on pure salt observed at the static decimeter-scale of the lab to the dynamic km-scale of wireline-inferred observations in a salt allochthon in the real world of the offshore in deepwater salt Gulf of Mexico. According to Ghanbarzadeh et al., 2015, the three-part gray background in Figure 9 corresponds to an upper no-percolation zone (dark grey), a transition zone (moderate grey) and a lower percolation zone (light grey). This they then infer to be related to changes in dihedral angle in the halite sampled in the well (right side column). Across the data columns, what the data in the GC8 well show is:  A) Gamma log; allochthon salt has somewhat higher API values at depths shallower than 5000 m; B) Resistivity log, a change in resistivity to higher values (i.e., lower conductivity) with a change in the same cross-salt depth range as seen in the gamma log, beginning around 5100 m; C) Gas (from sniffer), shows a trend of decreasing gas content from the base of salt (around 6200 m) up to a depth around 4700 m, then relatively low values to top salt, with an interval that is possibly shalier interval (perhaps a suture - see below)  that also has a somewhat higher gas content ; D) Gas chromatography, the methane (CH4) content mirrors the total gas trends, as do the other gas phases, where measured; E) Mud Log (fluorescence response), dead oil is variably present from base of salt up to 5000 m, oil staining, oil cut and fluorescence (UV) are variably present from base salt up to a depth of 4400 m.

On the basis of the presented log data, one can infer the lower kilometer of the 4 km salt section contains more methane, more liquid hydrocarbons, and more organic material/kerogen compared to the upper 3 km of salt. Thus, the lower section of the salt intersected in the GC8 well is likely to be locally rich in zones of dark or anomalous salt, compared with the overlying 3 km of salt. What is not given in figure 9 is any information on likely levels of non-organic impurities in the salt, yet this information would have been noted in the same mud log report that listed hydrocarbon levels in the well. In my opinion, there is a lack of lithological information on the Gulf of Mexico salt in the Ghanbarzadeh et al. paper, so one must ask; "does the lower kilometer of salt sampled in the GC8 well, as well as containing hydrocarbons also contain other impurities like shale, pyrite, anhydrite, etc. If so, potentially leaky intervals could be present that were emplaced by sedimentological processes unrelated to changes in the dihedral angle of the halite (see next section).


Giving information that is standard in any mud-log cuttings description (such as the amount of anhydrite, shale, etc that occur in drill chips across the salt mass), would have added a greater level of scientific validity to to Ghanbarzadeh et al.'s inference that observed changes in hydrocarbon content up section, was solely facilitated by changes in dihedral angle of halite facilitating ongoing leakage from below the base of salt and not due to the dynamic nature of salt low as the allochthon or fused allochthons formed.  Lithological information on salt purity is widespread in the Gulf of Mexico public domain data. For example, Figure 10 shows a seismic section through the Mahogany field and the intersection of the salt by the Phillips No. 1 discovery well (drilled in 1991). This interpreted section, tied to wireline and cuttings information, was first published back in 1995 and re-published in 2010. It shows intrasalt complexity, which we now know typifies many sutured salt allochthon and canopy terrains across the Gulf of Mexico salt province. Internally, Gulf of Mexico salt allochthons, like others worldwide, are not composed of pure halite, just as is the case in the onshore structures discussed in the context of dark salt zones in article 1. Likely, a similar lack of purity and significant structural and lithological variation typifies most if not all of the salt masses sampled by the Gulf of Mexico wells listed in the Ghanbarzadeh et al. paper, including the key GC8 well (Figure 9). This variation in salt purity and varying degrees of local leakage is inherent to the emplacement stage of all salt allochthons world-wide. It is set up as the salt flow (both gravity spreading and gravity gliding) occurs at, or just below the seafloor, fed by varying combinations of extrusion or thrusting, which moves salt out and over the seabed (Figure 11).

 


 

Salt, when it is flowing laterally and creating a salt allochthon, is in a period of rapid breakout (Figure 11; Hudec and Jackson, 2006, 2007; Warren 2016). This describes the situation when a rising salt sheet rolls out over its base, much in the same way a military tank moves out over its track belt. As the salt spreads, the basal and lateral salt in the expanding allochthon mass, is subject to dissolution, episodic retreat, collapse and mixing with seafloor sediment, along with the entry of compactional fluids derived from the sediments beneath. Increased impurity levels are particularly obvious in disturbed basal shear zones that transition downward into a gumbo zone (Figure 12a), but also mantle the sides of subvertical salt structures, and can evolve by further salt dissolution into lateral caprocks and shale sheaths (Figure 6).

In expanding allochthon provinces, zones of non-halite sediment typically define sutures within (autosutures; Figure 12b) or between salt canopies (allosutures; Figure 12c). These sutures are encased in halite as locally leaky, dark salt intervals, and they tend to be able to contribute greater volumes of fluid and ongoing intrasalt dissolution intensity and alteration where the suture sediment is in contact with outside-the-salt fluids. Allochthon rollout, with simultaneous diagenesis and leakage, occurs across intrasalt shear zones, or along deforming basal zones. In the basal part of an expanding allochthon sheet the combination of shearing, sealing, and periodic leakage creates what is known as “gumbo,” a term that describes a complex, variably-pressured, shale-rich transition along the basal margin of most salt allochthons in the Gulf of Mexico (Figure 12a). Away from suture zones, as more allochthon salt rolls out over the top of earlier foot-zones to the spreading salt mass, the inner parts of the expanding and spreading allochthon body tend toward greater internal salt purity (less non-salt and dissolution residue sediment, as well as less salt-entrained hydrocarbons and fluid inclusion).

At the salt's upper contact, the spreading salt mass may carry its overburden with it, or it may be bare topped (aka open-toed; Figure 11). In either case, once salt movement slows and stops, a caprock carapace starts to form that is best developed wherever the salt edge is flushed by undersaturated pore waters (Figure 6). Soon after its emplacement, the basal zone of a salt allochthon acts a focus for rising compactional fluids coming from sediments beneath. So, even as it is still spreading, the lower side of the salt sheet is subject to dissolution, and hydrocarbon entry, often with remnants of the same hydrocarbon-entraining brines leaking to seafloor about the salt sheet edge. As the laterally-focused subsalt brines escape to the seafloor across zones of thinned and leaky salt or at the allochthon edge, they can pond to form chemosynthetic DHAL (Deepsea Hypersaline Anoxic Lake) brine pools (Figure 3a). Such seep-fed brine lakes typify the deep sea floor in the salt allochthon region of continental slope and rise in the Gulf of Mexico and the compressional salt ridge terrain in the central and eastern Mediterranean. If an allochthon sheet continues to expand, organic-rich DHAL sediments and fluids become part of the basal shear to the salt sheet (Figure 12a).


Unfortunately, Ghanbarzadeh et al., 2015 did not consider the likely geological implications of salt allochthon emplacement mechanisms and how this likely explains much of the geological character seen in wireline signatures across wells intersecting salt in the Gulf of Mexico. Rather, they assume the salt system and the geological character they infer as existing in the lower portions of Gulf of Mexico salt masses, are tied to post-emplacement changes in salt's dihedral angle in what they consider as relatively homogenous and pure salt masses. They modeled the various salt masses in the Gulf of Mexico as static, with upward changes in the salt purity indicative of concurrent hydrocarbon leakage into salt and facilitated by altered dihedral angles in the halite. A basic tenet of science is "similarity does not mean equivalence." Without a core from this zone, one cannot assume hydrocarbon occurrence in the lower portions of Gulf of Mexico salt sheets is due to changes in dihedral angle. Equally, if not more likely, is that the wireline signatures they present in their paper indicate the manner in which the lower part of a salt allochthon has spread. To me, it seems that the Ghanbarzadeh et al. paper argues for caution in the use of salt cavities for nuclear waste storage for the wrong reasons.

Is nuclear waste storage in salt a safe, viable long-term option?

Worldwide, subsurface salt is an excellent seal, but we also know that salt does fail, that salt does leak, and that salt does dissolve, especially in intrasalt zones in contact with "outside" fluids. Within the zone of anthropogenic access for salt-encased waste storage (depths of 1-2km subsurface) the weakest points for potential leakage in a salt mass, both natural and anthropogenic, are related to intersection with, or unplanned creation of, unexpected fluid transmission zones and associated entry of undersaturated fluids that are sourced outside the salt (see case histories in Chapter 7 and 13 in Warren, 2016). This intersection with zones of undersaturated fluid creates zones of weakened seal capacity and increases the possibility of exchange and mixing of fluids derived both within and outside the salt mass. In the 1-2 km depth range, the key factor to be discussed in relation to dihedral angle change inducing percolation in the salt, will only be expressed as local heating and fluid haloes in the salt about the storage cavity. Such angle changes are tied to a thermal regime induced by long-term storage of medium to high-level radioactive waste.  

I use an ideal depth range of 1-2 km for storage cavities in salt as cavities located much deeper than 2 km are subject to compressional closure or salt creep during the active life of the cavity (active = time of waste emplacement into the cavity). Cavities shallower than 1 km are subject to the effects of deep phreatic circulation. Salt-creep-induced partial cavity closure, in a salt diapir host, plagued the initial stages of use of the purpose-built gas storage cavity known as Eminence in Mississippi. In the early 1970s, this cavity was subject to a creep-induced reduction in cavity volume until gas storage pressures were increased and the cavern shape re-stabilised. Cavities in salt shallower than 1 km are likely to be located in salt intervals that at times have been altered by cross flows of deeply-circulating meteoric or marine-derived phreatic waters. Problematic percolation or leakage zones (aka anomalous salt zones), which can occur in some places in salt masses in the 1-2 km depth range, are usually tied to varying combinations of salt thinning, salt dissolution or intersection with unexpected regions of impure salt (relative aquifers). In addition to such natural process sets, cross-salt leakage can be related to local zones of mechanical damage, tied to processes involved in excavating a mine shaft, or in the drilling and casing of wells used to create a purpose-built salt-solution cavity. Many potential areas of leakage in existing mines or brine wells are the result of poorly completed or maintained access wells, or intersections with zones of “dark salt,” or with proximity to a thinned salt cavity wall in a diapir, as documented in articles 1 and 2 (and detailed in various case studies in Chapters 7 and 13 in Warren 2016).

In my opinion, the history of extraction, and intersections with leakage zones, during the life of most of the world’s existing salt mines means conventional mines in salt are probably not appropriate sites for long-term radioactive waste storage. Existing salt mines were not designed for waste storage, but to extract salt or potash with mining operations often continuing in a particular direction along an ore seam until the edge of the salt was approached or even intersected. When high fluid transmission zones are unexpectedly intersected during the lifetime of a salt mine, two things happen; 1) the mine floods and operations cease, or the flooded mine is converted to a brine extraction facility (Patience Lake) or, 2) the zone of leakage is successfully grouted and in the short term (tens of years) mining continues (Warren, 2016).

For example, in the period 1906 to 1988, when Asse II was an operational salt mine, there were 29 documented water breaches that were grouted or retreated from. Over the long term, these same water-entry driven dissolution zones indicate a set of natural seep processes that continued behind the grout job. This is true in any salt mine that has come “out of the salt” and outside fluid has leaked into the mine. “Out-of-salt” intersections are typically related to fluids entering the salt mass via dark-salt or brecciated zones or shale sheath intersections (these all forms of anomalous salt discussed in article 1 and documented in the case studies discussed in Chapter 13 in Warren 2016).

I distinguish such “out-of-salt” fluid intersections from “in-salt” fluid-filled cavities. When the latter is cut, entrained fluids drain into the mine and then flow stops. Such intersections can be dangerous during the operation of a mine as there is often nitrogen, methane or CO2 in an "in-the-salt” cavity, so there is potential for explosion and fatalities. But, in terms of long-term and ongoing fluid leakage “in-salt” cavities are not a problem.

Ultimately, because “out-of-salt” fluid intersections are part of the working life of any salt mine, seal integrity in any mine converted to a storage facility will fail. Such failures are evidenced by current water entry problems in Asse II Mine, Germany (low-medium level radioactive waste storage) and the removal of the oil formerly stored in the Weeks Island strategic hydrocarbon facility, Texas. Weeks Island was a salt mine converted to oil storage. After the mine was filled with oil, expanding karst cavities were noticed forming at the surface above the storage area. Recovery required a very expensive renovation program that ultimately removed more than 95% of the stored hydrocarbons. And yet, during the active life of the Weeks Island Salt Mine, the mine geologists had mapped “black salt” occurrences and tied them to unwanted fluid entries that were then grouted. Operations to block or control the entry of fluids were successful, and salt extraction continued apace.  This information on fluid entry was available well before the salt mine was purchased and converted to a federal oil storage facility. However, in the 1970s when the mine was converted, our knowledge of salt properties and salt's stability over the longer term was less refined than today.

Worldwide, the biggest problem with converting existing salt mines to low to medium level nuclear waste storage facilities is that all salt mines are relatively shallow, with operating mine depth controlled by temperatures where humans can work (typically 300-700 m and always less than 1.1 km). This relatively shallow depth range, especially at depths above 500 m, is also where slowly-circulating subsurface or phreatic waters are dissolving halite to varying degrees, This is where fluids can enter the salt from outside and so create problematic dark-salt and collapse breccia zones within the salt. In the long-term (hundreds to thousands of years) these same fluid access regions have the potential to allow stored waste fluids to escape the salt mass,

Another potential problem with long-term waste storage in many salt mines, and in some salt cavity hydrocarbon storage facilities excavated in bedded (non-diapiric) salt, is the limited thickness of a halite beds across the depth range of such conventional salt mines and storage facilities. Worldwide, bedded ancient salt tends to be either lacustrine or intracratonic, and individual halite units are no more than 10-50 m thick in stacks of various saline lithologies. That is, intracratonic halite is usually interlayered with laterally extensive carbonate, anhydrite or shale beds, that together pile into bedded saline successions up to a few hundred metres thick (Warren 2010). The non-halite interlayers may act as potential long-term intrasalt aquifers, especially if connected to non-salt sediments outside the halite (Figure 13). This is particularly true if the non-salt beds remain intact and hydraulically connected to up-dip or down-dip zones where the encasing halite is dissolutionally thinned or lost. Connection to such a dolomite bed above the main salt bed, in combination with damaged casing in an access well, explains the Hutchison gas explosion (Warren, 2016). Also, if there is significant local heating associated with longer term nuclear waste storage in such relatively thin (<10-50 m) salt beds, then percolation, related to heat-induced dihedral angle changes, may also become relevant over the long-term (tens of thousands of years), even in bedded storage facilities in 1-2 km depth range.


Now what?

Creating a purpose-built mine for the storage of low-level waste in a salt diapir within the appropriate depth range of 1-2 km is the preferred approach and a much safer option, compared to the conversion of existing mines in diapiric salt, but is likely to be prohibitively expensive. To minimise the potential of unwanted fluid ingress, the entry shaft should be vertical, not inclined. The freeze-stabilised “best practice” vertical shaft currently being constructed by BHP in Canada for its new Jansen potash mine (bedded salt) is expected to cost more than $1.3 billion. If a purpose-built mine storage facility were to be constructed for low to medium level waste storage in a salt diapir, then the facility should operate at a depth of 800-1000m. Ideally, such a purpose-built mine should also be located hundreds of metres away from the edges of salt mass in a region that is not part of an area of older historical salt extraction operations. At current costings, such a conventionally-mined purpose-built storage facility for low to medium level radioactive waste is not economically feasible.

This leaves purpose-built salt-solution cavities excavated within thick salt domes at depths of 1-2 km; such purpose-built cavities should be located well away from the salt edge and in zones with no nearby pre-existing brine-extraction cavities or oil-field exploration wells. This precludes most of the onshore salt diapir provinces of Europe and North America as repositories for high-level nuclear waste, all possible sites are located in high population areas and can have century-long histories of poorly documented salt and brine extraction and petroleum wells. Staying "in-the-salt" over the long-term would an ongoing problem in these regions (see case histories in chapters 7 and 13 in Warren, 2016 for a summary of some problems areas).

References

Alsharhan, A. S., and C. G. S. C. Kendall, 1994, Depositional setting of the Upper Jurassic Hith Anhydrite of the Arabian Gulf; an analog to Holocene evaporites of the United Arab Emirates and Lake MacLeod of Western Australia: Bulletin-American Association of Petroleum Geologists, v. 78, p. 1075-1096.

Andresen, K. J., M. Huuse, N. H. Schodt, L. F. Clausen, and L. Seidler, 2011, Hydrocarbon plumbing systems of salt minibasins offshore Angola revealed by three-dimensional seismic analysis: AAPG Bulletin, v. 95, p. 1039-1065.

Bowman, S. A., 2011, Regional seismic interpretation of the hydrocarbon prospectivity of offshore Syria: GeoArabia, v. 16, p. 95-124.

Cartwright, J., M. Huuse, and A. Aplin, 2007, Seal bypass systems: American Association Petroleum Geologists - Bulletin, v. 91, p. 1141-1166.

Davison, I., 2009, Faulting and fluid flow through salt: Journal of the Geological Society, v. 166, p. 205-216.

Dooley, T. P., M. R. Hudec, and M. P. A. Jackson, 2012, The structure and evolution of sutures in allochthonous salt: Bulletin American Association Petroleum Geologists, v. 96, p. 1045-1070.

Ghanbarzadeh, S., M. A. Hesse, M. Prodanović, and J. E. Gardner, 2015, Deformation-assisted fluid percolation in rock salt: Science, v. 350, p. 1069-1072.

Gillhaus, A., 2010, Natural gas storage in salt caverns - Summary of worldwide projects and consequences of varying storage objectives and salt formations, in Z. H. Zou, H. Xie, and E. Yoon, eds., Underground Storage of CO2 and Energy, CRC Press, Boca Raton, Fl., p. 191-198.

Harrison, H., and B. Patton, 1995, Translation of salt sheets by basal shear: Proceedings of GCS-SEPM Foundation 16th Annual Research Conference, Salt Sediment and Hydrocarbons, Dec 3-6, 1995, p. 99-107.

Holly Harrison, Dwight ‘Clint’ Moore, and P. Hodgkins, 2010, A Geologic Review of the Mahogany Subsalt Discovery: A Well That Proved a Play (The Mahogany Subsalt Discovery: A Unique Hydrocarbon Play, Offshore Louisiana): Search and Discovery Article #60049 Posted April 28, 2010, Adapted from presentation at AAPG Annual Convention, 1995, and from an extended abstract prepared for presentation at GCS-SEPM Foundation 16th Annual Research Conference, “Salt, Sediment and Hydrocarbons,” December 3-6, 1995.

Hudec, M. R., and M. P. A. Jackson, 2006, Advance of allochthonous salt sheets in passive margins and orogens: American Association Petroleum Geologists - Bulletin, v. 90, p. 1535-1564.

Hudec, M. R., and M. P. A. Jackson, 2007, Terra infirma: Understanding salt tectonics: Earth-Science Reviews, v. 82, p. 1-28.

Jackson, C. A. L., and M. M. Lewis, 2012, Origin of an anhydrite sheath encircling a salt diapir and implications for the seismic imaging of steep-sided salt structures, Egersund Basin, Northern North Sea: Journal of the Geological Society, v. 169, p. 593-599.

Kettanah, Y. A., 2013, Hydrocarbon fluid inclusions in the Argo salt, offshore Canadian Atlantic margin: Canadian Journal of Earth Sciences, v. 50, p. 607-635.

Kukla, P., J. Urai, J. K. Warren, L. Reuning, S. Becker, J. Schoenherr, M. Mohr, H. van Gent, S. Abe, S. Li, Desbois, G. Zsolt Schléder, and M. de Keijzer, 2011, An Integrated, Multi-scale Approach to Salt Dynamics and Internal Dynamics of Salt Structures: AAPG Search and Discovery Article #40703.

Leitner, C., F. Neubauer, J. L. Urai, and J. Schoenherr, 2011, Structure and evolution of a rocksalt-mudrock-tectonite: The haselgebirge in the Northern Calcareous Alps: Journal of Structural Geology, v. 33, p. 970-984.

Lewis, S., and M. Holness, 1996, Equilibrium halite-H2O dihedral angles: High rock salt permeability in the shallow crust: Geology, v. 24, p. 431-434.

Schoenherr, J., J. L. Urai, P. A. Kukla, R. Littke, Z. Schleder, J.-M. Larroque, M. J. Newall, N. Al-Abry, H. A. Al-Siyabi, and Z. Rawahi, 2007, Limits to the sealing capacity of rock salt: A case study of the infra-Cambrian Ara Salt from the South Oman salt basin: Bulletin American Association Petroleum Geologists, v. 91, p. 1541-1557.

Terken, J. M. J., N. L. Frewin, and S. L. Indrelid, 2001, Petroleum systems of Oman: Charge timing and risks: Bulletin-American Association of Petroleum Geologists, v. 85, p. 1817-1845.

Thrasher, J., A. J. Fleet, S. J. Hay, M. Hovland, and S. Düppenbecker, 1996, Understanding geology as the key to using seepage in exploration: the spectrum of seepage styles, in S. D., and M. A. Abrams, eds., Hydrocarbon migration and its near-surface expression, AAPG Memoir 66, p. 223-241.

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., 2016, Evaporites: A Compendium (ISBN 978-3-319-13511-3) Released Feb. 22 2016: Berlin, Springer, 1854 p.


 

Salt as a Fluid Seal: Article 3 of 4: When it doesn't leak - Seals to hydrocarbons

John Warren - Saturday, March 12, 2016

This, the third article in the series of four on salt leakage, discusses how and when salt acts as a seal. The fourth article will place this discussion in real world situations where the various salt units (especially “dark” salt) have lost long-term seal integrity and what this means in terms of CO2 geosequestration and waste storage.

Evaporite seal character

Unlike thick shales, subsurface evaporites in the diagenetic realm better fit Hunt’s (1990) definition of an actual pressure seal, which he defined as an impermeable rock with zero transmissivity maintained over long periods of geologic time. Very little subsalt fluid can escape through a salt mass that, until breached, tends to hold back all the compactional and thermohaline waters, gases or liquid hydrocarbons rising from below. In contrast, shale-seals consistently leak all these fluids to varying degrees.


In the realm of subsurface hydrocarbon exploration and development, salts (especially halite) are second only to clathrates in ability to form an effective seal to circulating pore waters and hydrocarbons, including methane. (Figure 1; Warren, 2016). Natural methane clathrates (methane-ice mixtures) are more efficient seals, but in the diagenetic realm, clathrate occurrence is limited by the inherent low-temperature stability requirement. This means clathrates act as hydrocarbon seals onshore in permafrost regions, such as some Siberian gas fields, or below clathrate layers down to depths of a hundred meters or so beneath the modern deep-sea floor, as occurs below the cold waters of the slope and rise across the halokinetic Gulf of Mexico or the non-halokinetic sediment of offshore Brunei (Warren, 2016; Warren et al., 2011a).

Like clathrates, evaporite layers can generate overpressures at very shallow burial depths, unlike clathrates they do not dissolve and dissipate in response to rising temperatures of the diagenetic burial realm. Evaporites create the highest and sharpest depth-related pressure differentials known in sedimentary basins in both overpressured and underpressured settings (Fertl 1976). Salt-sealed overpressured intervals can be as shallow as a few hundred meters below the surface or deeper than 6,000 m.


Unlike the low temperatures requirements for a clathrate seal, evaporite seals, with their extremely high entry pressures, ductility, very low permeabilities and large lateral extents, can maintain seal integrity over wide areas, even when exposed to a wide range of subsurface temperature and pressure conditions. A typical shale seal has permeability ≈ 10-1 to 10-5 md, with extreme values as low as 10-8 md (Figure 2). Quantitative measurement of evaporite permeability is beyond the capacity of standard instruments used in the oil industry and is mostly a topic of study for engineers working with waste-storage caverns. Their work shows the permeability of undisturbed halite is a nanodarcy or less, that is, undamaged subsurface salt has measured permeabilities that are less than 10-21 m2 (10-6md) with some of the tighter halite permeabilities ≈10-7 to 10-9 md. In contrast, typical massive anhydrites ≈10-5 md (Beauheim and Roberts, 2002). This explains a general “rule of thumb” used in the oil industry that a halite bed should be at least 2 m thick to be considered a possible seal, while and anhydrite bed should be at least 10 m thick. Equally important is the reliability of the geological model of the evaporite that is used to extrapolate lateral continuity in the seal (Warren, 2016). Pore pressures in thick sealing halites can approach lithostatic (Ehgartner et al., 1998) and when exceeded salt can locally fracture and leak (as discussed in the previous article).

Massive thick bedded pure halite units in the diagenetic realm usually contain few, if any, interconnected pore throats. The distance between NaCl lattice units is 2.8 x 10-10 m, while the smallest molecular diameter of a hydrocarbon molecule (methane) is 3.8 x 10-10 m. The most frequent way that hydrocarbons migrate through an unfractured undissolved halite bed is if the halite contains impurities that render it locally porous and make it brittle during deformation.

 


 

Seal capacity in flowing pure salt

Halite’s very high ductility and its ability to flow, re-anneal and re-establish lattice bonding by solution creep when subject to stress give it a low susceptibility to fracturing even when it is deforming (Figure 3). This is why cross-salt fault and fracture patterns, as seen in most salt-entraining basins, make the industry considers salt a “crack-stopper” (Figure 4). Worldwide, seismic imaging of halokinetic realms shows that the salt has flowed, while adjacent carbonate and siliciclastics sequences fractured. Halite’s ability to maintain seal integrity under stress, and so prevent the escape of hydrocarbons, reflects a combination of an ability to flow and re-anneal, and the small size of molecular interspaces in its ionically-bonded NaCl crystal lattice (Figure 5; see detailed discussion see Warren, 2016, chapters 6 and 10).


 

This propensity to flow under stress (tendency toward Newtonian flow) is why many laboratory tests and measurements consistently under-represent salt’s flow and subsurface seal integrity responses. Inherently any lab experiment is tied to short time frames of up to weeks or a few years. However, such laboratory tests are likely more relevant to real-world subsurface situations where salt in the vicinity of any wellbore is damaged by the nearby passage of the drill bit and its associated fluids. The applicability of laboratory measurements to real-world subsurface situations is a philosophical quandary inherent to many natural science experiments with a time-related possible-error component. By putting equipment into a natural subsurface salt region, or by removing salt samples from their natural deep subsurface environment to take measures in the lab, or by growing salt crystals in the lab to work on, we always alter things and so get outcomes that can never be 100% accurate with respect to the original unaltered subsurface salt setting. That is, within observational errors, how do we quantify random versus systematic errors when we are always altering the samples and the surrounds via the process of gaining access?

Whether, during catagenesis, buried halite beds that enclose organic intrabeds can release volatiles to sediments outside the salt mass is still a matter of some discussion among organic geochemists. The long-term lack of fracture or pore throats in buried salt beds is why organic-rich intrasalt carbonate or shale laminites tend to be inefficient source rocks in style 1a source rocks (Warren et al., 2011a). Likewise, possible flushing and maturation effects are poorly understood in subsurface situations where encased organic-rich beds are in contact with hydrated salts converting to their anhydrous equivalents (such as gypsum to anhydrite or carnallite to sylvite, mirabilite to thenardite). Loss of water of crystallisation in shallow burial (<0.5km) has the potential to allow organic-rich fluids to escape early as the hydrated salts transform to their anhydrous forms. Usually, such burial transformations are near complete in the first kilometre of burial and so may only allow immature hydrocarbons to escape into adjacent more porous sediments (Hite and Anders, 1991). There they must be stored, mature and remigrate during later burial if they are to act as hydrocarbon source rocks (Warren, 1986). Many intrasalt organic-rich beds survive well into the metamorphic realm and evolve into graphitic quartzites and marbles encased in meta-evaporitic albitites and scapolites.

As a general rule, even as a halite bed fractures, its inherent lack of strength and the consequent ability to flow means any microscale intercrystalline fractures quickly re-anneal by a combination of flow and pressure-solution induced recrystallisation. (Figure 5). Current consensus in the oil and gas industry is that some thin impurity-rich salt beds, interlayered with carrier beds, do leak small amounts of volatiles, but much less efficiently than thicker organic-rich mudstones and shales; whereas organics encased in thicker salt beds probably cannot leak from the unit until the enclosing salt dissolves or natural hydrofracturing occurs (as in the Ara Salt of Oman). Evaporite beds and salt allochthons constitute some of the strongest long-term subsurface barriers to the vertical migration of hydrocarbons in a sedimentary basin both as a seal to hydrocarbons and in COsequestration.

The next and final article in this series on salt leakage will consider; how and where does a salt seal leak in the real world of the subsurface?

References

Beauheim, R. L., and R. M. Roberts, 2002, Hydrology and hydraulic properties of a bedded evaporite formation: Journal of Hydrology, v. 259, p. 66-88.

Downey, M. W., 1984, Evaluating seals for hydrocarbon accumulations: Bulletin American Association of Petroleum Geologists, v. 68, p. 1752-1763.

Ehgartner, B. L., J. T. Neal, and T. E. Hinkebein, 1998, Gas Releases from Salt: SAND98-1354, Sandia National Laboratories, Albuquerque, NM, June 1998.

Fertl, W. H., 1976, Abnormal Formation Pressures: Amsterdam, Elsevier Scientific, 382 p.

Hite, R. J., and D. E. Anders, 1991, Petroleum and evaporites, in J. L. Melvin, ed., Evaporites, petroleum and mineral resources, v. 50: Amsterdam, Elsevier Developments in Sedimentology, p. 477-533.

Hunt, J. M., 1990, Generation and migration of petroleum from abnormally pressured fluid compartments: Bulletin American Association of Petroleum Geologists, v. 74, p. 1-12.

Ter Heege, J. H., J. H. P. De Bresser, and C. J. Spiers, 2005, Dynamic recrystallisation of wet synthetic polycrystalline halite: dependence of grain size distribution on flow stress, temperature and strain: Tectonophysics, v. 396, p. 35-57.

Urai, J. L., Z. Schléder, C. J. Spiers, and P. A. Kukla, 2008, Flow and Transport Properties of Salt Rocks, in R. Littke, ed., Dynamics of complex intracontinental basins: The Central European Basin System, Elsevier, p. 277-290.

Warren, J. K., 1986, Shallow water evaporitic environments and their source rock potential: Journal Sedimentary Petrology, v. 56, p. 442-454.

Warren, J. K., 2011b, Evaporitic source rocks: mesohaline responses to cycles of “famine or feast” in layered brines, Doug Shearman Memorial Volume, (Wiley-Blackwell) IAS Special Publication Number 43, p. 315-392.

Warren, J. K., 2016, Evaporites: A compendium (ISBN 978-3-319-13511-3) Released Feb. 2016: Berlin, Springer, 1854 p.

Warren, J. K., A. Cheung, and I. Cartwright, 2011a, Organic Geochemical, Isotopic and Seismic Indicators of Fluid Flow in Pressurized Growth Anticlines and Mud Volcanoes in Modern Deepwater Slope and Rise Sediments of Offshore Brunei Darussalam; Implications for hydrocarbon exploration in other mud and salt diapir provinces (Chapter 10), in L. J. Wood, ed., Shale Tectonics, v. 93: Tulsa OK, AAPG Memoir 93 (Proceedings of Hedberg Conference), p. 163-196.

 

 

 

 

 

 

Salt as a Fluid Seal: Article 2 of 4: Internal fluid source

John Warren - Wednesday, January 20, 2016

 

Black Salt: as an indicator of overpressure

The previous article in this series on salt leakage focused on black and dark salt created by ingress or interaction of undersaturated waters with relatively shallow halokinetic salt masses, with entry zones often tied to intervals of salt shear. The resulting black or dark salt textures are one style of “anomalous” salt. This article looks at fluid entry into salt in subsurface intervals of high pore pressure, exemplified by the “black salt” in the Ara salt seals of Oman. Such intervals are often tied to burial-pressure and temperature-related changes to the dihedral angle of salt (halite).


Dihedral angle changes and the permeability of salt

Permeability in intercrystalline pore networks in re-equilibrating and crystallising subsurface salt is tied to the dihedral angle  at solid-solid-liquid triple junctions (Figure 1; Lewis and Holness, 1996, Holness and Lewis, 1997). When the halite dihedral angle is higher than 60° under static laboratory conditions, this contact angle equates to the maintenance of closure of polyhedral grain boundaries by halite precipitation, and so at these lower temperatures both bedded and halokinetic recrystallized salt is impermeable (Schenk and Urai, 2004; Holness and Lewis, 1996). In this temperature range, the small amount of brine present in the salt is distributed in micrometer-sized isolated fluid inclusions at termini of salt crystal polygon apices. In contrast, when the solid-solid-liquid interfaces of increasingly heated and pressurised polyhedral halite attain dihedral angles that are less than 60° then the fluid-inclusion filled intercrystal cavities link up and the salt mass becomes permeable.


At burial temperatures >100°-150°C and pressures of 70 MPa or more, the dihedral angle has decreased to values <60°, driving a redistribution of the fluid into a thermodynamically stable network of connected, fluid-filled channels or fused fluid strings at grain boundary triple junctions. This transition may be related to the observation by Peach and Spiers (1996) that, during natural deformation of rocksalt at great depths, salt undergoes natural hydraulic fracturing or dilatancy. The dihedral angle is, therefore, a thermodynamic property that changes with pressure P and temperature T. Holness and Lewis’s experiments demonstrated that buried salt masses, subject to high pressures and elevated temperatures, can acquire intercrystalline or polyhedral permeability comparable to associated with intergranular permeability in sand.

This typically occurs at higher temperatures and pressures where intercrystal water positions link within flowing or static, but texturally re-equilibrated, salt and so creates continuous fluid strings along evolving intercrystalline junctions in the burial-recrystallised salt. The newly attained intercrystal configuration allows penetration and throughflow of hot, dense brines or hydrocarbons into and through the altered mass of salt polyhedrons. In Oman has created characteristic haloes of black salt about pressurised salt-encased carbonate slivers (next section).

At the same time as a recrytallising salt mass passes into the earlier stages of the greenschist facies, the salt is dissolving and altering to sodic scapolite (Warren, 2016; Chapter 13). Thus, through the later stages of diagenesis and into early to medium grades of metamorphism, the salt and its daughter products may act as sources and conduits for flow of chloride-rich metalliferous brines and salt slurries. This occurs as bedded and halokinetic salt evolves from a dense impermeable salt mass into permeable salt with higher dihedral angles and so explains salt’s significant role in the creation of many of massive base metal and IOCG deposits (Warren, 2016; Chapters 15, 16).


Black salt and overpressure in Oman

The transition in dihedral angle with increasing pressure and temperature explains the occurrence of black (bitumen-charged) haloes in salt encasing some carbonate-sliver reservoirs in the Ara Salt of Oman (Figure 2; Kukla et al., 2011a, b). Once this recrystallization occurs, the previous lower P&T mosaic halite loses its ability to act as an aquitard or aquiclude (seal) and can instead serve as a permeable conduit for escaping highly-pressurised and hydrocarbon-rich formation waters. According to Lewis and Holness, the depth at which the recrystallization occurs may begin as shallow as a few kilometres (Figure 1). But, their pressure bomb laboratory-based static-salt experiments did not completely encompass the ability of natural salt to pressure creep and self-seal by longer-term diffusion-controlled pressure solution (Warren 2016, Chapter 6). Even if the changing dihedral angles alter and open up permeability at such shallow depths, there is no guarantee that subsequent flowage associated with pressure solution will not re-anneal these new pores. The ability of salt to continue to act as a highly efficient hydrocarbon seal to depths of 6-10 km means, in my opinion, that bedded salt does may become a relative aquifer until attaining depths of 6-10 km or more. This occurs certainly at temperatures and pressures where the sequence is entering the greenschist realm. In extremely overpressured situations the transition of dihedral angles is much shallower, as in the 40-50m thick black salt rims that typify the salt-encased hydrocarbon-charged carbonate stringers in the Ara Salt of Oman (Kukla et al., 2011b). Once it does transform into polyhedral halite, a former aquiclude becomes an aquifer flushed by chloride-rich brines, likely carrying other volatiles.

A release of entrained inclusion (±intercrystalline) water at temperatures > 300-400°C (early greenschist) influences the textures of deeply buried halite. Most of the inclusions in chevron halite and other inclusion-rich cloudy primary salts are due to entrained brine inclusions and not mineral matter. Figure 3 plots the weight loss of various types of halite during heating. It clearly shows cloudy (inclusion-rich) halite releases up to 5 times more brine (0.2-0.5 wt%) than clear coarsely crystalline halite. An analysis of all fluids released during heating shows carbon dioxide and hydrogen contents are much lower than the water volumes: CO2/H2O < 0.01 and H2/H2O < 0.005. Organic compounds, with CH4, are always present (<0.05% H2O), and are twice as abundant in cloudy halite. There are also traces of nitrogen and, in some samples, hydrogen sulphide and sulphur dioxide (Zimmermann and Moretto, 1996).


The influence of overpressure driving changes in the dihedral angle of pressurised salt is most clearly seen in black-salt encased Late Neoproterozoic to early Cambrian intra-salt Ara (stringer) reservoirs of the South Oman Salt Basin (Figures 2, 4, 5; Kukla et al., 2011b). These carbonate bodies are isolated in salt and frequently contain low-permeability dolomites and are characterised by high initial hydrocarbon production rates due to overpressure. But not all stringers are overpressured, and a temporal relationship is observed defined by increasingly overpressured reservoirs within stratigraphically younger units. There are two separate pressure trends in the stringers; one is hydrostatic to slightly-above hydrostatic, and the other is overpressured from 17 to 22 kPa.m−1, almost at lithostatic pressures (Figure 4).


The black staining of the halite is caused by intragranular microcracks and grain boundaries filled with solid bitumen formed by the alteration of oil (Figures 2, 5). The same samples show evidence for crystal plastic deformation and dynamic recrystallization. Subgrain-size piezometry indicates a maximum differential paleostress of less than 2 MPa. Under such low shear stress, laboratory-calibrated dilatancy criteria suggest that oil can only enter the rock salt at near-zero effective stresses, where fluid pressures are very close to lithostatic. In Schoenherr et al.’s (2007b) model, the oil pressure in the carbonate stringer reservoirs reservoir increases until it is equal to the fluid pressure in the low, but interconnected, porosity of the Ara Salt, plus the capillary entry pressure (Figure 5). When this condition is met, oil is expelled into the rock salt, which dilates and increases its permeability by many orders of magnitude. Sealing capacity is lost, and fluid flow will continue until the fluid pressure drops below te minimal principal stress, at which point rock salt will reseal to maintain the fluid pressure at lithostatic values. Inclusion studies in the halite indicate ambient temperatures at the time of entry were more than 90°C, implying hydrocarbons could move into interconnected polyhedral tubes in the halite. These conduits were created in response to changes in the polyhedral angle in the halite in response to elevated temperatures (Lewis and Holness, 1996).


Hydrocarbon-stained “black salt” can extend up to 100 metres from the pressurised supplying stringer into the Ara salt of Oman (Figure 2, 5). It indicates a burial-mesogenetic pressure regime and is not the same process set as seen in the telogenetic “black salt” regions of the onshore Gulf of Mexico. The latter is created by dissolution, meteoric water entry, and clastic contamination, as in the crests of nearsurface diapirs such as Weeks Island (Warren 2015). An Ara stringer enclosed by oil-stained salt but now below the lithostatic gradient likely indicates a later deflation event that caused either complete (C) or partial (E) loss of overpressures. Alternatively, stringers showing overpressure, but below the lithostatic gradient (E), might be explained by regional cooling or some other hitherto unexplained mechanism (Figure 4a; Kukla et al., 2011a, b).

Structural, petrophysical and seismic data analysis suggests that overpressure generation in the Ara is driven initially by rapid burial of the stringers in salt, with a subsequent significant contribution to the overpressure from thermal fluid effects and kerogen conversion of organic-rich laminites with the stringer bodies. If the overpressured stringers come in contact with a siliciclastic minibasin, they will deflate and return to hydrostatic pressures (A) in Figure 4. When the connection between the minibasin and the stringers is lost, they can regain overpressures because of further oil generation and burial (A’). If hydrocarbon production in undeflated stringers stops relatively early, the fluid pressures do not reach lithostatic pressures (B). If hydrocarbon generation continues, the fluid pressures exceed the lithostatic pressure (red star), leading to dilation and oil expulsion into the rock salt to what is locally known as “black salt” (D and E).

As well as these examples of overpressure associated with older evaporites, overpressure readily develops in salt-sculpted Tertiary basins. For example, overpressure occurs in salt shear (gumbo) transitions beneath some, but not all, shallow salt allochthons in Green and Mahogany Canyon regions in the Gulf of Mexico (Beckman, 1999: Shaker 2008). Where salt allochthons are climbing the stratigraphy, subsalt sealing and associated overpressure can occur beneath the salt mass at shallower levels than is observed in overpressured shale basins.

In terms of extension and compression regimes within a single allochthon tongue, Shaker (2008) noted that in extensional regions in halokinetic basins the magnitude and direction of the principal stresses are controlled by sediment load, salt thickness, and salt emplacement-displacement history. Therefore, the maximum principal stress is not necessarily represented by the sheer weight of the overburden, as is usually assumed in quiescent terranes. Salt buoyancy often acts upward and has the tendency to accelerate and decelerate the principal stress above and below the salt, respectively. A distinctive shift of the pore pressure envelopes and normal compaction trends takes place across the salt body in several wells drilled trough salt below minibasins in the Mississippi Canyon, Green Canyon, and Garden Banks areas of the Gulf of Mexico. A lower pore pressure gradient has been observed below the salt and a higher gradient above the salt barrier. On the salt-rooted minibasin scale, a high-gradient was also observed in areas where the salt was emplaced and a lower gradient where the salt withdrew (Shaker and Smith, 2002). On the other hand, in the compressional portion of a salt allochthon system, lateral stress generated by the salt movement piling up salt at the foot of the slope acts as the maximum principal stress, whereas the load of sediment represents the minimum stress.


Extreme overpressuring is commonplace in subsalt settings in the Gulf of Mexico at depths of 3000-4000 m and its variability creates drilling problems, as evidenced by the BP Horizon spill and explosion on April 20, 2010. Gas generated at greater depths in these regions can be trapped under the salt seal at pressures approaching lithostatic. It means drilling under the allochthonous salt on the Gulf Coast slope can intersect undercompacted sediments that are moderately to extremely overpressured and friable (Hunt et al., 1998). The influence of highly effective Jurassic salt seals on pressure gradients in the Neogene stratigraphy of the Gulf of Mexico is seen in the increased mud weights typically required for safe drilling, once an evaporite allochthon is breached by the drill (Table 1). Many wells intersecting salt allochthons in the deepwater realm of the Gulf of Mexico and the circum-Atlantic Salt basins are overpressured at some depth below the base of salt with mud weights controlling pressures ranging from 14 to 17.5 ppg.

Implications

This and the previous article (Warren, 2015) demonstrate that black salt is a form of anomalous salt that indicates salt has leaked, however, the locations and conditions where leakage has occurred are distinct. The black salt encountered in the salt mines of the US Gulf Coast are indicative of meteoric water entry in relatively shallow conditions in regions where the salt is in contact with the surrounding shales of muds that enclose the diapir salt core. In other words, fluid entry is from the outside of the salt mass and fluids move into the salt from its edges and likely enhance  the porosity in the intercrystalline salt. In contrast, the black salt occurrences in the Ara Salt of Oman are indicative of overpressure haloes, generated internally via hydrocarbon and fluid expulsion in carbonate slivers, which are are fully encased in salt. This creates naturally hydrofractured envelopes in the salt mass in zones where pressure and temperature induced changes in the dihedral angle has generated intercrystalline fluid strings within the recrystallised polyhedral halite. The two settings of black salt formation are distinct.

There is not a single mechanism that creates black salt in a halokinetic salt mass. We shall discuss the implications of this in the next article which will include a look at leakage models in halokinetic salt systems both in terms of their seal integrity and the implications for short and  long term storage of hydrocarbons and nuclear waste. 

References

Beckman, J., 1999. Study reveals overpressure sources in deep-lying formations. Oil and Gas Journal, September: 137.

Holness, M.B. and Lewis, S., 1997. The structure of the halite-brine interface inferred from pressure and temperature variations of equilibrium dihedral angles in the halite-H2O-CO2 system. Geochimica et Cosmochimica Acta, 61(4): 795-804.

Hunt, J.M., Whelan, J.K., Eglinton, L.B. and Cathles III, L.M., 1998. Relation of shale porosities, gas generation, and compaction to deep overpressures in the US Gulf Coast. In: B.E. Law, G.F. Ulmishek and V.I. Slavin (Editors), Abnormal pressures in hydrocarbon environments. American Association Petroleum Geologists Memoir 70, Tulsa, OK, pp. 87-104.

Kukla, P., Urai, J., Warren, J.K., Reuning, L., Becker, S., Schoenherr, J., Mohr, M., van Gent, H., Abe, S., Li, S., Desbois, Zsolt Schléder, G. and de Keijzer, M., 2011a. An Integrated, Multi-scale Approach to Salt Dynamics and Internal Dynamics of Salt Structures. AAPG Search and Discovery Article #40703 (2011).

Kukla, P.A., Reuning, L., Becker, S., Urai, J.L. and Schoenherr, J., 2011b. Distribution and mechanisms of overpressure generation and deflation in the late Neoproterozoic to early Cambrian South Oman Salt Basin. Geofluids, 11(4): 349-361.

Lewis, S. and Holness, M., 1996. Equilibrium halite-H2O dihedral angles: High rock salt permeability in the shallow crust. Geology, 24(5): 431-434.

O'Brien, J. and Lerche, I., 1994. Understanding subsalt overpressure may reduce drilling risks. Oil and Gas Journal, 92(4): 28-29,32-34.

Peach, C. and Spiers, C.J., 1996. Influence of crystal plastic deformation on dilatancy and permeability development in synthetic salt rock. Tectonophysics, 256: 101-128.

Schenk, O. and Urai, J.L., 2004. Microstructural evolution and grain boundary structure during static recrystallization in synthetic polycrystals of sodium chloride containing saturated brine. Contributions to Mineralogy and Petrology, 146: 671-682.

Schoenherr, J., Littke, R., Urai, J.L., Kukla, P.A. and Rawahi, Z., 2007a. Polyphase thermal evolution in the Infra-Cambrian Ara Group (South Oman Salt Basin) as deduced by maturity of solid reservoir bitumen. Organic Geochemistry, 38(8): 1293-1318.

Schoenherr, J., Urai, J.L., Kukla, P.A., Littke, R., Schleder, Z., Larroque, J.-M., Newall, M.J., Al-Abry, N., Al-Siyabi, H.A. and Rawahi, Z., 2007b. Limits to the sealing capacity of rock salt: A case study of the infra-Cambrian Ara Salt from the South Oman salt basin. Bulletin American Association Petroleum Geologists, 91(11): 1541-1557.

Shaker, S., 2008. The double edged sword: The impact of the interaction between salt and sediment on subsalt exploration risk in deep water. Gulf Coast Association of Geological Societies Transactions, 58: 759-769.

Warren, J.K., 2015. Salt as a fluid seal: Article 1,  Salty Matters blog; First published on Dec 19, 2015; www.saltworkconsultants.com.

Warren, J.K., 2016. Evaporites: A compendium (ISBN 978-3-319-13511-3) Released Feb. 2016. Springer, Berlin, 1854 pp.

Zimmermann, J.L. and Moretto, R., 1996. Release of water and gases from halite crystals. European Journal of Mineralogy, 8(2): 413-422.


 

 

Salt as a Fluid Seal: Article 1 of 4: External fluid source

John Warren - Saturday, December 19, 2015

 

Introduction

In the next few articles, I plan to discuss salt’s ability to act as a fluid seal in a variety of halokinetic settings, as well as looking at the nature of the sealing salt. Of particular interest are formative mechanisms driving textural and permeability variations in zones that typify the salt side of the sealing boundaries in sub-vertical salt stems versus the lower contact transitions in sub-horizontal allochthons. In the first few articles, we shall focus on local-scale scenarios and salt seal textures, including situations where salt has leaked, and the intercrystalline or tetrahedral/polyhedral pores contain fluid or mineralogical evidence of leakage and crystal boundary dissolution. Within the salt mass, this is usually indicated by occurrences of “black” or “dark” salt in anomalous salt zones, some of which are intersected by workings in a number of salt mines. In contrast, in most oil exploration scenarios we only have wireline signatures to interpret the deeper and typically offshore seal horizons. Following on from that discussion, we shall look at more regional examples of cross-formational leakage. Finally, we will discuss implications of possible leakage in terms of understanding and predicting outcomes with respect to both waste storage and hydrocarbon sealing

“Black” or “dark” salt in anomalous salt zones

The geological term “black salt” covers a variety of salt textures and associated mechanisms of formation. The term “black” salt also has a non-technical culinary association (kala namak[1]) but, other than in the footnote, I will not discuss it further in this series of articles. The geological descriptor “black” or “dark” salt is widely used in the US salt mining industry as a pointer to possible zones of current or past natural fluid entry into the salt mass. Colouring fluids can be brine, oil or gas, often with solid impurities dominated by shale, anhydrite or calcite-dolomite. These intrasalt “black” or “dark” salt zones in a mine were also referred to as “shear” zones and considered pointers to what are often unstable regions, liable to fluid entry, gassy outbursts and roof or wall collapse. “Shear”, “black” and “dark” salt zones are better described under the broader term “anomalous salt” zones, many of which were or are  in fluid contact with the enclosing non-salt sediment mass (Kupfer, 1990).

In a somewhat related fashion, the term “black salt” is used by the oil industry in Oman and Europe to indicate subsurface zones of overpressured salt, where natural hydrofracturing has occurred, and hydrocarbons have penetrated up to 100 m into the sealing salt mass. Hence its dark color (naturally hydrofractured salt and its textures are the focus of the second article in this series on salt leakage). Fluid entry in this type of “black’ salt is ascribed to temperature-related changes in the dihedral angle of the halite crystals in “black” salt zones. In a similar fashion, the term “black” salt was used in a recent paper in Science by Ghanbarzadeh et al. (2015) and the dihedral angle changes are ascribed to temperature increases in halokinetic salt intervals in the offshore Gulf of Mexico. There the authors argue temperature increases have changed the intercrystalline dihedral angle in a salt mass, and so facilitated the entry of fluids from adjacent strata into the salt body.

So, the term “black” salt is used in the geological community without reference to geological criteria that can separate what I consider are at least two distinct styles of “black” or “dark” salt formation and leakage. One type of salt leakage occurs when the salt is relatively shallow and subject to dissolution driven by the entry of meteoric and other near-surface undersaturated waters into folded and refolded shear (anomalous) salt zones in and about salt stems and décollements. This typically occurs when the flowing salt crest is relatively shallow and tends to occur in regions where the leaking “black” salt zone is in contact with the nonsalt boundary edges of the halokinetic salt mass. This process set ultimately leads to an accumulation of insoluble residues (clays, anhydrite, gypsum, calcite, etc.) that define a unit called a caprock. The term “cap” is somewhat of a misnomer as “caprock” units also form on the sides and undersides of a salt mass, wherever the salt unit is in contact with undersaturated cross-flowing formational waters (Warren, 2016). The other type of “black salt, exemplified by the Ara salt in Oman is related to deeper salt burial, salt flow and an association with intrasalt pressurized fluids (a focus of next article in this series on salt leakage). Accordingly, if we are not to confuse styles of “black” salt genesis (meteoric or undersaturated fluid entry versus intrasalt overpressures) then a better non genetic term should be used to describe zones of "black" or "dark" salt. Although less euphonious, the better term is “anomalous” salt. This describes all zones within halite-dominant intervals with features that are not typical of the bulk of the main diapiric salt mass (Kupfer, 1990).

In this first article we look at various types of anomalous salt in salt mines, largely related to the entry of, or interaction with, undersaturated relatively shallow formation waters. The next article focuses on salt leakage and "black" salt related to overpressure. Then, as we shall see in the third article on salt leakage, there are significant implications of occurrences of anomalous salt with respect to practicalities of safe intrasalt storage and fluid contamination with respect to separating the two types of black salt. This is especially so when working in the subsurface without the luxury of core or mine wall exposures. Ignoring the origins of “black” or “dark” salt, and the associated implications for wireline interpretation, means any conclusions in terms of waste storage outcomes and/or hydrocarbon seal potential, by generalizing lab-based experimental results on leaking salt to all “black” salt occurrences in halokinetic settings, will be somewhat confused (e. g. Ghanbarzadeh et al., 2015).

Black salt (dark salt) in anomalous salt in response to undersaturated fluid entry

Intervals of “black” or dark salt are described in US Gulf Coast salt mines in publications by Balk (1953), Kupfer (1976, 1990) and Looff et al., (2010), the following observations are largely based on their work. Nearsurface (<1-2 km) portions of mined or cored diapirs with “black” salt zones in the Gulf Coast USA and Germany are segmented into a number of intradiapir zones showing differential movement between adjacent salt spines or flowing masses. The more homogenous intervals of consistently mineable salt ore are separated by anomalous zones, formerly called “shear” zones. This association of homogenous spines separated by narrower shear or anomalous zones was first mapped in mine walls in the Jefferson Island salt dome by Balk (1953). His work was one of a series of classic papers mapping the internal structural complexities and shears in various mined salt diapirs in the US Gulf Coast and the Zechstein of Germany. Subsequent work by Kupfer (1976) on the same US Gulf Coast Five-Island salt mines (Jefferson Island, Avery Island, Weeks Island, Cote Blanche and Belle Isle) further refined notions of internal shear and occurrences of “black” or “dark” salt in diapirs. Today, only the Cote Blanche and Avery Island salt mines are still in operation along the Five Island Salt Dome Trend (Figure 1)

A shear zone in a diapiric structure forms where adjacent parts of a salt structure are moving (rising or falling) at different rates. Such zones tend to dominate the perimeter of a salt structure across which salt mass is rising or falling with respect to the adjacent sediment and so grade outward from the salt spine into a boundary “shale sheath”. Older shear zones and shale sheaths also are commonplace in re-folded intervals within a salt stem. Mapping of these zones by Balk (1953), Muehlberger and Claubaugh (1968) and Kupfer (1976) across many salt mines showed salt in a diapir must flow at different rates at different times. Otherwise, the complex and highly variable internal refolded drape and napkin folds seen in diapirs in all the world’s salt mines could not form. Figure 2 illustrates some internal complexities the diapir scale typifying various diapiric salt structures across the world and the dominantlyvertical flow fabrics in diapir stems and subhorizonatal flow textures in overhangs and salt tongues. Figure 3 shows that same vertical dominance (biaxial elongation) of salt crystals from cores collected in diapir stems cored various salt mines, while Figure 4 shows the typical vertical banding fold style that typifies diapir stems.



Walden and Jacoby (1963) were the first to call attention to a Gulf Coast anomalous salt zone. They documented a fault zone in the Avery Island salt mine that separated the region of salt being mined, across an anomalous zone, from the domal core. To call attention to the zonal ductile, not brittle, nature of intradiapir salt flow, Kupfer, 1974 changed the description of such anomalous zones from "fault” zones to "shear” zones and concluded most intradiapir shear zones were not faulted zones (defined by brittle fracture offset). In a later paper, he suggested abandoning of the genetic and misleading term "shear zone" and proposed replacement with the broader nongenetic term "anomalous salt zone" (Kupfer, 1990).


The term “anomalous salt,” as defined by Kupfer (1990), is based on his then more than twenty years experience in various salt mines in the US Gulf Coast. An anomalous salt zone is defined broadly as a zone of anomalous features in salt of whatever origin. He noted that typical anomalous salt zone features are different to the majority of features in the adjacent salt and involve varying combinations of anomalous features that include:

Textures--Coarse-grained, piokiloblastic, friable

Inclusions--Sediments, hydrocarbons, brine, gases

Structures--Sheared salt, gas outbursts, brine leaks, undue roof and wall slabbing, jointing, voids, and slight porosity development

Compositions--Potash/magnesium, high anhydrite content, very black salt (made up of disseminated fluid and solid impurities.

The terms “anomalous salt” and “anomalous zones” as defined by Kupfer (1990), are based on observations across the various Five Island salt mines of South Louisiana (Figure 1). As later refined in Kupfer et al. (1998), anomalous salt is a rocksalt zone that deviates from what are considered typical domal salt. Typical Gulf Coast rocksalt according to Kupfer is reasonably pure halite (97%+/- 2%), with minor amounts of disseminated anhydrite (CaSO4) being the primary non-halite impurity. Grainsize is considered to be uniform with grain diameters of 3 – 10 mm (0.12 – 0.39 inches). With continued mapping of Five Island mines, Kupfer et al. (1998) and Looff et al. (2010) documented an even wider variety of anomalous salt zone characteristics and concluded that the creation of anomalous zones need not be related to faulting or shearing, but also can be related to fluid entry and salt dissolution (Figure 5). Anomalous salt can be defined by impurity content, structure, colour, or other features. Anomalous features may not have sharp contacts or uniform thickness, and most are not continuous over long distances. Individual anomalous features commonly disappear for tens of metres (hundreds of feet) only to appear over some horizontal distance. The internal salt fabric of a salt dome is always composed of both typical (volumetrically dominant) and anomalous salt. Kupfer (op. cit.) noted that other salt deposits, including horizontally bedded nonhalokinetic salt deposits in the Permian of West Texas and the Devonian of Western Canada, all have anomalous zones of various origins.


Further work in both the salt mines and salt cavern storage industry has increasingly invoked the concept of anomalous features, anomalous zones and boundary shear zones although there is still a significant confusion over the appropriate use of the terminology (Looff et al., 2010). Because of the flow experienced by diapiric salt, most anomalous salt features parallel the near vertical internal banding of the salt. Many anomalous salt features may create zones of differing creep, strength, or dissolution characteristics that can impact the solution mining and operation of a salt storage cavern and some are tied to zones of problematic fluid entry in a mine. An anomalous zone is any zone in a salt diapir that contains 3 or more dissimilar anomalous features (Kupfer, 1990). The term “anomalous” implies nothing regarding the genesis of the zone. While many anomalous zones may extend laterally over hundreds of metres in length, they are variable in nature, near vertical, and parallel to layering (Figure 5). Typical widths are poorly known but are commonly in the order of 30-50m; however individual structures or anomalous features within the anomalous zone may be as thin as millimetres.

Boundary Shear Zones (BSZ) and Edge Zones (EZ) are the two types of anomalous zones that have a genetic interpretation (Looff et al., 2010). Boundary shear zones are those zones that bound an active salt spine where the salt experiences increased shear stress due to differential salt movement. An edge zone is similar to a boundary shear zone except, instead of being internal within the dome, it is confined to the periphery of the salt stock. Anomalous salt is not restricted to shear zones, however within and about as diapir edge one can expect most anomalous salt to be associated with shear zones (Kupfer, 1990; Looff, 2000).

Anomalous zones within a salt spine are in many cases the remnants of relict BSZ’s from older spines incorporated into younger active salt spines and this especially obvious with those boundary zones associated with clastic impurities (Figure 6). Boundary shear zones and edge zones around the dome tend to be more problematic for storage caverns as they are likely to contain greater occurrences of anomalous salt, higher impurity content (including gas and brine) and structural features that may degrade salt quality and enable leakage. Thus salt caverns can be constructed within boundary salt zones, but if possible, they should be avoided as they can result in non-optimal operating conditions, long-term operational difficulties and in the most severe cases contribute to the loss of cavern integrity (Looff et al., 2010). In the case of edge zones, additional distance to the edge of the salt dome needs to be maintained not only to cover any uncertainty regarding the placement of the edge of salt with respect of mine workings but also to account for the potential for degraded salt quality and to provide a sufficient pillar of good quality salt between the mine or cavern wall and the edge of dome.


A top-of-salt boundary between aggradational and dissolutional components atop diapirs in the Five Islands salt landscape typically coincides with underlying anomalous zones of differential shear within the underlying diapir typically indicated by “black” or “dark” salt zones in the various diapirs (Kupfer, 1976; Lock, 2000). Where such interior anomalous “black” salt zones have intersected the edge of the salt mass, they tend to create intervals with a greater propensity for water entry or gas outbursts and unstable roof zone liable to slabbing and collapse (Figure 6). Such anomalous zones can leak water into a mine, and over the longer term create stability problems, as illustrated by problems in; the now abandoned Weeks Island oil storage facility, the Avery Island Salt Mine, and the likely association of a subvertical zone with anomalous salt, and the enhanced fluid entry that occurred during the Lake Peigneur collapse, which was tied to 1980 flooding of the former Jefferson Island Salt Mine. Today, only two of the former mines in the Five Island Salt Dome trend remain unflooded. For a more detailed discussion of these and other salt leakage scenarios tied to undersaturated fluid entry into salt mines and caverns, see case histories in Chapter 13, Warren 2016.

“Black” or “Dark” Salt zones and leakage into the former Weeks Island storage facility

In the walls of the now-flooded Weeks Island salt mine, Kupfer (1976) noted that wide black beds of the internal “shear” zone are unusual and not found over most of the rest of the mine where salt was extracted. In places, the anomalous zone beds contain black clay (Room J-21), orange sandstone (S-20), and other fragments of clastic material (Paine et al., 1965). These clastic remnants typically occur as balls or roundish blebs ranging in size up to tens of cm in diameter. Petroleum leaked out of seams in this black salt zone and seams in the surrounding salt; the escaping fluid ranged from yellow grease and heavy, blue oil to very light, straw-yellow distillate. Methane and carbon dioxide were also common. The width (surmised) and structural complexity of the anomalous zone suggest that internal salt movement continued after a clastic boundary sheath-zone was incorporated into the salt stock (Figure 7).


The cause of the drainage and abandonment of the Weeks Island oil storage facility was an active subsidence sinkhole some 10 metres across and 10 metres deep, first noted near the edge of the SPR facility in May 1992, and perhaps reaching the surface about a year earlier. The growing doline depression was located on the south-central portion of the island, directly over a subsurface trough, which was obvious in the top-of-salt contours based on former mine records before conversion to a hydrocarbon storage facility (Figure7; Neal and Myers, 1995). Earlier shallow exploratory drilling around the Department of Energy service and production shafts in 1986 had identified the presence of irregularities and brine-filled voids along the top of salt mass across this region. A second, much smaller sinkhole was noticed in early February 1995, but it did not constitute a serious threat as it lay outside the area of cavern storage.

The first sinkhole occurs in a position of sharp change in landform slope (transition from high island to gully fill) and lies atop the projected alignment of what is known as Shear Zone E (a dark salt zone) in the underlying salt (Autin and McCulloh, 1995). Neal (1994) pointed out that Kupfer’s 1976 map of that part of the Weeks Island salt mine, located beneath the first sinkhole, was defined by black salt (also shown as Figure 8 which is based on the more recent Kupfer et al. (1998) map). Miners always avoided such “black” salt or “dark” salt zones in the various subsurface workings and the lateral extent of workings in many of the Five Island mines extended only as far as intersections with significant “black” or “dark” salt regions (Figure 6 & 7).


The volume of the first Weeks Island sinkhole (estimated as 650 m3 when first noted), its occurrence over a trough in the top of salt, and its position directly above the oil-filled mine caverns, meant it was of urgent concern to the SPR authorities, especially in terms of the stability of the roof of the storage cavern. This feature did not form overnight; it lies atop a shear zone that formed during the diapiric rise of the salt and capped by a rockhead valley containing Pleistocene sediment fill. Salt extraction during mine operations probably created tension across the shear zone, thereby favouring fracture enlargement in the anomalous salt zone, as early perhaps as 1970 (Figure 6; Waltham et al., 2005). Eventually, an incursion of undersaturated groundwater traversed the fracture zone across some 107 m, from a level equivalent to the rockhead down to the mine where it emerged. Over time, ongoing dissolution enlarged a void at the top of the anomalous salt zone, creating the collapse environment for the sinkhole first noted at the land surface in 1991. Investigations were undertaken in 1994 and 1995 into the cause of active at-surface sinkholes verified that water from the aquifer above the Weeks Island salt dome was seeping into the underground oil storage chamber at the first sinkhole site (Figures 7& 8; Neal and Myers, 1995; Neal et al., 1995, 1997). Drainage and decommissioning of the Weeks Island facility followed.

Beginning in 1994, and continuing until the abandonment of the facility, saturated brine was injected directly into the throat of the first sinkhole, which lay some 75 metres beneath the surface. This essentially arrested further dissolution and bought time for DOE (Department of Energy) to prepare for the safe and orderly transfer of crude oil to another storage facility. To provide added insurance during the oil transfer stage, a “freeze curtain” was constructed in 1995. It consisted of a 54 well installation around the principal sinkhole, which froze the overburden and uppermost salt to a depth of 67 metres (Figure 9; Martinez et al., 1998). Until the mine was filled with brine and its hydrocarbons removed, this freeze wall prevented groundwater flow into the mine via the region of black salt around the sinkhole. Dealing with this sinkhole was costly. Mitigation and the removal and transfer of oil, including the dismantling of infrastructure (pipelines, pumps, etc.), cost a total of nearly US$100 million; the freeze curtain itself cost nearly $10 million.


Following oil fill in 1980-1982, the Weeks Island facility had stored some 72.5 million BBL of crude oil in abandoned mine chambers. Then in November 1995, the Department of Energy (DOE) initiated oil drawdown procedures, along with brine refill and oil skimming, plus numerous plugging and sealing activities. In 1999, at the end of this recovery operation, about 98% of the crude oil had been recovered and transferred to other SPR facilities in Louisiana and Texas; approximately 1.47 MMBL remains in the now plugged and abandoned mine workings. In hindsight, based on an earlier leak into the mine, while it was an operational mine, and the noted presence of black salt in a shear zone in the mined salt, one might fault the initial DOE decision to select this mine for oil storage. In 1978 groundwater had already leaked into a part of the mine adjacent to the sinkhole and this was forewarning of events to come (Martinez et al., 1998). Injection of cement grout into the flow path controlled the leak into the operation mine at that time, but it could just as easily have become uncontrollable and formed a sinkhole then.


Black salt zones in the now-flooded Jefferson Island Salt Mine and the 1980 Lake Peigneur collapse

The most recently risen part of the Jefferson Island stock crest is now 250 m (800 ft) higher than the adjacent flat-topped salt mass, which is also overlain by a cap rock (Figure 10). The boundary separating the spine from the less active portion of the crest is a finer-grained and a “shale-rich” anomalous zone, penetrated by the former Jefferson Island mine workings. It defined a limit to the extent of salt mining in the diapir, which was focused on extracting the purer salt within the Jefferson Island spine. The spine and its boundary “shear” are reflected in the topography of the Jefferson Island landscape, with a solution lake, called Lake Peigneur, defining the zone of shallower salt created by the active spine. There on November 20, 1980, one of the most spectacular sinkhole events associated with oilwell drilling occurred atop the Jefferson Island dome just west of New Iberia. Lake Peigneur disappeared as it drained into an underlying salt mine cavern and a collapse sinkhole, some 0.91 km2 in area, developed in the SE portion of the lake (Figure 11; Autin, 2002; Warren 2016). In the 12 hours following the first intersection the underlying mine had flooded and the lake was completely drained. The lake is about 2.4 km in diameter, has a bean-shaped configuration, with a topographic promontory along the southeast shore of the lake rising to more than 23 m above sea level and the surrounding delta plain (Figure 10).

Drainage and collapse of the lake began when a Texaco oilrig, drilling from a pontoon in the lake, breached an unused section of the salt mine some 1000 feet (350 metres) below the lake floor (Figure 11). Witnesses working below ground described how a wave of water instantly filled an old sump in the mine measuring some 200 ft across and 24 feet deep. This old sump was in contact with a zone of anomalous “black” (shear zone) salt. The volume of floodwater engulfing the mine corridors couldn’t be drained by the available pumps. At the time of flooding the mine had four working levels and one projected future level. The shallowest was at 800 feet, it was the first mined level and had been exploited since 1922. The deepest part of the mine at the time of flooding was the approach rampways for a planned 1800 foot level. Some 23-28 million m3 of salt had been extracted in the preceding 58 years of mine life. The rapid flush of lake water into the mine, probably augmented by the drainage of natural solution cavities in the anomalous salt zone and associated collapse grabens below the lake floor, meant landslides and mudflows developed along the perimeter of the Peigneur sinkhole, so that post flooding the lake was enlarged by 28 ha.


With water filling the mine workings, the surface entry hole in the floor of Lake Peigneur quickly grew into a half-mile-wide crater. Eyewitnesses all agreed that the lake drained like a giant unplugged bathtub—taking with it trees, two oil rigs (worth more than $5 million), eleven barges, a tugboat and a sizeable part of the Live Oak Botanical Garden. It almost took local fisherman Leonce Viator Jr. as well. He was out fishing with his nephew Timmy on his fourteen-foot aluminium boat when the disaster struck. The water drained from the lake so quickly that the boat got stuck in the mud, and they were able to walk away! The drained lake didn’t stay dry for long, within two days it was refilled to its normal level by Gulf of Mexico waters flowing backwards into the lake depression through a connecting bayou (Delcambre Canal, aka Carline Bayou) former what was a waterfall with the highest drop in the Stat of Louisiana. Since parts of the lake bottom had slumped into the sinkhole during the collapse, the final water level in some sections of the lake was higher than before relative to previous land features. This ground movement and subsidence left one former lakefront house aslant under 12 feet of water.

Implications for other salt mines with anomalous salt zone intersections.

The Peigneur disaster had wider resource implications as it detrimentally affected the profitability of other salt mines in the Five Islands region (Autin, 2002). Even as the legal and political battles at Lake Peigneur subsided, safe mining operations at the nearby Belle Isle salt mine came into contention with public perceptions questioning the structural integrity of the salt dome roof. During ongoing operations, horizontal stress on the mineshaft near the level where the Louann Salt contacts the overlying Pleistocene Prairie Complex across a zone of anomalous salt had caused some mine shaft deterioration. Broad ground subsidence over the mine area was well documented and monitored, as was near continuous ground water leakage into the mine workings. The Peigneur disaster meant an increased perception of continued difficulty with mine operations and an increased risk of catastrophic collapse was considered a distinct possibility. In 1985, a controlled flooding of the Belle Isle Salt m\Mine was completed as part of a safe closure plan.

Subsidence over the nearby Avery Island salt mine (operated by Cargill Salt) has been monitored since 1986 when small bead-shaped sinkholes were initially noticed in the above mine region. Subsidence monitoring post-1986 defined a broad area of bowl-shaped subsidence, within associated areas of gully erosion (Autin, 2002). Avery mine is today the oldest operating salt mine in the United States and has been in continual operation since the American Civil War. The mine underwent a major reconstruction and a improved safety workover after the Lake Peigneur disaster. Subsidence is still occurring today along the active mine edge, which coincides with a topographic saddle above an anomalous salt zone, which is located inside the mined salt area. At times, ground water has seeped into the mine, and there are a number of known soil gas anomalies and solution dolines on the island. These are natural features that predate mining. Much of the subsidence on Avery Island is a natural process as differential subsidence occurs atop any shallow salt structure with the associated creation of zones of anomalous salt (Warren, 2016). Dating of middens and human artifacts around salt-solution induced, water-filled depressions atop the dome, shows dissolution-induced subsidence is a natural process, as are short episodes of lake floor collapse, slumping and the creation of water-filled suprasalt dolines (circular lakes). Such landscape events and their sedimentary signatures have histories that extend back well beyond the 3,000 years of human occupation documented on Avery Island (Autin, 2002).

Compared to the other salt domes of the Five Islands region of Louisiana, the Cote Blanche Island salt mine has benefited from a safe, stable salt mine operation throughout the mine life (Autin, 2002). Reasons for this success to date are possibly; (i) mining operations have not been conducted as long at Cote Blanche Island as other nearby domes, (ii) the Cote Blanche salt dome may have better natural structural integrity than other islands, thus allowing for greater mine stability (although it too has anomalous salt zones, a salt overhang, and other structural complexities), and (iii) the Cote Blanche Salt Mine is surrounded by more clayey (impervious) sediments than the other Five Islands diapirs, all with sandier surrounds, perhaps allowing for lower rates of undersaturated fluid crossflow and greater hydrologic stability.

Significance

And so, today, we know that anomalous salt zones near diapirs crests are often tied to subvertical fault or shear zones, and that many are also associated with the presence of past crossflows of undersaturated waters. Across the various US Gulf Coast mines (present and past) the anomalous (“shear”) salt zones within diapirs are known to be potential problematic leakage zones and so are avoided, if possible, during mining operations. This style of black salt distribution and the potential for intrasalt leakage must be taken into account when near-crestal and shallower portions of domes are to be utilised for any fluid or waste storage. Without an understanding of the significance of such “black” salt or anomalous salt layers, there are potential undefined leakage problems within some salt structures (Looff et al., 2010; Warren 2016).

References

 

Autin, W. J., 2002, Landscape evolution of the Five Islands of south Louisiana: scientific policy and salt dome utilization and management: Geomorphology, v. 47, p. 227-244.

Autin, W. J., and R. P. McCulloh, 1995, Quaternary geology of the Weeks and Cote Blanche islands salt domes: Gulf Coast Association of Geological Societies Transactions, v. 45, p. 39-46.

Balk, R., 1953, Salt Structure of Jefferson Island Salt Dome, Iberia and Vermilion Parishes, Louisiana: Bulletin American Association Petroleum Geologists, v. 37, p. 2455-2474.

Ghanbarzadeh, S., M. A. Hesse, M. Prodanović, and J. E. Gardner, 2015, Deformation-assisted fluid percolation in rock salt: Science, v. 350, p. 1069-1072.

Kupfer, D., 1976, Shear zones inside Gulf Coast salt stocks help to delineate spines of movement: Bulletin American Association of Petroleum Geologists, v. 60, p. 1434-1447.

Kupfer, D., 1980, Problems associated with anomalous zones in Louisiana salt stocks, USA, in A. H. Coogan, and H. Lukas, eds., Fifth Symposium on Salt (Hamburg, Germany, June 1978), v. 1: Cleveland OH, Northern Ohio Geological Society, p. 119-134.

Kupfer, D. H., 1974, Boundary shear zones in salt stocks: in Fourth Symposium on Salt. Northern Ohio Geological survey, v. 1, p. 215-225.

Kupfer, D. H., 1990, Anomalous features in the Five Islands salt stocks, Louisiana: Gulf Coast Association of Geological Societies Transactions, v. 40, p. 425-437.

Kupfer, D. H., B. E. Lock, and P. R. Schank, 1998, Anomalous Zones Within the Salt at Weeks Island, Louisiana: Gulf Coast Association of Geological Societies Transactions, v. 48, p. 181-191.

Lock, B. E., 2000, Geologic Mapping of Salt Mines in Salt Diapirs: Approaches and Examples from South Louisiana: Gulf Coast Association of Geological Societies Transactions, v. 50, p. 567-582.

Looff, K. M., 2000, Geologic and Microstructural Evidence of Differential Salt Movement at Weeks Island Salt Dome, Iberia Parish, Louisiana: Gulf Coast Association of Geological Societies Transactions, v. 50, p. 543-555.

Looff, K. M., K. M. Looff, and C. Rautman, 2010, Salt spines, boundary shear zones and anomalous salts: Their characteristics, detection and influence on salt dome storage caverns: Paper presented at Solution Mining Research Institute Spring 2010 Technical Conference, Grand Junction, Colorado, USA, 26-27 April 2010, 23 p.

Martinez, J. D., K. S. Johnson, and J. T. Neal, 1998, Sinkholes in Evaporite Rocks: American Scientist, v. 86, p. 38.

Muehlberger, W. R., and P. S. Clabaugh, 1968, Internal Structure and Petrofabrics of Gulf Coast Salt Domes: AAPG Memoir, v. 8, p. 90-98.

Neal, J. T., 1994, Surface features indicative of subsurface evaporite dissolution: Implications for storage and mining: Solution Mining Research Institute, Meeting paper, 1994 Spring meeting, Houston Texas.

Neal, J. T., S. Ballard, S. J. Bauer, B. L. Ehgartner, T. E. Hinkebein, E. L. Hoffman, J. K. Linn, M. A. Molecke, and A. R. Sattler, 1997, Mine-Induced Sinkholes Over the U.S. Strategic Petroleum Reserve (SPR) Storage Facility at Weeks Island, Louisiana: Geologic Mitigation Prior to and During Decommissioning, SAND96-2387A.: Presented at 6th Multidisciplinary Conference on Sinkholes and the Engineering & Environmental Impacts of Karst, Springfield, Missouri, April 6-9, 1997. Sandia National Laboratories, Albuquerque, NM.

Neal, J. T., S. J. Bauer, and B. L. Ehgartne, 1995, Sinkhole Progression at the Weeks Island, Louisiana, Strategic Petroleum Reserve (SPR) Site: Solution Mining Research Institute, Fall Meeting, San Antonio, Texas, October 1995. Sandia National Laboratories, Albuquerque, NM.

Neal, J. T., and R. E. Myers, 1995, Origin, Diagnostics, and Mitigation of a Salt Dissolution Sink-hole at the U,S. Strategic Petroleum Reserve Storage Site, Weeks Island Louisiana,: Sandia National Laboratories, Albuquerque, NM. Report Sandia SAND95-0222C Paper presented at the Fifth International Symposium on Land Subsidence, The Hague, October 1995. Proceedings of the Fifth International Symposium on Land Subsidence, IAHS Publ. No. 234.

Paine, W. R., M. W. Mitchell, R. R. Copeland Jr., and L. d. A. Gimbrede, 1965, Frio and Anahuac Sediment Inclusions, Belle Isle Salt Dome, St. Mary Parish, Louisiana: American Association Petroleum Geologists - Bulletin, v. 49, p. 616-620.

Walden, W., and C. H. Jacoby, 1963, Exploration by horizon­tal drilling at Avery Island, Louisiana, in A. C. Bersticker, ed., Symposium on Salt (First): Cleveland, OH, Northern Ohio Geo­logical Society, p. 367-376.

Waltham, T., F. Bell, and M. Culshaw, 2005, Sinkholes and Subsidence: Karst and Cavernous Rocks in Engineering and Construction: Berlin Heidelberg, Springer Praxis Books, 382 p.

Warren, J. K., 2016, Evaporites: A compendium (ISBN 978-3-319-13511-3) Released Jan-Feb. 2015: Berlin, Springer, 1854 p.

 

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[1] Pungent-smelling condiment Kala Namak (black salt) is widely used in South Asia, it consists primarily of sodium chloride with trace impurities of sodium sulphate, sodium bisulphate, sodium bisulphite, sodium sulphide. Kala Namak is also known as Himalayan Black Salt, Sulemani Namak, Bit Lobon , Kala Noon or as Bire Noon in Nepal. Its characteristic smell and taste is mainly due to its elevated sulfur content, which to a western nose is reminiscent of rotten eggs, largely due to the presence of greigite. The various iron impurities impart a brownish pink to dark violet colour to the coarse translucent crystals and, when ground into a powder, transform into a light purple to pink color.

Traditionally, mined salt was transformed from the raw natural form of salt into commercially-sold kala namak through a reductive chemical process. This heating transforms some of the naturally occurring iron oxidew and sodium sulfates in the raw salt into pungent hydrogen sulfide and sodium sulfide daughter products (along with greigite.[ The various sulphate salt impurities in the halite typify the partially recrystallised meteoric overprints that typify textures and structures in nearsurface salt residues in the Himalayan thrust belt (see Richards et al., 2015 for documentation of the geological and structural characteristics of this salt – this article can be downloaded from the publications page on this website).

Historically, the transformation of Himalayan thrust belt salt into kala namak involved firing the raw salt in a furnace for 24 hours, while sealed in a ceramic jar containing charcoal along with small quantities of harad seeds, amla, bahera, babul bark, or natron. The fired salt was then cooled, stored, and aged prior to sale. Kala namak is still prepared in this manner in northern India with production concentrated in Hisar district, Haryana. Although the kala namak can still be produced from natural salts with the required compounds, it is now common to now manufacture it synthetically using halite from non-Himalayan sources. This is done through combining sodium chloride with smaller quantities of sodium sulfate, sodium bisulfate and ferric sulfate, which are then chemically reduced with charcoal in a furnace. Reportedly, it is also possible to create similar products through reductive heat treatment of sodium chloride, 5–10% of sodium carbonate, sodium sulfate, and some sugar.


 

 

 

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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