Salty Matters

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

Stable isotopes in evaporite systems: Part II - 13C (Carbon)

John Warren - Thursday, May 31, 2018

 

Introduction

13C interpretation in most ancient basins focuses on carbonate sediment first deposited/precipitated in the marine realm. Accordingly we shall first look here at the significance of variations in 13C over time in marine carbonates and then move our focus into the hypersaline portions of modern and ancient salty geosystems. In doing so we shall utilize broad assumptions of homogeneity as to the initial distribution of 13C (and 18O) in the marine realm, but these are perhaps oversimplifications and associated limitations need to be recognized (Swart, 2015)

In the next article we shall look at the utility of crossplots of carbon and oxygen isotopes. Stable oxygen isotope values (d18O) crossplotted with respect to carbon isotope values (d13C) from, the same sample creates one of the most widely applied proxies used to infer palaeo-environmental conditions (depositional and diagenetic) in Holocene and ancient carbonate sediments. This is in large part due to kinetic fractionations that occur during evaporation (Leng and Marshall, 2004). It has long been known that as any liquid evaporates, the residual fluid becomes enriched in the less abundant heavy isotope(s) (see Horton et al., 2016 for detailed discussion).

Interpreting 13C

Over the Phanerozoic the standard paradigm for interpreting variations in variations in 13C values from modern and ancient marine carbonate is based on an integration of our understanding of the carbon cycle with the following arguments. Most of the carbon in Earth’s near-surface systems is stored in sedimentary rocks with only about 0.1% in living organisms and the atmosphere-hydrosphere (Figure 1). Oxidized carbon occurs primarily as marine carbonates and reduced carbon as organic matter in sediments. In the carbon cycle, CO2 from the oceans and atmosphere is transferred into sediments as carbonate carbon (Ccarb) or organic carbon (Corg), the former of which monitors the composition of the oceans (Figure 1). The cycle is completed by uplift and weathering of sedimentary rocks and by volcanism, both of which return CO2 to the atmosphere.


There are two stable carbon isotopes, carbon 12 (6 protons and 6 neutrons) and carbon 13 (6 protons and 7 neutrons). Photosynthetic organisms incorporate disproportionately more CO2 containing the lighter carbon 12 than the heavier carbon 13 (the lighter molecules move faster and therefore diffuse more easily into cells where photosynthesis takes place). During periods of high biological productivity, more light carbon 12 is locked up in living organisms and in resulting organic matter that is being buried and preserved in contemporary sediments. Consequently, due the metabolic (mostly photosynthetic) activities of a wide variety of plants, bacteria and archaea, the atmosphere and oceans and their sediments become depleted in carbon 12 and enriched in carbon 13 (Figure 2)


It is assumed that the carbon isotopic ratio in calcareous shells of marine organisms is in equilibrium with that of seawater. So as more carbon 12 is held in biomass during times of high primary productivity, and increased burial of organic carbon, calcareous (CaCO3) skeletal materials become enriched in carbon 13. In contrast during periods of low biological productivity and decreased burial of organic carbon, for example following mass extinctions, marine calcareous skeletal materials become enriched in carbon 12.

Hence plotting variations in carbon isotopes in marine carbonates and organic matter over time offers a way to trace the growth of the crustal reservoir of reduced carbon (Des Marais, 1997). That is, the relative abundance of carbon isotopes is controlled chiefly by: 1) equilibrium isotopic effects among inorganic carbon species, 2) fractionation associated with the biochemistry of organic matter, and 3) the relative rates of burial of carbonate and organic carbon in sediments (Condie 2016).

Because organic matter preferentially incorporates 12C over 13C, there should be an increase in the 13C/12C ratio (as measured by δ13C) in buried carbon with time, and indeed this is what is observed (Des Marais, 1997; Worsley & Nance, 1989). δ13Corg increases from values < -40‰ in the Archaean to modern values of -20 to -30‰. On the other hand, seawater carbon as tracked with δ13Ccarb remains roughly constant with time, with δ13Ccarb averaging about 0%.

Variation in fluxes over time within the carbon cycle can be monitored by an isotopic mass balance (Des Marais, 1997), whereby;

δin = fcarbδ13Ccarb + forgδ13Corg

δin represents the isotopic composition of carbon entering the global surface environment comprised of the atmosphere, hydrosphere, and biosphere. The right side of the equation represents the weighted-average isotopic composition of carbonate (δ13Ccarb) and organic (δ13Corg) carbon buried in sediments, and fcarb and forg are the fractions of carbon buried in each form (fcarb = 1 - forg). For timescales longer than 100 Myr, δin = -5‰, the average value for crustal and mantle carbon (Holser et al., 1988). Thus, where values of sedimentary δ13Ccarb and δ13Corg can be measured, it may be possible to determine forg for ancient carbon cycles. Higher values of δ13Ccarb indicate either a higher value of forg or a greater negativity of average δ13Corg.


During the Phanerozoic, there are several peaks in δ13Ccarb, the largest at about 110, 280, 300, 400, and 530 Ma (Figure 3). These peaks are widely interpreted to reflect an increase in burial rate of organic carbon (Des Marais et al., 1992; Frakes et al., 1992). This is because organic matter selectively enriched in 12C depletes seawater in this isotope, raising the δ13C values of seawater. In the late Paleozoic (300-250 Ma), the maxima in δ13Ccarb correspond to the rise and spread of vascular land plants, which provided a new source of organic debris for burial (Condie 2106, Berner, 1987, 20 01). Also conducive to preservation of organic remains at this time were the vast lowlands on Pangea, which appear to have been sites of widespread swamps where bacterial decay of organic matter is minimized. The drop in δ13Ccarb at the end of the Permian is not understood. Perhaps, large amounts of photosynthetic O2 generated by Carboniferous forests led to extensive forest fires that destroyed large numbers of land plants in the Late Permian (Condie, 2016). However, the reasons for the oscillations in δ13Ccarb are not yet unequivocally resolved and, as in all sciences, the tenet "...perceived correlation does not necessarily equate to causation"must always be at the forefront in the scientific mindset.


Across the Precambrian and the Phanerozoic, the initiation of glaciation on a global scale, as in the Cryogenian ‘Snowball Earth’, has been interpreted to be dependent on parameters like the latitudinal extent of continents and oceanic circulations (Figure 4; Condie, 2016). The main drive for an onset of global glaciation is believed to be the lowering of atmospheric CO2. It likely also requires a continental landmass to be covering one of the earth's polar positions. More recently, cooling related to an increase in the earth's albedo due to widespread evaporites (saline giants) has been added to the list of possible drivers to the onset of glaciation.

Climate modelling studies imply that CO2 concentrations as low as 100–150 ppm are required to initiate global glaciation (e.g. Liu et al., 2013; Feulner and Kienert, 2014). One potential cause of lowered CO2 is drawdown of CO2 during intense silicate weathering in equatorial regions (Hoffman and Schrag, 2002; Goddéris et al., 2003). Photosynthesis provides another mechanism for CO2 drawdown, via conversion of CO2 to O2 and rapid burial of organic carbon, which is reflected in a positive δ13C excursion for carbonates (Pierrehumbert et al., 2011). Additionally, long term cloud cover (Feulner et al., 2015), fluctuations in atmospheric-ocean heat transport, the earth's albedo, or solar luminosity (Pierrehumbert et al. (2011) are also proposed as potential causes of the onset of glaciation (ice-house mode climate).

In a recent paper, Schmid 2017 focused on the cause of the Bitter Springs carbon isotope anomaly, she argues the cause of the pre-glacial, globally recognised, carbon and oxygen isotope variations in carbonate sediments tied to the Bitter Springs anomaly is a response to widespread fractional evaporation of dissolved CO2. This carbon isotope anomaly ties to a well defined correlation with the distribution of Neoproterozoic evaporite basins. She also shows volcanism occurred during the onset of the Bitter Springs Stage (811–788 Ma) and associated widespread evaporite distribution across Australia.


Schmid (op. cit.) argues that the albedo effect began with of the widespread deposition of Rodinian supercontinent evaporites in very shallow marine to epicontinental sedimentary successions beginning ≈810 Ma, increased siliciclastic redbed weathering. This and continuing evaporite deposition and exposure between ≈780 and 720 Ma drove a worldscale increase in Earth's albedo. Such highly reflective salt deposits defined a saline giant across an area that today covers one-third of the Australia continent. Thus, this and other penecontemporaneous saline giants over the Rodinian supercontinent played a potentially significant role in the onset of atmospheric cooling via a significant increase in albedo (Figure 5). These salt beds occur in periods that typify the onset of local (750 Ma) and then global glaciation (720 Ma).


Schmid (2017) goes on to note that the degree of evaporation in the Bitter Springs group sediments is related to the δ13C signature in variably concentrated waters (Figure 6). That is the Tonian Bitter Springs Group (≈830–750 Ma), within the Amadeus Basin in central Australia consists of thick halite and sulphate evaporite accumulations and associated carbonates. The deposition of halite occurred in shallow marine, lagoon (salina) environment (Gillen Formation), and developed into sulphate-dominated supratidal sabkha during sea level regression (Johnnys Creek Formation). The overall regression was interrupted by a transgressive phase lasting at least 20 Ma and leading to deposition of basin-wide stromatolitic dolostone (Loves Creek Formation). The salinity and high evaporation is reflected in positive δ13C in the intercalated carbonates (+4 to +6‰ VPDB) of the evaporitic units, while the shallow marine stromatolitic incursion of the Loves Creek Formation (−2‰ δ13C) show typical marine carbonate isotopic values (Figure 7).

This salinity controlled isotopic separation supports the observations of Stiller et al. (1985) who noted extreme enrichment of 13C in the dissolved inorganic carbon pool in evaporating brines up with δ13C values of up to + 16.5‰ under natural abiotic, oxic conditions in Dead Sea evaporation ponds (Figure 7). The systematic increase in 13C values in highly evaporated waters from the various bittern ponds of the Dead Sea Saltworks is thought to result from a nonequilibrium gas-transfer isotope fractionation. The process of ongoing evaporation leads to CO2 loss within the evaporative brine as less and less gas can held in solution (see Warren 2016, Chapter 9). CO2 exchange in a concentrating surface brine occurs directly between the water column and air, resulting in direct CO2 loss through evaporation. In a sabkha environment. CO2 is released from the hypersaline groundwater through sediments before being released to air as evaporites may form intrasediment precipitates. Overall, atmospheric CO2 uptake in hypersaline settings fed by shallow marine water is diminished compared to the normal marine settings.


Precipitated carbonates modern salinas and sabkhas are mainly aragonite, and formed in association with such evaporative brine, are consistently13C enriched, as seen in nearby Solar Lake and Sabkha Gavish (Figure 2; Stiller et al., 1985; Schidlowski et al., 1984). In a similar fashion, Palaeoproterozoic interbedded shallow marine carbonates, redbeds and evaporites have values up to δ13C + 17.2‰ (Melezhik et al., 1999). Permian and Triassic (Schmid et al., 2006a) redbeds and evaporite sequences also have 13C-rich carbonates (up to +7‰) and enrichment is partly attributed to evaporation and associated CO2 loss (Beauchamp et al., 1987). In modern oceans, atmospheric CO2 is consumed by biological activity and carbonate production originates from mainly marine organisms, leading to near atmospheric to organic negative δ13C signatures in the precipitated sediment(Andersson, 2013).

If increasing salinity leads to unfavourable conditions for photosynthesising organisms to survive (Lazar and Erez, 1992), carbonate through to bittern precipitation becomes increasingly abiotic and evaporation driven, especially at the upper end of the evaporation series. The loss of Ca during evaporation of a brine, via aragonite and calcium sulphate precipitation, leads to an increase in Mg/Ca ratio and an increase in residual brine density. This can result in primary dolomite precipitation or widespread reflux dolomitisation (Schmid et al., 2006, Warren 2000, 2016.

In summary, the typical δ13C signature in normal marine carbonate sediment across much of geological time centres around 0 ‰ and ranges between a few parts per mille on either side of the zero line reflecting precipitation by calcifying and photosynthesising organisms (e.g. algae), while abiotic, evaporation induced carbonates tend to have δ13C values above +1‰. More positive δ13C values (+4 to +6‰) tend to typify dominantly abiotic carbonates (and local methanogenic carbonates with even more positive values) and support the notion of evaporation-driven 13C-enrichement in times of widespread evaporitic epeiric and basinwide carbonates. In the Precambrian, widespread marine stromatolitic units such as, algal Loves Creek Formation reflects δ13C values for biogenic carbonate precipitation under shallow marine, non-hypersaline conditions. The change from a shallow hypersaline lagoon towards evaporitic mudflats and salterns suggests an increase in aridity and continentality/hydrographic isolation, with associated more positive δ13C values.

Implications for some types of 13C anomaly

The Bitter Springs Group chemostratigraphy has been correlated globally and the negative excursion was named previously after this unit (Bitter Springs Stage anomaly). However, the mechanism of evaporation-driven fractionation of δ13C is different from the commonly proposed inorganic-organic carbon fractionation, and challenges the views on interpreting global chemostratigraphic anomalies or excursion and their cause. Evaporite basins covered vast regions worldwide prior to the Sturtian glaciation, e.g. the Australian evaporites would have covered a third of the continent. The light surface of evaporites and associated carbonates would have had a high albedo and effectively cause less surface heat absorption. This subsequently would have triggered temperature decrease on a continental and possibly global scale. The Schmid paper hypothesises that the deposition of evaporites worldwide would have contributed to global cooling starting ≈100 Ma prior to Snowball Earth and would have played an important role in the onset of global glaciation.

References

Andersson, A.J., 2013. The oceanic CaCO3 cycle. In: T. Holland (Editor), Treatise on Geochemistry, 2nd ed. Elsevier, pp. 519-542.

Beauchamp, B., Oldershaw, A.E. and Krouse, H.R., 1987. Upper Carboniferous to Upper Permian 13C-enriched primary carbonates in the Sverdrup Basin, Canadian Arctic: comparisons to coeval western North American ocean margins. Chem. Geol. , 65: 391-413.

Berner, R.A., 1987. Models for carbon and sulfur cycles and atmospheric oxygen; application to Paleozoic geologic history. American Journal of Science, 287: 177-196.

Berner, R.A., 2001. Modeling atmospheric O2 over Phanerozoic time. Geochimica et Cosmochimica Acta, 65: 685-694.

Condie, K.C., 2016. Earth as an Evolving Planetary System (3rd edition). Elsevier, 350 pp.

Des Marais, D.J., 1997. Isotopic evolution of the biogeochemical carbon cycle during the Proterozoic Eon. Organic Geochemistry, 27(5): 185-193.

Des Marais, D.J., Strauss, H., Summons, R.E. and Hayes, J.M., 1992. Carbon isotope evidence for the stepwise oxidation of the Proterozoic environment. Nature, 359: 605-609.

Feulner, G., Hallmann, C. and Kienert, H., 2015. Snowball cooling after algal rise. Nat. Geosci. , 8: 659-662.

Feulner, G. and Kienert, H., 2014. Climate simulations of Neoproterozoic snowball Earth events: similar critical carbon dioxide levels for the Sturtian and Marinoan glaciations. Earth Planet. Sci. Lett., 404: 200-205.

Frakes, L.A., Francis, J.E. and Syktus, J.L., 1992. Climate modes of the Phanerozoic. Cambridge University Press, New York, 274 pp.

Goddéris, Y., Donnadieu, Y., Nédélec, A., Dupré, B., Dessert, C., Grard, A., Ramstein, G. and François, L.M., 2003. The Sturtian ‘snowball’ glaciation: fire and ice. Earth Planet. Sci. Lett. , 211: 1-12.

Hoffman, P.F. and Schrag, D.P., 2002. The snowball Earth hypothesis: testing the limits of global change. Terra Nova, 14: 129-155.

Holser, W.T., Schidlowski, M., Mackenzie, F.T. and Maynard, J.B., 1988. Geochemical cycles of carbon and sulfur. In: C.B. Gregor, R.M. Garrels, F.T. Mackenzie and J.B. Maynard (Editors), Chemical cycles in the evolution of the earth. John Wiley, New York, pp. 105–173.

Horton, T.W., Defliese, W.F., Tripati, A.K. and Oze, C., 2016. Evaporation induced 18O and 13C enrichment in lake systems: A global perspective on hydrologic balance effects. Quaternary Science Reviews, 131: 365-379.

Lazar, B. and Erez, J., 1992. Carbon geochemistry of marine-derived brines: I. 13C depletions due to intense photosynthesis. Geochim. Cosmochim. Acta, 56: 335-345.

Leng, M.J. and Marshall, J.D., 2004. Paleoclimate interpretation of stable isotope data from lake sediment archives. Quaternary Science Reviews, 23(811-831).

Liu, Y., Peltier, W.R., Yang, J. and Vettoretti, G., 2013. The initiation of Neoproterozoic ‘‘snowball” climates in CCSM3: the influence of paleocontinental configuration. Climate Past, 9: 2555-2577.

Melezhik, V.A., Fallick, A.E., Medvedev, P.V. and Makarikhin, V.V., 1999. Extreme 13Ccarb enrichment in ca. 2.0 Ga magnesite-stromatolite-dolomite-‘red beds’ association in a global context: a case for the world-wide signal enhanced by a local environment. Earth Sci. Rev., 48: 71-120.

Pierrehumbert, R.T., Abott, D.S., Voigt, A. and Koll, D., 2011. Climate of the neoproterozoic. Annu. Rev. Earth Planet. Sci., 39: 417-460.

Schidlowski, M., Matzigkeit, U. and Krumbein, W.E., 1984. Superheavy organic carbon from hypersaline microbial mats; Assimilatory Pathway and Geochemical Implications. Naturwissenschaften, 71(6): 303-308.

Schmid, S., 2017. Neoproterozoic evaporites and their role in carbon isotope chemostratigraphy (Amadeus Basin, Australia). Precambrian Research, 290: 16-31.

Schmid, S., Worden, R.H. and Fisher, Q., 2006. Carbon isotope stratigraphy using carbonate cements in the Triassic Sherwood Sandstone Group: Corrib Field, west of Ireland. Chem. Geol., 225: 137-155.

Stiller, M., Rounick, J.S. and Shasha, S., 1985. Extreme carbon-isotope enrichments in evaporating brines. Nature, 316: 434.

Swart, P.K., 2015. The geochemistry of carbonate diagenesis: The past, present and future. Sedimentology, 62(5): 1233-1304.

Warren, J.K., 2000. Dolomite: Occurrence, evolution and economically important associations. Earth Science Reviews, 52(1-3): 1-81.

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

Worsley, T.R. and Nance, R.D., 1989. Carbon redox and climate control through Earth history: A speculative reconstruction. Paleogeography, Paleoclimatology, Paleoecology, 75: 259-282.

 

Stable isotopes in evaporite systems: Part I: Sulphur

John Warren - Monday, April 30, 2018

 

Introduction

The sulphur isotopic composition of sulphate dissolved in modern seawater (SW), and the relationship with the associated modern and ancient sulphate precipitates, has been studied for more than five decades. An understanding of the controlling factors is fundamental in any interpretation of the origin of modern and ancient sedimentary calcium sulphates.

So, we shall look at the significance of sulphur isotopes, first by reviwing what is known in terms of the isotopic evolution of marine sulphate salts across the evaporation series from gypsum to the bitterns, and then across a time perspective via the evolution of oceanic sulphate and sulphide signatures from the Archean to the present.


Sulphur isotopes across the bittern series

The accepted d34S value of modern seawater-derived calcium sulphate (gypsum) is + 20.0 ±0.2‰ (Sasaki, 1972; Zak et al., 1980 and references therein). This is a average value, based on numerous analyses across the range ( +19.3 to +21.4‰). Notably, Rees et al. (1978) obtained a mean of +20.99 ± 0.09‰, using the SF6 method, which has a better reproducibility than the conventional S02 method. Mediterranean seawater gave a d34S value of +20.5‰ (Nielsen, 1978).

Measured values in natural gypsum from seawater show initial precipitates have a d34S value slightly higher than that of its source brine (Figure 1). The highest isotope differential for gypsum naturally precipitated from seawater, as recorded in the literature, is +4.2‰ (Laguna Madre, Texas, U.S.A.; Thode, 1964). Most reported d34Sgypsum-sw differentials lie in range from 0 to + 2.4‰ (Ault and Kulp, 1959; Thode et al., 1961; Thode and Monster, 1965; Holser and Kaplan, 1966).

Prior to Raab and Spirto (Figure 1; 1991), laboratory experiment data on d34Sgypsum-solution are scarce, especially for solutions mimicking initial precipitation of gypsum from natural seawater and passing into halite saturation. Harrison (1956) measured a d34Sgypsum-solution value of ~ + 2‰ for gypsum precipitated from an artificial solution, that was saturated with respect to gypsum. Thode and Monster (1965) calculated a K-value [ (32S/34S)solution/ (32S/34S)gypsum] of 1.00165 from a measured a d34Sgypsum-solution value of + 1.65‰ for a CaSO4.2H2O -saturated solution, evaporated under reduced pressure and allowed to age and equilibrate for 24 months at room temperature. An experiment using natural seawater was carried out by Holser and Kaplan (1966), who sampled the products of evaporating seawater in a tank with continuous refilling (green circles in Figure 1). The results show “only a small difference between brine and gypsum precipitated” (Holser and Kaplan, 1966, p.97), resulting in a mean value of d34Sgypsum-seawater = +1.7‰ (+19.4 to +21.1‰). Harrison (1956) calculated from experimental vibrational frequencies for S04 in solution and in crystalline CaSO4.2H2O, a constant K = 1.001 for the reaction:

(Ca34S04.2H2O + 32SO4)SOLID = (Ca32S04.2H2O + 34SO4)SOLUTION

which means a 1‰ increase of d34S in the solid fraction. Nielsen (1978, p. 16-B-20), using Rayleigh-type fractionation curves indicates that, “...the gypsum/anhydrite of the sulphate facies should be slightly enriched in 34S with respect to the unaffected seawater sulphate”

In the geological record the evaporites of the later Mg- and K-Mg- sulphate bittern facies are depleted in 34S relative to the earlier, basal Ca-sulphates, as rseen in the geological record. Nielsen and Ricke (1964, p.582) give a mean value of +2‰ for the depletion in 34S in later bittern evaporite sulphates relative to the basal Ca-sulphates in the Upper Permian Zechstein Series (Hattorf and Reyershausen, Germany) whereas Holser and Kaplan (1966, pp. 116 and 117) give a value of -1.0±0.8‰ (their d34Spotash-magnesia facies sulphates - d34Sgypsum/anhydrite facies) for the Zechstein Basin (Germany) and -0.8±0.5‰ for the Upper Permian Delaware Basin (U.S.A.) evaporites (green circles in Figure 1).

Theoretical calculations of the behaviour of the sulphur isotopic fractionation during the late evaporation stages were made by Holser and Kaplan (1966, pp. 116 and 117, fig. 4) and by Nielsen (1978, p. 16-B-20, fig. 16-B-12) applying the Rayleigh distillation equation and using the same fractionation factor calculated from the initial gypsum (1.00165). Their curves are thus in a continuous line with those calculated for the Ca-sulphates. These show an increasing degree of depletion in 34S in the sulphates precipitated in the course of the progressive evaporation in a closed basin, relative to the first Ca-sulphate precipitated, up to the end of the carnallite facies. They explain it by the continuous depletion in 34S in the brines. Thus their calculated d34Scrystal-initial gypsum at the end of the halite facies is ~ -0.6‰, at the end of the Mg-sulphate facies -1.0‰, and at the beginning of the carnallite facies -3.8‰, and relative to the original seawater (their 34SC) the differences are +1.0, +0.4 and -2.2‰, respectively. Nielsen (1978) also plotted an extrapolated fractionation curve for the residual brines in a closed reservoir, indicating that the brine is constantly depleted by 1.65‰ relative to the associated precipitate.

Prior to the laboratory work of Raab and Spiro (1991), no experimental data pertaining to the isotopic behaviour of sulphate sulphur in the late evaporative stages of seawater was available in the literature. Raab and Spiro evaporated seawater, stepwise and isothermally at 23.5°C, for 73 days, up to a degree of evaporation of 138x by H2O weight. At various stages of evaporation the precipitate was totally removed from the brine and the brine was allowed to evaporate further. The sulfur isotopic compositions of the precipitates and related brines showed the following characteristics (Figure 1) where the initial d34S of the original seawater is +20‰. The d34S of both precipitates and associated brines decrease gradually across the gypsum field nd aup to the end of the halite field, where d34Sprecipitate = + 19.09‰ and d34Sbrine = + 18.40‰. The precipitates are always enriched in 34S relative to the associated brines in these fields, but the enrichment becomes smaller towards the end of the halite field. A crossover, where the d34S value of the brines becomes higher than those of the precipitates, occurs at the beginning of the Mg-sulfate field. The d34Sprecipitate increases from + 19.09‰ at the end of the halite field through +19.35‰ in the Mg-sulfate field to + 19.85‰ in the K-Mg-sulfate field, whereas the d34Sbrine increased from +18.40‰, through +20.91‰ to +20.94‰, respectively.

This evolution implies different values of fractionation factors (a) for the minerals precipitated in the late halite, Mg-sulphate and K-Mg-sulphate fields, other than that for gypsum (1.00165). The value of aprecipitate-residual brine would then be very slightly >1 in the late halite field and >1 in the two later fields.

The experimental pattern of evolution of the d34S-values of the precipitates from their experiment is in good agreement with data for natural anhydrites interbedded in halites, where d34S-values are lower relative to basal gypsum (and secondary anhydrite), and of primary minerals of the Mg- and K-Mg-sulfate facies, reported in evaporitic sequences, such as those of the Delaware (U.S.A.) and of the Zechstein (Germany) basins and so can be used to better interpret a marine origin of the sulphate bitterns.


Ancient oceanic sulphate

The element sulphur is an important constituent of the Earth’s exogenic cycle. During the sulphur cycle, 34S is fractionated from 32S, with the largest fractionation occurring during bacterial reduction of marine sulphate to sulphide. Isotopic fractionation is expressed as d34S, in a manner similar to that used for carbon isotopes and the longterm carbon curves related to the sulphur isotope curve across deep time (see next article). Sedimentary sulphates (mostly measured on anhydrite, but also baryte) typically are used to record the isotopic composition of sulphur in seawater (Figure 2). Mantle d34S is near 0‰, and bacterial reduction of sulphate to sulphides (mostly as pyrite) strongly prefers 32S, thus reducing d34S in organic sulphides to negative values (≈ -18‰), so leaving oxidized sulphur species with approximately equivalent positive values (+17‰; Figure 3).


Historically, the sulphur cycle has been interpreted as being largely controlled by the biosphere and in particular by sulphate-reducing bacteria that inhabit shallow marine waters (Strauss, 1997). Typically, sulphur occurs in its oxidized form as dissolved sulphate in seawater or as evaporitic sulphate and in its reduced form as sedimentary pyrite. The isotopic compositions of both redox states are sensitive indicators for changes of the geological, marine geochemical or biological environments in the past (Figure 2). The isotope record of marine sedimentary sulphate through time has been used successfully to determine global variations of the composition of seawater sulphate.

The isotopic composition of sedimentary (biogenic) pyrite reflects geochemical conditions during its formation via bacterial sulphate reduction. Sedimentary pyrite is, thus, an important record of evolutionary (microbial) processes of life on Earth. Both time records (anhydrite and pyrite) have been combined in an isotope mass balance calculation, and changes in burial rates of oxidized vs. reduced sulphur can be determined (Strauss, 1997). This, in turn, yields important information for the overall exogenic cycle (i.e. the earth's oxygen budget as discussed in the next article).

And so, values preserved in ancient marine sulphate evaporites are part of the broader world sulphur cycle across deep time that includes movements in and out of marine sulphides (dominantly pyrite) and marine baryte precipitates (Figure 2). Values based on evaporitic CaSO4 are consistent with the ranges seen in modern gypsum (Figure 3). A plot of ancient marine CaSO4 evaporites shows the oxxidised sulphur curve for seawater has varied across time from +30‰ in the Cambrian, to around +10‰ in the Permian and that it increased irregularly in the Mesozoic to its present value of +20‰ (Figure 4). Oxygen values show much less variability and will be discussed in more detail in the next article in this series. Time-consistent variations are reflected in all major marine sulphate evaporite deposits and were most likely controlled by major input or removal of sulphides from the oceanic reservoirs during changes driven by longterm variations in tectonic activity and weathering rates.

Historically, simple removal of oceanic sulphate via an increase in the volume of megasulphate deposition in a saline giant was not thought to be accompanied by dramatic isotopic effects. Rather, variations within the global sulphur cycle were thought to be controlled by a redox balance with stored sulphides and organics in more reducing environments, which are also linked to the carbon cycle and the atmospheric oxygen budget.

In this scenario the oxidative part of the global sulphur cycle is largely governed by continental weathering (especially of marine black shale), riverine transport and evaporite deposition, while the reduced part of the sulphur cycle is controlled by levels of fixation of reduced sulphur-bearing compounds in the sediment column, mostly as pyrite via bacterial sulphate reduction (Figure 2.). The latter process preferentially removes isotopically light sulphur from seawater and so increases the d34S value in the ocean, and any consequent precipitate.

However, more recent work question aspects of this older sulphur cycle/pyrite/organics model. As just discussed, variations in d34Ssulphate across the Phanerozoic are traditionally interpreted to reflect changes in the total amount of sulphur buried as pyrite in ocean sediments — a parameter referred to as fpyr and defined as (Hurtgen, 2012);

fpyr = [(pyrite Sburial)/(pyrite Sburial + evaporite S burial)].

However, Wortmann and Paytan (2012) conclude that the 5‰ negative d34Ssulphate shift in ~120-million- year-old rocks was caused by massive seawater sulphate removal, which accompanied large-scale evaporite deposition during the opening of the South Atlantic Ocean (Figure 4). In their model, the negative d34Ssulphate shift is driven by lower pyrite burial rates that result from substantially reduced marine sulphate levels in the world ocean, tied to megasulphate precipitation. The authors attribute a 5‰ positive d34Ssulphate shift in the world’s oceans about 50 million years ago to an abrupt increase in marine sulphate concentrations as a result of large-scale dissolution of freshly exposed evaporites; they argue that the higher sulphate concentrations in the ocean in turn led to more pyrite burial.


Likewise, Halevy et al. (2012 ) studied past sulphur fluxes to and from the ocean, but over a longer time-frame (the Phanerozoic). They quantified sulphate evaporite burial rates through time, then scaled these rates to obtain a global estimate of variation in sulphur flux. Their results indicate that sulphate burial rates were higher than previously estimated, but also greatly variable. When Halevy et al. (2012) integrated these improved evaporite burial fluxes with seawater sulphate concentration estimates and sulphur isotope constraints, their calculations implied that Phanerozoic fpyr values (fpyr = fraction of sulphur removed from the oceans as pyrite) were ~100% higher on average than previously recognized. These surprisingly high and constant pyrite burial outputs must have been balanced by equally high and constant inputs of sulphate to the ocean via sulphide oxidation (weathering). These relatively high and constant rates of pyrite weathering and burial over the Phanerozoic, as identified by Halevy et al. (2012, suggest that the consumption and production of oxygen via these processes played a larger role in regulating Phanerozoic atmospheric oxygen levels than previously recognized, perhaps by as much as 50%.

Both studies recognize the importance of episodic evaporite burial on the sulphur cycle, while Wortmann and Paytan (2012) clearly show that large-scale deposition and dissolution of sulphate evaporites over relatively short geologic time scales can have an enormous impact on marine sulphate concentrations, pyrite burial rates, and the carbon cycle and so probably play a more important role than previously recognised in regulating the chemistry of the ocean atmosphere system.

The 18O content in seawater sulphate fluctuates less than sulphur values over geologic time (see next article for detailed discussion). The isotopic composition of sulphate minerals varied only slightly from the Neoproterozoic to the Palaeozoic decreasing from +17 to +14‰ (Figure 4). Values then rose during the Devonian to reach +17‰ during the Early Carboniferous (Mississippian). Values then fell to =+10‰ during the Permian, mimicked by a similar decline in sulphur values in the Late Permian to Early Triassic. Since the rise to +15‰ in the Early Triassic, values of marine sulphate minerals have remained close to +14‰ (add 3.5‰ to mineral determined value to give ambient seawater value). Overall, oxygen values show little correlation with marine sulphate variation and are perhaps are more controlled by sulphide weathering reactions.

What is also significant is that, given the now well established sulphur isotope age curve, a comparison of a measured d34S value from an anhydrite or gypsum of known geological age to the curve allows an interpretation of a possible marine origin to the salt. A value which differs from the marine signature does not necessarily mean a nonmarine origin, but, at the least, it does mean diagenetic reworking or, more likely, a groundwater-induced recycling of sulphate ions into a nonmarine saline lake (Pierre, 1988). Such oxygen and sulphur isotopic crossplots have been used to establish the continental (nonmarine) origin of the Eocene gypsum of the Paris Basin and the upper Miocene gypsum of the Granada basin, with sulphate derived from weathering of uplifted Mesozoic marine evaporites (Fontes and Letolle, 1976; Rouchy and Pierre, 1979; Pierre, 1982).

Sulphur is largely resistant to isotopic fractionation during burial alteration and transformation of gypsum to anhydrite (Figure 5; Worden et al., 1997). For example, primary marine stratigraphic sulphur isotope variation is preserved in anhydrites of the Permian Khuff Formation, despite subsequent dehydration to anhydrite during burial (≈1,000m) and initial precipitation as gypsum from Permian and Triassic seawater. Gypsum dehydration to anhydrite did not involve significant isotopic fractionation or diagenetic redistribution of material in the subsurface. At depths greater than 4300 m, the same sulphur isotope variation across the Permian-Triassic boundary is still present in elemental sulphur and H2S, both products of the reaction of anhydrite with hydrocarbons via thermochemical sulphate reduction (Figure 5). Clearly, thermochemical sulphate reduction did not lead to sulphur isotope fractionation. Worden et al. also argues that significant mass transfer has not occurred in the system, at least in the vicinity of the Permian-Triassic boundary, even though elemental sulphur and H2S are both fluid phases at depths greater than 4300 m. Primary differences in sulphur isotopes have been preserved in the rocks and fluids, despite two major diagenetic overprints that converted the sulphur in the original gypsum into elemental sulphur and H2S by 4300 m burial and the potentially mobile nature of some of the reaction products. That is, all reactions occurred must have occurred in situ; there was no significant sulphur isotope fractionation, and only negligible sulphur was added, subtracted, or moved internally within the system.


The resistance to fractionation of sulphur isotopes in subsurface pore waters can also be utilised to determine the origin of saline thermal pore waters. In a study of sulphur isotopic compositions of waters in saline thermal springs, Risacher et al. (2011) came to the interesting conclusion that dissolution of continental sedimentary gypsum from the Tertiary-age Salt Cordillera was the dominant supplier of sulphate (Figure 6). The sulphate in the springs was not supplied by the reworking of volcanic sulphur in this active volcanic terrain. d34S values from 3 to 11‰ in continental gypsum and this also encompasses the range of d34S in pedogenic gypsum (5 to 8‰) and in most surface waters (3.4 to 7.4‰) including salt lakes (Rech et al., 2003). Frutos and Cisternas (2003) found isotope ratios ranging from 1.5 to 10.8‰ in five native sulphur samples. Figure 6 presents the sulphur isotope ratio of dissolved sulphate in thermal waters sampled by Risacher et al. (2011) and references therein. The d34S of sulphate in northern thermal springs is within the range of salt lakes waters and continental gypsum. In an earlier paper Risacher et al. (2003) showed that salar brines leak through bottom sediments and are recycled in the hydrologic system. Deep circulating thermal waters are dissolving continental gypsum in sedimentary layers below the volcanics associated with the present day salars. The exception to this observation is the sulphur in Tatio springs where Cortecci et al. (2005) proposed a deep-seated source for the sulphate, related to magma degassing (Figure 6).


References

Cortecci, G., Boschetti, T., Mussi, M., Herrera Lameli, C., Mucchino, C. and Barbieri, M., 2005. New chemical and original isotopic data on waters from El Tatio geothermal field, northern Chile. Geochemical Journal 39: 547-571.

Fontes, J.C. and Letolle, R., 1976. 18O and 34S in the upper Bartonian gypsum deposits of the Paris Basin. Chemical Geology, 18(4): 285-295.

Frutos, J. and Cisternas, M., 2003. Isotopic Differentiation in Volcanic-Epithermal Surface Sulfur Deposits of Northern Chile: d34S < 0‰ in “Fertile” Systems (Au-Ag-Cu Ore Deposits below), versus d34S ≥ 0‰ for “Barren” Systems. Short Papers - IV South American Symposium on Isotope Geology (Salvador, Brazil, 2003): 733-735.

Halevy, I., Peters, S.E. and Fischer, W.W., 2012. Sulfate Burial Constraints on the Phanerozoic Sulfur Cycle. Science, 337(6092): 331-334.

Holser, W.T. and Kaplan, I.R., 1966. Isotope geochemistry of sedimentary sulfates. Chemical Geology: 93-135.

Hurtgen, M.T., 2012. The Marine Sulfur Cycle, Revisited. Science, 337(6092): 305-306.

Nielsen, H., 1978. Sulfur isotopes in nature. In: K.H. Wedepohl (Editor), Handbook of Geochemistry Section 16B, pp. B1 - B40.

Nielsen, H. and Ricke, W., 1964. Schwefel-lsotopenverhältnissen von Evaporiten aus Deutschland; Ein Beitrag zur Kenntnis von d34S im Meerwasscr-Sulfat. Geochimica et Cosmochimica Act, 28: 577-591.

Pierre, C., 1982. Teneurs en isotopes stables (18O, 2H, 13C, 34S) et conditions de genese des evaporites marines; application a quelques milieux actuels et au Messinien de la Mediterranee. Doctoral Thesis, Orsay, Paris-Sud.

Raab, M. and Spiro, B., 1991. Sulfur isotopic variations during seawater evaporation with fractional crystallization. Chemical Geology: Isotope Geoscience section, 86(4): 323-333.

Rech, J.A., Quade, J. and Hart, W.S., 2003. Isotopic evidence for the source of Ca and S in soil gypsum, anhydrite and calcite in the Atacama Desert, Chile. Geochimica et Cosmochimica Acta 67(4): 575-586.

Rees, C.E., Jenkins, W.J. and Monster, J., 1978. The sulfur isotopic composition of ocean water sulphate. Geochimica et Cosmochimica Acta, 43: 377-381.

Risacher, F., Fritz, B. and Hauser, A., 2011. Origin of components in Chilean thermal waters. Journal of South American Earth Sciences, 31(1): 153-170.

Rouchy, J.M. and Pierre, C., 1979. Donnees sedimentologiques et isotopiques sur les gypses des series evaporitiques messiniennes d'Espagne meridionale et de Chypre. Rev. Geogr. Phys. Geol. Dyn., 21(4): 267-280.

Sasaki, A., 1971. Variation in sulfur isotope composition of oceanic sulfate. 14th Int. Geol. Congr. Sect. 1: 342-345.

Strauss, H., 1997. The isotopic composition of sedimentary sulfur through time. Palaeogeography Palaeoclimatology Palaeoecology, 132: 97-118.

Thode, H.D., 1964. Stable isotopes a key to our understanding of natural processes. Bulletin Canadian Petroleum Geologists, 12: 246-261.

Thode, H.G. and Monster, J., 1965. Sulfur-Isotope Geochemistry of Petroleum, Evaporites, and Ancient Seas, Fluids in Subsurface Environments. AAPG Memoir 4, pp. 367-377.

Worden, R.H., Smalley, P.C. and Fallick, A.E., 1997. Sulfur cycle in buried evaporites. Geology, 25(7): 643-646.

Wortmann, U.G. and Paytan, A., 2012. Rapid Variability of Seawater Chemistry Over the Past 130 Million Years. Science, 337(6092): 334-336.

Zak, I., Sakai, H. and Kaplan, R., 1980. Factors controlling the 18O/16O and 34S/32S isotopic ratios of ocean sulfates and interstitial sulfates from modern deep sea sediments. In: E.D. Goldberg, Y. Horibe and K. Saruhaki (Editors), Isotope Marine Chemistry. Geochem. Res. Assoc, Tokyo, pp. 339-373.


 

Well (wireline) log interpretation of evaporites: An overview

John Warren - Saturday, March 31, 2018

Introduction

Often, when I run an evaporite training program for a client in the hydrocarbon or the potash industries, I am asked to add a short training module on the identification of evaporites in a set of conventional wireline log outputs. This blog is an overview of what I discuss in such a module. But every evaporite basin has its own set of mineralogies and problems and a generalised discussion, as in this blog, must be refined to meet the needs of the drilling or mining program in a particular evaporite basin.

Significant thickness units of evaporites are rarely cored in oil and gas drilling, unless in error, while when drilling rock chips of the more soluble salt minerals are quickly dissolved in most drill muds; so only a small portion of any subsurface evaporite bed can be studied directly. The situation is somewhat different in Salt and potash mining where cores are commonly collected ahead of the mine face to ascertain ore extent and thickness. Increasingly core calibrated well logs are replacing the need for extensive coring when ascertaining and predicting ore quality.

Many evaporite properties can be ascertained by examining a suite of conventional wireline logs. Many evaporite beds contain only one or two dominant saline minerals, they lack free pore fluids and have negligible porosity. This dramatically simplifies log interpretation and enhances the reliability of inferences with respect mineralogy. Thick clean evaporites will show the same characteristic set of log responses, not only locally but according to some authors worldwide (e.g., Serra, 1984, p. 173; Warren, 2006, Chapter 10). The most commonly available logs for the study of evaporites are the logs measuring hole diameter, electrical properties, bulk density, neutron porosity logs, sonic logs and, if significant levels of potash salts are present, both gamma and multispectral gamma logs.


Evaporites as seen in well logs

Well-logs are a continuous recording of a geophysical paragf meter along a borehole, where the value of the measurement is continuously plotted against depth in the borehole. Currently, the well logging industry is transitioning from wireline or cable-based well logging tools (Figure 1) to the increasing use of well-log tools designed for use in directional drilling. Wireline or cable tools can only be utilised in vertical to steeply inclined wells. The same set of conventional well log measurements are now increasingly collected using MWD (measurement while drilling) and LWD (logging while drilling) tools. With MWD/LWD, measurements are made by a suite of well-logging tools that reside immediately behind the advancing drill-bit. Part of the data collected by these tools is sent to the surface in real time (MWD) by mud pulsing or some other method of telemetry. The remaining portion of the collected data (LWD) is stored on a hard disk and recovered typically when a worn drill-bit is pulled to the surface to be replaced.

Although there are numerous well-logging tools and measurements that can be used in the study of evaporites, this section deals with only a few of the more conventional logging methods. Rider (1996) is an excellent overview from a geological, not petrophysical, perspective of the general principles of well-log interpretation. For a more comprehensive discussion of the geological applications of well-logs, there are many logging-company manuals, as well as excellent books and articles such as Kruger (2014), Crain (2010), Ellis and Singer (2007), Nelson (2007), Rider (1996), Nurmi (1978), and Alger and Crain (1966), Crain and Anderson (1966).


Electrical properties

Electrical resistivity, the reciprocal of electrical conductivity, is the degree with which a formation opposes the flow of electrical current. Onshore, a log of the spontaneous potential of a formation is run at the same time as a resistivity log. In reality, the measured resistivity is dependent on the combined resistivity of both the rock matrix and any contained fluids. Most solid rock materials are insulators, while their enclosed fluids are conductors. Hydrocarbons are the exception to fluid conductivity; they are infinitely resistive, and this is the basis for the quick look-identification of hydrocarbons and the use of Archies Law to determine water saturation levels in potential hydrocarbon reservoirs. In terms of evaporite identification, most evaporite units contain little if any pores or free water and so have very high resistivities compared to other more porous units (Table 1).

When the evaporite unit is relatively pure and monomineralogic, it creates a distinctive blocky log shape, whereas when it entrains beds of thin more porous lithologies (mudstones, shale, sands, limestone, dolomite) or perhaps contains brine-filled cavities and vugs, a muh spikier log is seen across a saline interval. The actual wireline log signature depends on the content of brine, sand, clay, bitumen and other variables. Within a local area in a basin, an elevated resistivity signature, although it does not allow a first indication of the presence of evaporite salts, can subsequently confirm it. When bitumens and salts are present in the same interval (as in salt-encased EoCambrian carbonate-slither reservoirs in the South Oman Salt Basin), the co-occurrence of halite cement, anhydrite cement and bitumen complicates a reliable interpretation of movable hydrocarbons.

Total & Spectral gamma-ray logs

The gamma-ray or gamma log is a record of the formation’s radioactivity. The radiation emanates from uranium, thorium and potassium which occur naturally in the formation. A simple gamma-ray log measures the radioactivity of the three radiogenic elements (U, K, Th) combined, while the spectral gamma log shows the amount of each radiogenic element contributing to a formation's radioactivity.

As a first indicator of lithology in non-evaporitic intervals, the gamma log is extremely useful in suggesting where shale may be expected in a formation; worldwide, elevated gamma readings in a sandstone-mudstone succession are typically used to indicate shaliness of the formation (Vclay). Clays can contain high levels of potassium-containing minerals, thorium (another radiogenic mineral) tends to be fixed in shales (c.f. sands), and that clays typically “fix” marine uranium into the sediment in three main ways i) chemical precipitation in acid (pH 2.5 - 4.0) or reducing environments, ii) adsorption by organic matter in the clays, iii) adsorption by phosphates in the clays. Uranium can also be mobilised and “re-fixed in the subsurface across redox interfaces. High gamma readings can also be due to the elevated potassium content in glauconite-rich sands, or the secondary movement of uranium to form “hot” cement and fissure fills in a number of Middle East reservoirs.


More importantly for our purposes, high gamma signatures can be associated with those evaporites which contain high proportions of the potassium salts such as sylvite, carnallite, and polyhalite (Table 1; Figure 2). In the potash-entraining salts, there is between 10% and 50% potassium by weight. When it is considered that the average shale contains only 2.7% potassium, the very strong radioactivity indicative of the potassium salts in an evaporite suite is understandable and means potash beds can be distinguished from the somewhat-elevate uranium-derived kicks of marine shale and the low radiogenic content of any adjacent halite, anhydrite or carbonate beds.

In contrast to halite units containing potassium salts, the more common evaporites, such as halite and anhydrite, give very low readings on the gamma log scale (Figure 2). Once an initial tie-back to core-determined assay values is done, it is possible to reliably estimate the percentage of K2O from the gamma response. As a general "rule of thumb," Edwards et al. (1967) showed that for a 6.25-inch, liquid-filled hole there was a correlation of 12.6 API units per 1% K2O.


Bulk Density Log

The bulk density or density log is related to the electron density of a formation and is the near-numerical equivalent of the formation's specific gravity (gm/cc); that is it is considered to measure variations in the average total density of the formation. A tool-measured value includes the density of the solid rock matrix and the density of fluids enclosed in the pores. The bulk density log is a measure of the degree of scattering or attenuation of gamma rays by electrons in the formation (Compton scattering). The electron density of a formation (electrons/cc) is closely related to the common density (gm/cc) and is typically used as a direct indicator of common density. Unfortunately, some minerals, including halite and sylvite, have electron densities that are not directly proportional to their specific gravities. Such minerals require the use of apparent bulk density for interpretation. Fortunately, many of the evaporite minerals have sufficient differences in bulk density to be recognised especiallry when crossplotred against Pef or NPHI (neutron) values (Figure 3, 4 and 5).

Many evaporite units are relatively pure and often mono- or bi-mineralic. Because of this, their lithological composition can be suspected, if not positively identified from the density log. However, when impure, the densities will fluctuate. Fortunately, most relatively pure evaporites thicker than a metre tend to give intervals of constant density with only minor variation. When this occurs, densities near the expected value in a clean evaporite unit can be easily identified and correlated to mineralogy using the bulk density log.

Neutron Logs

The neutron porosity index or neutron log provides a continuous record of a formation’s reaction to fast neutron bombardment. It is primarily a measurement of the hydrogen concentration in the formation, whether from the water of hydration, as in the case of hydrated salts such as gypsum and carnallite, or from water or oil in the more commonly understood non-evaporite situation. Quantitatively, the neutron log is used to measure porosity (in limestone-equivalent porosity units), qualitatively, it is a good discriminator between oil and gas in intervals without hydrated salts or other minerals. Geologically, it can be used to identify gross lithology, and so define evaporites (negative porosity values), hydrated minerals, and volcanic rocks and zeolites. A crossplot of formation bulk density versus neutron-log measurement is an extremely valuable tool for identifying various subsurface evaporite lithologies (Figure 2).

For example, in thick evaporite successions, a neutron log can distinguish between various evaporite salts on the basis of water of crystallisation (Rider, 1996). Gypsum is the most common of the evaporites containing water of crystallisation. However, carnallite, polyhalite, and kainite also contain the water radical (Table 1). In a neutron-density (NPHI-RHOB) crossplot, all these hydrated salts have high neutron-log values and characteristic tightly-clustered apparent bulk densities, which separates them from other anhydrous evaporites such as salt or anhydrite, which contain no water and hence have NPHI values near zero (Figure 4).


Sonic or Acoustic Logs

The sonic or acoustic log shows a formation’s interval transit time, designated ∆t, measured in microseconds/ft or microseconds/m (∆t is the reciprocal of sonic velocity * 1000). It is a measure of a formation’s capacity to transmit sound waves. Geologically this capacity typically varies with lithology and rock texture, notably porosity. Once again, because most subsurface evaporites have extremely low porosities and are often relatively pure, the sonic log can be used to reliably identify evaporites, once an initial identification has been made by some other means (Table 1). The seeming precision of the figures given in Table 1 are illusory as the actual transit times in thick evaporites can be strongly influenced by compositional variation, temperature and confining pressure.

Rock salt is formed mainly composed of the mineral halite and is a lithology whose density is effectively constant with depth (Warren, 2016; Chapter 1). Since density is probably the most critical factor in determining acoustic velocity, the ∆tma of a thick halite unit tends to be relatively constant over a wide depth range. For pure halite, the interval transit time is 68 µsec/ft (14,625 ft/sec). However, many halite units contain varying levels of impurities, usually anhydrite, either as interbeds or disseminated throughout the sequence. Anhydrite has an interval transit time of 50 µsec/ft (20,000 ft/sec). The velocity variations in a binary system of halite and anhydrite are related linearly (either by weight or volume) to the densities of the various mixtures of the end member values allowing semiquantitative determinations of the purity of the units. Sonic logs are widely used in the oil industry for correlation and the construction of synthetic seismograms. When considering representative velocities in interpreting seismic lines, it must be remembered that the presence of bedded anhydrite and carbonate units within the total rock salt interval can have an appreciable effect on the average seismic velocity through a salt interval.

Basic identification conventional wireline log outputs

The gamma log ()aka as the lithology log) measures the natural or spontaneous radioactivity of a formation. In a sand-shale basin, the measured gamma values are used to infer clay contet. In an evaporite basin, the gamma log (especially the spectral gamma log) is a reliable indication of the presence or absence of potash salts.

In a classic quick-look analysis of any potential hydrocarbon reservoir, the sonic, density and neutron logs are used both individually and in combination to estimate the porosity of likely reservoir strata. These three logs are referred to as the “porosity logs”. Although they are typically used to indirectly infer porosity, they actually reflect variations in rock properties related to the passage of sound, induced gamma radiation and high energy neutron bombardment. The fact there is negligible porosity in most subsurface evaporites means the “porosity logs” in combination with each other, or with a spectral gamma log, can be used to identify evaporite mineralogies.


With conventional wireline log suites in bedded and halokinetic sequences worldwide I use following quick-look procedure to identify various major anhydrite, halite and potash units (refer to Table 1 and figure 5):

1) Tentatively define evaporite intervals as zones dominated by lowest gamma-ray values (some carbonates also show very low, but typically slightly higher gamma values). Remember potash beds encased in halite, or less often anhydrite, will have high gamma values.

2) Confirm pure anhydrite intervals (thicker than a metre -tool resolution dependent) using

a) Sonic - ∆t ≈ 50 microseconds

b) Bulk Density - Log value of 2.98 gm/cc. Anhydrite densities in log curve greater than 2.95 typically indicate anhydrite (but be aware of possible metal sulphides (pyrite, galena) and barite cement in some evaporite masses, especially in the caprock to halokinetic structures. NPHI porosities of anhydrite tend to hover at zero or on the negative side (in standard limestone porosity units)

3) Confirm pure halite intervals (thicker than a metre) using a combination of density and NPHI (neutron) logs. Halite-dominated zones show a consistent combination of bulk densities around 2.1 gm/cc, negative NPHI porosities and sonic (delta T) values around 64 - 70 µs/ft.

4) Use the caliper - a curve tracking the nominal bit size to confirm anhydrite versus halite (±potash salts). A caliper value much larger than bit-size indicates borehole washout, it is probably due to the intersection of salt, not the less-soluble anhydrite. Anhydrite beds tend to show an "in-gauge" caliper profiles and also tend to be slower-drilled units compared to halite (penetration data can be seen in a mudlog or well completion report). However, carbonate intrasalt beds can also show slow drill penetration rates and an "in-gauge" profile

5) Use the resistivity log - Salt like anhydrite has high resistance to current flow (Table 1), but, due to wash-out, salt units often shows lower apparent resistivity values, especially in the microresistivity and shallow reading curve outputs.

6) If there are high gamma (K-rich) intervals within thick halite beds or high-density values (>2.8gm/cc) adjacent to anhydrite consider these intervals to be possible zones with elevated levels of salts such as sylvite, carnallite or polyhalite.

7) Zones of very-low apparent bulk densities (<1.4 gm/cc) and low gamma values within a thick halite may indicate beds dominated by a non-potash evaporite minerals such as bischofite.


8) An overlaid combination of a density and a gamma log in the same track can also useful in separating what are often two co-associated high density subsurface sulphate salts, namely anhydrite and polyhalite (Figure 6). Polyhalite is a potash salt with elevated density (2.8 gm/cc) and elevated gamma values. Anhydrite has an even higher density but lacks potash and hence exhibits relatively low Gr values. This diffence in potash response in what are both characteristically high-density minerals allows for their differentiation.

9) A lack of porosity and the near linear response of the spectral gamma log to mineral proportions means GR outputs when tied to assay values can be reliably used to infer K2O ore grades (Figure 7).

In summary, most nonporous, thick relatively monomineralic evaporite units are readily identified using wireline logs, and often the proportions of minerals can be reliably determined using relevant crossplots. However, any mineralogical interpretation based on a well log outputs is just that, an interpretation, and whenever possible should be checked against rock evidence such as chips or core. When confirming a log suite interpretation of an evaporite interval, you should keep in mind that chips composed of the more soluble evaporite minerals are often completely dissolved in the drilling mud before they reach the shale shaker. In this case, the wireline logs can give a better indication of actual mineralogy than the mud chips.


References

Alger, R.P. and Crain, E.R., 1966. Defining evaporite deposits with electrical well logs. In: L.L. Raymer, W.R. Hoyle and M.P. Tixier (Editors), Second Syposium on Salt. North Ohio Geol. Soc. , pp. 116-130.

Biehl, B.C., Reuning, L., Strozyk, F. and Kukla, P.A., 2014. Origin and deformation of intra-salt sulphate layers: an example from the Dutch Zechstein (Late Permian). International Journal of Earth Sciences, 103(3): 697-712.

Crain, E.R., 2010. Potash redux. InSite CWLS Magazibe, 29(2): 17-26.

Crain, E.R. and Anderson, W.B., 1966. Quantitative log evaluation of the Prairie Evaporite Formation in Saskatchewan. J. Can. Pet. Technol., 5(3): 145-152.

Ellis, D.V. and Singer, J.M., 2007. Well logging for Earth Scientists (Second Edition). Elsevier.

Fuzesy, L.M., 1982. Petrology of potash ore in the Esterhazy Member of the Middle Devonian Prairie Evaporite in southeastern Saskatchewan. NDGS/SKGS-AAPG; Fourth International Williston Basin Symposium, October 5-7, 1982: 67-73.

Hill, D.G., 1993. Multiple log potash assay. Journal of Applied Geophysics, 30(4): 281-295.

Kruger, N., 2014. The Potash Members of the Prairie Formation in North Dakota. Report of Investigation - North Dakopta Geologica; Survey No. 113.

Nelson, P.H., 2007. Evaluation of potash grade with gamma-ray logs. U.S. Geological Survey Open-File Report 2007-1292, 14 p.

Nurmi, R.D., 1978. Use of well logs in evaporite sequences. In: W.E. Dean and B.C. Schreiber (Editors), Marine evaporites. . Soc. Econ. Paleontol. Mineral., Short Course notes, Tulsa, OK, pp. 144-176.

Rider, M., 1996. The geological interpretation of well-logs, second edition. Whittles Publishing, Caithness, 268 pp.

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


Salt Dissolution (5 of 5): Metals and saltflow-focused fluids

John Warren - Wednesday, February 28, 2018

 

Introduction

Most subsurface evaporites ultimately dissolve and, through their ongoing dissolution and alteration, can create conditions suitable for metal enrichment and entrapment in subsurface settings ranging from the burial diagenetic through to the metamorphic and igneous realms. This article looks at a few examples tied to halokinesis, a more comprehensive set of examples and more detailed discussion is given in Chapters 15 and 16 in Warren (2016). Because most, if not all, of any precursor salt mass that helped form these metalliferous deposits via dissolution, has gone, the resulting metal and other accumulations tend to be at or near the edges of salt basins, or in areas where most or all of the actual salts are long gone (typically via complete subsurface dissolution or metamorphic transformation, so that only breccias, weld and indicator mineral suites remain).

A lack of a direct co-occurrence with evaporite salts is perhaps why the metal-evaporite association is not recognised by some in the economic geology community. The significance of disappearing salt masses in focusing and enhancing metal precipitation, via the creation of chloride-rich and sulphate-rich brines, may not be evident without the conceptual tools needed to recognise the former presence of evaporites, post-salt halokinetic structural geometries, and meta-evaporite mineral associations.

The various ore tonnage-grade plots in Warren (2016), shows that many metal accumulations with an evaporite association tend to plot at the larger end of their respective deposit groupings.



Evaporite dissolution helps create "prepared ground."

I am not saying all large metal accumulations require evaporites or the highly-saline subsurface fluids that they can generate. Although, some recent papers do argue for a widespread role of evaporites in a Pb-Zn association (Fusswinkel et al., 2013; Wilkinson et al., 2009) and in sedimentary redbed copper deposits (Rose, 1976; Hitzman et al., 2010).

Typically, the conceptualisation of an evaporite in the economic geology literature is as a bedded evaporite and brine source (Figure 1). Likewise, this article and the relevant chapters in Warren (2016) detail a number of megagiant ore deposits where dissolving evaporite bodies have contributed in some way to a metal accumulation (Table 1). However this current article, like Warren 2016) focuses on the mechanisms and indicators tied to a halokinetic-ore association. Halokinesis is an aspect of evaporites that is not widely discussed in the field of ore deposit models.


Not all sediment-associated ore deposits are associated with evaporites. Only in those ore deposits classified as anorogenic and/or continental margin can subsurface evaporite masses can be involved in the same unusual concentration and alteration conditions that lead to the creation of metalliferous ore deposits (evaporite associations are indicated by E in Figure 2). At other times and locations hydrothermal mineral salts, especially anhydrite (CaSO4)which can supply sulphur as it dissolves, can be an integral part of the ore accumulation, but their occurrence may be unrelated to aridity. Hydrothermal anhydrite and other burial/magmatic hydrothermal salts tend to form in high salinity conditions inherent to the ore-forming environment and not necessarily to the presence of precursor evaporites; as in the formation of carbonatites (e.g. Afrikanda and Bayan Obo; Wu, 2008), or pegmatites and some IOCG deposits (hydrothermal anhydrite is indicated by HA in Figure 2). In some such hot subsurface settings the role of any nearby buried true” evaporite may be, via its dissolution or alteration, to aid in the creation of highly-saline high-temperature basinal brines (Chapter 16; Warren, 2016). According to whether the resulting brines are chloride or sulphate-rich, they can act as either enhanced metal carriers or fixers.

The role of evaporites creating metalliferous ores is two-fold; 1) In solution (halite-dominant precursor) they can act as chloride-rich metal carriers and 2) Locally, asCaSO4 beds or masses alter and disintegrate, their dissolution products, especially if trapped, can supply sulphur (mostly via bacteriogenic or thermogenic H2S). Dissolutional interfaces set up chemical interfaces that act as foci during brine mixing so manufacturing conditions suitable for precipitation of metal sulphides or native elements. As a consequence, most evaporite-associated ore systems tend to epigenetic, rather than syngenetic. Subsurface salt beds and masses are merely the solid part of a sizeable ionic recycling system, dissolved metals are another part, and zones of mixing between the two are typically sites where metal sulphides tend to gather.

At the world-scale, both evaporite and ore systems are driven by plate tectonics. Halite-dominated sequences, deposited in the drawdown basin centres, tend to dissolve in burial, and so supply chloride ions to the brine system. Salt beds that are thick enough tend to flow and thus focus the upward and centripetal passage of basinal and hydrothermal fluid flows. Dissolving gypsum or anhydrite beds, typically deposited higher on the basin platform or diagenetically accumulated along salt dissolution edges and salt welds (touchdowns) can supply sulphur, via bacterial or thermochemical sulphate reduction, while simultaneously focusing the subsalt metalliferous brine flows into the precipitation interface.

When the chemistries of the dissolving salt beds and the metal carriers interact so that redox fronts, salinity contrasts, and other precipitative interfaces are set up, an ore deposit can form. Thus, in base and precious metal exploration in evaporitic terranes, we are ultimately searching for those parts of a subsurface ionic cycling system where the salt dissolution, salt beds and metal systems have interacted to create economic levels of metalliferous precipitates.

Modelling

Conceptually, this evaporite-related notion of regional fluid flow in a sedimentary/metasedimentary host is somewhat different to the internal process and local mineralised halo models that dominate our understanding of those world-class ore deposits related to the interior workings of igneous systems. The latter is known as an orthomagmatic system where internal igneous processes of fractional crystallisation and liquid immiscibility largely control ore formation. Ores are deposited in an evolving framework of world-scale tectonics and magmatism across time, from Archaean greenstones to those of present-day sialic plate tectonism. Examples, where buried evaporites have been assimilated into a magma chamber, are discussed chapter 16 in Warren, 2016. Then there are the various ore deposits that are external to (paramagmatic) or unrelated to the emplacement of igneous bodies (nonmagmatic). In both cases, the mineralisation is typically part of an ongoing long-term sedimentary burial history, tied to dissolving and flowing salt masses and associated hydrothermal circulation.

Evidence for hydrothermally-induced low-moderate temperature mineralisation is often best preserved in textures in the hydrothermally altered rock matrix, typically located outside the actual ore deposit (in its hydrothermal alteration halo). From the hydrothermal fluid perspective, one should see the role of evaporites and metal sulphides as each contributing its part to a larger scale “mineral systems” paradigm; much in the same way as, in a petroleum system, the integration of concepts of source, carrier, seal and trap are fundamental requirements to understand and predict economic oil and gas accumulations.

This holistic ore systems approach is not fully encompassed in some economic geology studies that use sequence stratigraphic sedimentological approaches for ore deposit prediction in greenschist terrains (Ruffel et al. 1998; Wilkinson and Dunster, 1996). In my opinion, this approach can shift the interpretation paradigm too far into the depositional realm. The problem with classic sequence stratigraphic criteria, when trying to understand ore genesis, is that sequence stratigraphy does not handle well the concept of a mobile ephemeral subsurface salt body that climbs the stratigraphy via autochthonous and allochthonous process sets (halokinesis). As the salt flows, it dissolves and so brings with it the associated epigenetic influences of brine-driven diagenesis and metasomatism.

Current sequence stratigraphic paradigms in the economic geology realm are dominated by the assumption that the geometry of the units in the depositional system, and associated fault characteristics, are relatively static within the buried sediment prism. Yet, in terms of most sediment-hosted hypersaline ore deposits, what is most important in understanding the metal-evaporite association is the understanding of; 1) Evaporite dissolution and halokinesis, 2) Migration of subsurface fluids, 3) Creation of shallower or lateral-flow redox fronts along with, 4) Opening and closing of fault/shear focused fluid conduits, typically tied to, the coming and going of bedded and halokinetic salts. These factors, rather than primary sediment wedge geometries, are the dominant controls as the mineralising system passes from the diagenetic into the metamorphic realm.

It is interesting that in a benchmark paper, discussing and classifying the world’s ore deposits in a plate-tectonic-time framework, Groves et al., (2007) list almost all the major ore categories shown in Figure 2b as belonging to the group of “...sediment-hosted deposits of non-diagnostic or variable geodynamic setting.” Into this category, they place all stratiform to stratabound sediment-hosted deposits with variable proportions of Pb, Zn, Cu (including Zambian Copper belt, Kupferschiefer and SedEx deposits). They go on to note (p. 26) that, although there is general agreement that the majority of these various deposits formed during active crustal extension, either in intracratonic rift basins or passive margin sediment hosts, there is considerable controversy concerning their broader scale tectonic setting at the time of mineralisation and the driving force for hydrothermal fluid flow at the time of their mineralisation.

Perhaps this lack of model specificity in the varied interpretations of sediment-hosted deposits reflects the fact that one piece of significant information is missing from many ore genesis models. Namely, that the greater majority of these poorly classified sediment-hosted deposits sat atop, or adjacent to, or beneath, what were once thick evaporite sequences (Table 1). In many cases, the salt mass is long gone. It was the dissolution of these salt masses, either bedded or halokinetic-allochthonous, that focused much of the ore-fluid flow in the sedimentary-diagenetic realm. The loss of salt as the basin sediments passed from the low temperature diagenetic into the metamorphic realm, and as the metalliferous fluid flow was focused into permeable conduits about, below or above the dissolving and retreating, or flowing salt edges, is how salt-related ore deposits form.

This is why the majority of these salt-aided deposits tend to occur outside salt basins that retain substantial salt masses still in the diagenetic realm. The deposits are a response to the dissolution and flow of evaporites, or the residual seawater bitterns created in underlying and subjacent settings as the salt beds were deposited, not to the presence of actual undissolved primary evaporite masses. As we see in Proterozoic and Archaean meta-evaporites and most Precambrian evaporite associations, the original salt mass is long gone from the hosting succession, via varying combinations of halokinesis, dissolution and metasomatism (Warren, 2016, Chapter 13; Salt Matters blog, August 28, 2016).

Ore deposits of Precambrian tend to be linked to evaporite alteration products and residues and rarely preserve actual sedimentary salts (other than local remains of minor hydrothermal anhydrite). In younger Phanerozoic deposits, such as Kupferschiefer, the Atlantis II deep and Dzhezkazgan, portions of actual salt (brine source) can remain in the more deeply buried parts of the basin.

Metal sulphide precipitates are not rare or unique in the subsurface diagenetic fluid milieu, what is essential in the prediction of ore-grade levels of metal sulphide buildups is understanding where and why the metal precipitation system is focused into particular structurally-controlled positions and encompass time frame/fluid volumes sufficient to build an ore deposit.

That is, evaporite-associated ore deposits are no more than ancient subsurface hydrology-specific associations where the precipitation system was stable enough, for long enough, to allow higher, ore-grade levels of metals sulphides to accumulate from carrier brines at particularly favourable and stable chemical and temperature interfaces. As such, metal precipitation sites are part of an ore-forming process set, spread across the epigenetic and syngenetic realms (Table 1). They are part of the regional evolution of the fluid plumbing from the time of deposition, into burial, and on into the realm of metamorphic transformation. This means to understand the ore system tied an evaporite-entraining system holistically; one must integrate local ore paragenesis with various aspects of the basin-scale geology, sedimentology, sequence stratigraphy, diagenetic-metamorphic-igneous facies, fluid flow conduits and structural evolution of the evaporitic basin.

Metals with a halokinetic focus

To illustrate the importance of salt dissolution tied to halokinetic fluid focusing I have chosen two well-known deposits, one is a stratiform redbed copper association (Corocoro deposit), the other a SedEx style Pb-Zn deposit (McArthur River or HYC deposit)

Corocoro and other sandstone-hosted deposits of the Central Andes

Stratabound deposits of copper (±Ag), hosted by variably-dipping continental clastic sedimentary rocks, occur in Central Andean intermontane basins and are known to postdate compressive deformation/uplift events in the region (Flint, 1986, 1989). The deposits are relatively small with variable host-rock depositional ages and include; Negra Huanusha, central Peru (Permo-Triassic); Caleta Coloso, northern Chile (Lower Cretaceous); Corocoro, northwestern Bolivia (Oligo-Miocene); San Bartolo, northern Chile (Oligo-Miocene); and Yasyamayo, northwestern Argentina (Miocene-Pliocene).


The Corocoro area has produced the largest amount of copper in these Andean examples, something like 7.8 million tonnes of copper at a grade of 7.1% (Cox et al., 2007). The location of mineralisation is controlled by structurally-focused redox fronts in bedded sediment hosts, which abut a steeply-dipping translatent thrust fault (Figure 3). Deposits are irregular, usually elongate lenses of native metal, sulphides, and their oxidation products. Typically, deposits are hosted in alluvial fan and playa sandstones or conglomerate facies that also contain abundant gypsum and lesser halite. The undersides of some copper sheets at Corocoro even preserve mudcrack polygons and bed-parallel burrow traces (Savrda et al., 2006). Ore mimicry of mudcracks is not a feature controlled by on-for-one-replacement of organic material deposited in a sandstone; rather it is following pre-existing permeability/redox contrasts.

Corocoro deposits have been mined sporadically since they were first exploited by the local Indians, prior to the Spanish invasion in the 16th century and were largely exhausted in half a century of more intense mining operations that began in 1873 (Figure 3). Sandstone and conglomerate matrices show evidence of bleaching and leaching of the original redbed host with numerous red-greybed redox interfaces visible in the mined sequences. Ore minerals (dominantly native copper) are secondary fills within secondary intergranular pores created by the dissolution of earlier carbonate and sulphate masses and intergranular cement. Twelve grey sandstone beds, which were host to the long worked-out native copper ores, occur within a stratigraphic thickness of 60 m, in a unit known as the Ramos Member that still hosts abundant CaSO4 as gypsum (Figure 3).

Ores are stratabound, but not necessarily stratiform, and the larger masses of native copper are typically shallower and present as vein fills. Sometimes the copper pseudomorphs large orthogonal-ended aragonite prisms, which can be several centimetres across. There are two main styles of mineralization; 1) Ore minerals as a matrix to stratiform detrital silicates, typically low dipping and commonly highlighting primary sedimentary structures, such as cross stratification, 2) Ores in stacked channelized sand bodies, that show steep dips in structurally complex and folded zones with local brecciation (Figure 3). Native copper commonly fills thin laterally extensive sheets in tectonic fractures in the limbs of tight folds. Ljunggren and Meyer (1964) interpreted these folded diagenetic sheets of copper as a remobilization products precipitated during deformation of earlier matrix-pore filling copper.

Critical factors in Corocoro ore genesis include (Flint, 1989; Aliva-Salinas, 1990): 1) Stratigraphic association of evaporites, organic-rich lacustrine mudstones, clastic reservoir rocks, and orogenic, igneous provenance areas for both basin-fill sediments and metals; and 2) Intrabasinal evolution of metal-mobilising saline brines derived from the buried and dissolving lacustrine evaporites that flush volcaniclastics, volcanics and feldspathic sediments. The same saline diagenetic fluids also caused the dissolution of early, framework-supporting cement and large aragonite prisms, all now pseudomorphed by native copper. Avila-Salinas (1990) notes the presence of a salt-cored décollement and its likely tie to some of the highly saline sodium chloride brines found at depth in the vicinity of the Toledo Mine (Figure 3b)

The ore-hosting clastic horizons are consistently located in the highly gypsiferous Vetas Member of the Ramos Formation, which was deposited as redbeds in braidplains or fluviodeltaic playa margins centripetal to the edges of saline evaporitic lakes that were accumulating gypsum and halite (Figure 4; Flint, 1989). Abundant gypsum is still present in the Ramos Member as nodules and satinspar vein fills. Both are secondary evaporite textures likely implying the dissolution of previously more voluminous CaSO4 and NaCl beds and masses. Gypsum along with celestite are the most common gangue minerals associated with native copper veins in all the Corocoro deposits (Singewald and Berry, 1922). In the geological analysis of the first two decades of last century, the copper-bearing beds of the westerly-dipping series were called "vetas" and those of the easterly-dipping beds "ramos" and, as a matter of convenience, the names became attached to the rocks themselves. The term "veta" is Spanish for vein and "ramo" the Spanish for branch (native copper). The 1922 paper by Singewald and Berry noted that the veta horizons were traceable continuously for over 5 km in outcrop, but they found no apparent primary trends related to ramos outcrops (Figure 3).

Six mineralised layers of each kind were in exploited in mining during the first two decades of last century, the thicknesses of which varied from a few centimetres to 7 meters (Figure 3). Sheets and masses of native copper, called charque, were up to 600 pounds in weight, but more significant volumes of copper were extracted from vetas sandstones where copper was found as diffuse minute grains, pellets, or granular masses of the native metal. Associated with the enriched copper zones were more oxidised minerals as malachite, chrysocolla, azurite, domeykite, and chalcocite. Singewald and Berry (1922) noted gypsum and salt were the principal gangue minerals, while silver minerals were rare. The vetas sediment hosts tended to be coarser grained, often conglomeratic; whereas the ramos sediment hosts were finer-grained with copper present as smaller particles and masses.


The currently accepted interpretation of the Corocoro copper is that it formed during early diagenesis within a saline playa depositional environment, and in combination with dissolution of the adjacent bedded lacustrine evaporites (Figure 4). This bedded combination is thought to have controlled the formation, transport and precipitation of the copper ore (Flint, 1989). Playa sandstones, sealed between impervious evaporitic mudstone layers, created the plumbing for focused metalliferous fluid migration toward the basin margin. It is argued that the carbonaceous material at Corocoro was likely concentrated in the sandstones and conglomerates and not in the shalier members of the sedimentary sequence (Eugster, 1989).

The organics were considered strata-entrained as primary plant matter (e.g. spores) preferentially in the sandstones, along with later possible catagenic/hydrothermally cracked products migrating as hydrocarbons out of the basin. This created locally reducing pore environments in the aquifers wherever these reduced fluids met with somewhat more oxidising updip pore waters. This updip migration of saline reducing waters, in combination with sulphur supplied as H2S from the adjacent dissolving calcium sulphate beds and nodules, as well as from dissolving intergranular sulphate cement, precipitated copper in the newly created secondary porosity. The pore water chemistry and flow hydrology of this sandstone-hosted Cu system is thought to show many affinities with diagenetic uranium-redox precipitating systems, as defined by Shockey and Renfro (1974).

However, there is, in my mind, a possible anomaly in this model, which assumes organics were deposited in fluvial sandstones at the time of deposition. It is highly unusual to have higher plant material accumulating in large volumes in sandstone in a setting that is sufficiently arid and oxidising to precipitate ongoing interbeds of halite and gypsum. Such settings are typically too dry to allow abundant higher plant growth. Also, groundwaters that are flowing basinward through bajada sandstones in Neogene sediments of the Andes are ephemeral or too oxidising to facilitate the long-term reducing conditions needed to preserve significant volumes of high plant remains in the sandstone aquifers.

What is also interesting in this sedimentological/diagenetic model of Tertiary age cupriferous redbeds deposits in the Andes, centred on Corocoro, but not considered in any detail in the published literature base, is the question..., What controlled the folding, and the associated brecciation and perhaps even subsurface brine interfaces responsible for the Cu precipitation? All the stratabound Bolivian Cu deposits accumulated in sediment hosts that were deposited in fault-bound intermontane groundwater sumps. All are located in hydrologic lows in the crustal shortening tectonic scenario that typifies the Tertiary history of the Andes.

The variable ages of the host sediments and the predominance of evaporite indicators including gypsum in outcrop (often as diagenetic residues, not primary, features in the fluvial hosts) and all intimately tied to the Corocoro ore forces the question...., “was the fluid focusing driving the Cu precipitation a response to compression-driven halokinesis in an evolving salt-lubricated thrust belt?” Did this on-ground scenario occur in a halokinetic hydrology, that was possibly related to a combination of thrust-driven telogenesis, redox setup, evaporite dissolution and aquifer focusing of brines with dissolution aiding local slumping? This, along with associated strike-slip prisms, could better explain the stability of redox interfaces in sandstone aquifers across timeframes needed to accumulate significant native copper volumes. After all, most of the ore textures are passive precipitates, mainly in pre-existing porosity. If so, perhaps these deposits are not a variation on a roll-front uranium theme, which is predicated on dispersed primary organic material in the host sandstones (Shockey and Renfro, 1974).


When one plots the position of Corocoro and other redbed copper across the region, the 1000-lb gorilla that has been standing in the corner of the room for the past century becomes obvious. The Corocoro redbed copper deposit is located on a salt-cored fault system linked across less than a kilometre to an outcropping gypsum-capped remnant of a salt diapir which crosscuts the anticlinal axis of a saline redbed/greybed Corocoro sequence and ties to the saline decollement of the Corocoro Fault (Figure 5). The same tie to salt-cored decollement and diapir proximity is true of other nearby redbed copper deposits to the south-southeast, such as Veta Verde and Callapa. It is highly likely that the saline fluid interfaces forming the redbed Cu deposits of Corocoro, Veta Verde and Callapa were halokinetically focused. A similar-salt lubricated set of thrusts and strike-slip faults typifies halokinetic anticline outcrops in Central Iran.

It is highly likely that much of the structuration that is controlling Corocoro ore positioning is a response to salt flow related uplift, brine conduits and fracture creation. Metal precipitation occurred at redox interfaces induced and controlled by regional salt-lubricated compressional tectonism, and the associated salt-structuration has driven the brine-interface redox hydrology.

Work by Rutland (1966) did make an observation that the Corocoro ore deposits are related to an unconformity between the Ramos and Vetas Formations. Previously, the unconformity was interpreted as directly due to the outcrop of the Corocoro Fault. He noted that the fault and the unconformity were one and the same. In the 1960s there was no notion of a salt weld but it was nonetheless a highly astute observation by Roy Rutland. He went on to note a similar unconformity is tied to the growth of the Chuquichambi salt diapir, some 100 km southeast of Corocoro. Unfortunately, the halokinetic implications of Rutland's work were not considered 20 years later in Flint's key 1989, paper inferring a mostly clastic sedimentological origin for the Corocoro and other similar SSC deposits.

A possible halokinetic/weld association also leads to the question... Were the salt lakes, that are considered an integral part of the depositional and saline ore-precipitation systems at Corocoro by Flint, also a response to dissolution of the same nearby diapiric structures, when they were active in the mid to late Tertiary? This tie, between diapir/weld brines sourced in the drainage hinterland and bedded evaporite - lacustrine mud interbeds accumulating in the groundwater outflow sumps, is the case with groundwater inflow for the Salar de Atacama infill, as it is in other Quaternary salt lakes in the region. The are many diapir remnants across the Andes region. It seems that the Corocoro style of Cu mineralisation is perhaps another example of suprasalt redox focusing in a halokinetic setting.

Whether the halokinetic scenario, or the currently accepted non-halokinetic bedded arid-lacustrine evaporite scenario, explains the Cu mineralisation Corocoro is yet to be tested. But in terms of future copper exploration for similar deposits, it probably requires an answer. A halokinetic association offers an exploration targeting mechanism, utilising satellite imagery and aerial/gravimetric data, prior to the acquisition of on-ground land positions and geochemical surveys.

McArthur River (HYC), Ridge II and Cooley II deposits, Australia

This material on the HYC deposit will be expanded upon in an upcoming paper by Lees and Warren (in prep.). Before mining, the McArthur River (or HYC) Pb-Zn-Ag deposit, contained 227 million tonnes of 9.2% Zn, 4.1% Pb, 0.2% Cu and 41 ppm Ag (Logan et al., 1990; Pirajno, and Bagas, 2008). The deposit is hosted in the HYC Pyritic Shale member and lies adjacent to the Emu Fault in the McArthur Basin and adjacent to what are currently sub-economic base metal deposits in the Emu Fault zone known as the Cooley II and the Ridge II deposits (Figure 6a). Across all these deposits, major ore sulphides are pyrite, sphalerite and galena, with lesser chalcopyrite, arsenopyrite and marcasite. The mineralised region has an area of two km2 and averages 55 m in thickness (Figure 6b). It is elongated parallel to the major Emu growth Fault, which lies 1.5 km to the east, but is separated from the main ore mass by carbonate breccias of the Cooley Dolostone Member (Figure 6a-d).


The sequence at McArthur River comprises dolomites of the Emmerugga Dolostone (with the Mara Dolostone and Mitchell Yard members), overlain by the Teena Dolostone with abundant aragonite splays indicative of a normal-marine tropical Proterozoic carbonate. Overlying the Teena Dolostone in the vicinity of the HYC deposit is the somewhat deeper water Barney Creek Formation and its equivalents, containing the W-Fold Shale member, while the ore is hosted in carbonaceous shales, with multiple lenses of fine-grained galena-sphalerite-pyrite, separated by inter-ore sedimentary breccias (Large et al., 1998). This unit contains numerous sedimentary features indicative of a deeper-water anoxic setting. For example, comparison with d13C values from isolated kerogen in the HYC laminites confirms that n-alkanes in Bitumen II are indigenous to HYC, indicating that the deposit formed under euxinic conditions. This supports a generally-held model for Sedex deposits the region, whereby lead and zinc reacted in a stratified water column with sulphide produced by bacterial sulphate reduction (Holman et al., 2014).

The ore-hosting organic-rich 1,643-Ma HYC Pyritic Shale Member of the Barney Creek Formation is much thicker in the HYC sub-basin than elsewhere in the Batten Trough Fault Zone (e.g., Glyde River Basin) and consists mainly of dolomitic carbonaceous siltstones (Figure 7; Davidson and Dashlooty, 1993; Bull 1998). I would argue this thickening reflects a combination of long-term local basinfloor subsidence, related to salt withdrawal, and brine stratification due to ongoing salt dissolution and focused outflow. Indicators of former salt allochthon tiers are widespread in the vicinity of the HYC deposit, but are absent in the Glyde River Basin.


Breccias in and around HYC

In the HYC mine area, the ore interval is overlain by the HYC pyritic shale member and made up of pyritic bituminous and dolomitic shales and polymict breccias (Figure 7). Importantly, when contacts are walked out in outcrop, the polymict breccias are significantly transgressive to bedding, while drilled intersections in the vicinity of the HYC deposit and in the mine itself show the breccias are stratabound. Another interesting feature of these breccias is that they can contain mineralised clasts. More broadly, a variety of sedimentary breccias occur throughout the Barney Creek Formation stratigraphy, especially along the eastern margin of the HYC half graben and tend to pass updip into the breccias of the Cooley Dolostone (Figure 6a).

Williams 1976, defined three breccia types (I, II and III) in the HYC area. Type I breccia beds occur in the lower half of the HYC Pyritic Shale Member and contain clasts characteristic of lithologies in formations of the McArthur Group below the Barney Creek Formation (Table 2). In the northern end of the sub-basin, the breccias are of a chaotic nature with no sorting and minor grading of clasts (Figure 6b). The underlying shale beds are frequently contorted and squeezed between the breccia fragments, which reach a maximum size of approximately 10 m. Toward the south, the thickness and maximum clast size of individual breccia beds decrease (Figure 6b). All breccia units are thickest adjacent to the Emu Fault Zone and likely record sediment sinks controlled by rapid fault-controlled basin subsidence during Barney Creek time. Inter-ore breccias amalgamate and thicken to the north-north-east of HYC, and occupy a position toward the foot of what is interpreted as a more substantial breccia lens, dominated by sediment gravity flow deposits (Figure 6d; Logan et al., 2001).


In a subsequent study, Ireland et al. (2004a) identified four distinct sedimentary breccia styles within Type I breccias: framework-supported polymictic boulder breccia; matrix-supported pebble breccia; and gravel-rich and sand-rich graded turbidite beds (Table 2). The boulder breccias can be weakly reverse-graded and show rapid lateral transition into the other facies, all of which are interpreted as more distal manifestations of the same sedimentary events. The flow geometry and relationships between these breccia styles are interpreted by Ireland et al. (2004a) to reflect mass-flow initiation as clast-rich debris flows, with transformation via the elutriation of fines into a subsequent turbulent flow from which the turbidite and matrix-supported breccia facies were deposited.

All the Type 1 mass-flow facies contain clasts of the common and minor components of the in-situ laminated base-metal mineralised siltstone. Texturally these clasts are identical to their in-situ counterparts and are distinct from other sulphidic clasts that are of unequivocal replacement origin. In the boulder breccias, intraclasts may be the dominant clast type, and the matrix may contain abundant fine-grained sphalerite and pyrite. Dark-coloured sphaleritic and pyritic breccia matrices are distinct from pale carbonate-siliciclastic matrices, are associated with a high abundance of sulphidic clasts, and systematically occupy the lower parts of breccia units. Consequently, clasts that resemble in-situ ore facies are confirmed as genuine intraclasts incorporated into erosive mass flows before complete consolidation. Disaggregation and assimilation of sulphidic sediment in the flow contributed to the sulphide component of the dark breccia matrices. The presence of laminated sulphidic intraclasts in the mass-flow facies constrains mineralisation at HYC to the uppermost part of the seafloor sediment pile, where this material was susceptible to erosion by incoming clast-rich mass flows. That is, the presence of laminated sulphidic intraclasts in the mass-flow facies constrains mineralisation at HYC to the uppermost part of the seafloor sediment pile, where this material was susceptible to erosion by incoming clast-rich mass flows (Ireland et al., 2004a).

Type II breccia beds occur throughout the HYC Pyritic Shale Member but are most common in the upper half of the Member. Clasts are predominantly grey dololutite which occasionally contain radiating clusters of acicular crystal pseudomorphs (“coxcos”) indicative of tropical Proterozoic shelf carbonates. The clasts are similar to lithologies in the Emmerugga and Teena Dolomites and are considered to have been derived from these formations. A characteristic of this breccia type, which differentiates it from Type I and III breccias is the absence of green and red clasts, signifying that clasts in Type II breccias were not derived from the Tooganinie or lower formations, but mostly derived by erosion and collapsed of updip shallow-water cemented shelf carbonate layers. Type II breccias lack the well-developed grading seen in Type I breccias. Isopach maps (Figure 6c) and maximum clast-size plots of individual breccia beds show a close correlation and indicate the type II breccias dominate in the southeast of the HYC subbasin.

Type III breccia beds are confined to the uppermost breccia unit of the HYC Pyritic Shale Member in the HYC sub-basin and are equivalent to the Upper Breccia of Murray (1975). This unit consists exclusively of Type III breccias with the exception of several shale beds near the base. The top of the Upper Breccia is not exposed in the sub-basin, and the unit reaches a maximum known thickness of 210 m. Clasts within the breccias are completely chaotic, and there is no recognisable grading or sorting. Clasts range in size from a few millimetres up to several tens of metres. The fragment lithologies are identical to those in the Type I breccias with the notable exception that they also contain clasts of sandstone, quartzite and potash-metasomatized quartz dolerite—lithologies that are characteristic of the underlying Masterton Formation. The fragments are therefore considered to be derived from the McArthur Group (below the Barney Creek Formation) and the Masterton Formation. According to Walker et al. (1977), the most likely source of the clasts from the Masterton Formation is erosion uplifts and horsts in the Emu Fault Zone. But the same authors also state the exact source area and the direction of movement of the clasts could not be identified. In my opinion, Type III breccias are salt-ablation derived and so contain a variety of clasts lithologies plucked by the rising salt as it rose toward the surface to feed an at-seafloor allochthon.

More broadly, breccias of the updip Cooley Dolostone member, that interfinger and also overlie the HYC deposit (Figure 6a) are usually regarded as part of the Barney Creek Formation. The Cooley Dolostone is interpreted, historically, as a talus slope breccia (Walker et al. 1977, Logan 1979), containing clasts eroded from the Teena and Emmerugga Dolostones. Hinman (1995) regarded the Cooley Dolostone as a tectonic breccia, formed along reverse faults within the steep to overturned, brittle dolomitic lithologies of Teena, Mitchell Yard and Mara Dolostones(members of the Emmerugga Dolostone) as they were overthrust against and over Barney Creek Formation lithologies. Perkins & Bell (1998) interpret the Cooley Dolostone as an in situ alteration body, contiguous with, and derived from, the HYC sequence, rather than being separated from it by a thrust fault. I interpret much of the Cooley as a salt allochthon breccia derived from a salt-cored basin edge fault system, now evolved into a salt weld (Table 2).

Brine haloes and mineralisation

Regional-scale potassic alteration of Tawallah Group dolerites and sediments were documented by Cooke et al. (1998), Davidson (1998, 1999). These authors describe fluids responsible for this alteration as oxidised, low-temperature (100˚C), saline (> 20wt % NaCl equiv), Na-K-Ca-Mg-rich brines, and argue that the high salinities and the presence of hydrocarbons are consistent with brine derivation from nearby evaporitic carbonates during diagenesis.

I suggest that saline fluids feeding these haloes came not from the dissolution of evaporites in adjacent bedded carbonate hosts, but from the decay of former fault-fed thick salt allochthon tongues in positions that now are indicated by salt allochthon breccias. These breccias tie back to what were salt-lubricated fault and salt welds. The presence of salt and diagenetic haloes in these features focused tectonic movement and fluid supply in both initial extensional and subsequent compressional stages. As such, this interpretation supports a salt dissolution origin of the brine origins proposed by both Logan (1979) and Hinman (1995). The difference with their interpretations is that I envisage the brine being derived during salt flow emplacement and dissolution, tied to focused fault conduits in a mobile, suprasalt fault complex, atop or adjacent to the now-dissolved flowing and tiered salt mass. I do not think the nearby platform carbonates (with coxcos and smooth-walled cherts) ever contained significant volumes of primary evaporites.

Worldwide and across deep time, most halokinetic basinwide evaporite associations are typified by an initial extensional and loaded set of diapirs evolving into salt-cored fault welds, with subsequent reactivation of these features in compression (Warren, 2016; Chapter 6). Such a framework typifies long-term salt tectonics with inherently changing structural foci across most Phanerozoic halokinetic salt realms, as in the North Sea, the Persian (Arabian) Gulf and most circum-Atlantic salt basins. It is indicative of continental plate-edge evaporites caught up in the Wilson cycle (Warren, 2010).

Near the HYC deposit, Mn-enrichment, particularly of dolomite and ankerite in the W-fold Shale beneath the ore zone, is considered to be related to exhalation of Mn-bearing brines, associated with rifting and basin deepening, before the onset of zinc-lead mineralisation (Large et al. 1998). This too, is consistent with the salt-focused mineralisation hydrology of diagenetic ferroan and Mn-bearing hydrologies of the modern Red Sea halokinetic deeps (Schmidt et al., 2015) and the Danakhil depression in the Quaternary, when it was a marine-fed saline system (Bonatti et al., 1972).

Ridge and Cooley deposits

In the area to the east of to McArthur River HYC basin, a number of currently sub-economic Zn-Pb-Cu deposits occur, typified by the nearby Ridge and Cooley deposits (Figure 6a; Walker et al. 1977; Williams 1978). Both are similar to the Coxco deposit, being described as MVT deposits mainly hosted by dolomitic breccias, but with minor, shale-hosted concordant mineralisation in the Ridge II deposit (Figure 8; Williams 1978). Likewise, the Coxco deposit contains several million tonnes at 2.5% Zn and 0.5% Pb, in coarse-grained, stratabound galena-sphalerite-pyrite-marcasite, hosted by dolomitic breccias containing clasts of the Mara Dolostone Member, Reward Dolostone, and the Lynott Formation of the McArthur Group, within the Emu Fault Zone (Walker et al. 1977, Walker et al. 1983). Mineralisation comprises veins, “karst” and dissolution breccia fill likened to Mississippi Valley Type (MVT) mineralisation (Walker et al. 1977).

According to Williams (1978), the Emmerugga Dolostone hosts the discordant mineralisation of Cooley II deposit, while Cooley Dolostone breccias contain the Ridge II deposit (Figure 8). The Emmerugga Dolostone at Cooley II consists of massive to laminated dolostone and contains carbonaceous matter, stromatolites, oncolites, and ooids, indicating that it was deposited in a shallow-water normal-marine environment with high biologic productivity. Similarly, the Cooley Dolostone host at Ridge II is a breccia composed of randomly oriented dolostone clasts varying in diameter from a few millimetres up to several tens of metres. Some clasts have near-identical lithologies to those comprising the Emmerugga Dolostone, whereas others contain coxcos and were likely derived from the fragmentation of Teena Dolostone. The Cooley Dolostone breccia contains little depositional matrix. Clast boundaries are marked by sudden changes in features such as dolostone type and bedding-core angles, indicating that the breccia was mostly clast-supported at the time of formation. Most interestingly, drilling in the vicinity of the deposit (DDHR210) intersected a large clast of “out of sequence” dolerite (Figure 8a). Similar large salt-buoyed clasts (up to 100’s meters across) composed of Eocene dolerite occur in the salt allochthon breccias at Kuh-e-Namak-Qom (Salty Matters blog, March 10, 2015).


Two major phases of crosscutting brecciation in the area are recognised by Williams (1978) in drill core samples of discordant mineralisation from both the Emmerugga and Cooley Dolostone hosts. First generation breccias, formed during the earlier phase of brecciation, consist of angular clasts of dolostone (< 1 mm to at least 1 m in diameter) in a dark colored matrix of tiny ( < 1 µm to 20µm) anhedral dolomite grains, disseminated euhedral pyrite crystals (<50 µm in diameter) and reddish brown carbonaceous matter). The identical nature of the first generation breccias in both the Emmerugga and Cooley Dolostone hosts suggests that brecciation occurred simultaneously in both, via the same mechanism (Williams, 1978). At the time this interpretation was made, there was no “data” (paradigm) available to determine whether the brecciation in the Cooley Dolostone occurred in situ or whether it took place in the dolostone before its removal from the Western Fault Block. Today, we would likely interpret these features as reworked salt ablation breccias on the deep seafloor with infiltrated suspension clays and early-diagenetic pyrite.

Second generation breccias, formed during a later phase of brecciation, consist of angular clasts of first-generation breccias (< 1 mm to at least 10 cm in diameter) in a matrix of either veins filled with sulphide minerals and dolomite, or fine-grained (10 µm to 100 µm in diameter) anhedral dolomite grains, disseminated to massive sulfide minerals, small (on the average 500 µm x 20 µm) interlocking laths of barite or dolomite pseudomorphs after barite, and brown carbonaceous matter (Williams, 1978). Second generation breccias, although coincident with the first generation breccias, are less widespread than the earlier breccias. Again, according to Williams (op. cit.), the similarity of the second generation breccias in both the Emmerugga and Cooley Dolostones suggests a common origin. Again, they concluded there was no “data” (paradigm) available to establish the time of this brecciation relative to the deposition of the Cooley Dolostone. I would argue these “second generation” breccias represent a less distally reworked salt ablation breccia, possibly with interspace anhydrite and gypsum at the time they formed. These calcium sulphate phases facilitated the shallow subsurface emplacement of metal sulphides via bacterial or thermochemical sulphate reduction, in a way not too dissimilar to the mechanisms emplacing Pb-Zn at Cadjebut or Bou Grine ores in Tunisia (Warren and Kempton et al., 1997; Warren 2016; Chapter 15).

Allochthon Interpretation

The origin of the HYC deposit and adjacent subeconomic mineralised accumulations is still somewhat controversial and equivocal (Figure 6a; Ireland et al. 2004a,b; Perkins and Bell, 1998; Logan, 1979; Walker et al., 1977). Large et al. 1998 summarised the alternative models: 1) a sedimentary-exhalative (‘sedex’) model was proposed by Croxford 1968 and Large et al. 1998; while, 2) a syndiagenetic subsurface replacement model was introduced by Williams 1978; Williams & Logan 1986; Hinman 1995 and Eldridge et al. 1993, the latter based on sulphur isotopes. In my opinion, a third factor, namely a now-dissolved salt allochthon system, should be considered in interpretations of ore genesis and associated breccias. I interpret ore-hosting laminites of HYC deposit as DHAL laminites, and the Ridge II and Cooley II were hosted in updip regions once dominated by salt tongues and salt ablation breccias within a fault-fed salt allochthon complex surrounded by updip normal-marine shoal-water platform carbonates (Figure 9).

That is, all three deposits are related to the ongoing and time-transgressive dissolution of shallow halokinetic salt tiers. The salt tongues periodically shed mass flow deposits, triggered by seafloor instability created by the interactions of salt flow, salt withdrawal and the dynamic nature of salt and fault welds. In my opinion, the lack of equivalent breccias, DHAL laminites and halo evidence in otherwise similar deepwater sediment in Barney Creek Formation in the Glyde River Basin, some 80 km to the south-east of HYC, is why this basin lacks economic levels of base metal mineralisation (Figure 7).


Assuming that the first and second generation breccias in Type 1 and III breccias in all of the stratigraphically discordant deposits (allochthon and weld breccia), first defined by Walker et al., 1977 (Table 2) had shared salty origins, the wider distribution of the first generation breccias suggests that they formed via seafloor reworking processes acting across the whole region as a rim to discordant mineralisation (Williams 1978). Therefore, Williams (op cit.) argued geologically reasonable causes of the brecciation in the Cooley Dolostone include; movement on the Western and Emu faults, slumping of debris off the Western Fault Block, and stratal collapse due to the dissolution of evaporite minerals. I would argue for all of the above, but add that the whole Cooley Dolostone breccia system at the time the first generation breccias formed was a massive salt-flow fault-feeder system that was salt-allochthon cored and salt-lubricated. Situated at and just below the deep seafloor, salt tongue dissolution created salt-ablation breccias, while the halokinetic-induced seafloor instability instigated periodic mass flows into a metalliferous brine lake; as occurs today in the modern Red Sea deeps, the Orca basin in the Gulf of Mexico and the various brine lakes (DHAL's) of the Mediterranean Ridges (Table 2).

Breccia textures in a halokinetic salt ablation system are always two stage (Warren, 2016); the first stage of brecciation occurs as the salt tongue is inflated and spreading over the surrounds, even as its edges dissolve into ablation breccias reworked by further salt tongue movements and accumulations of contemporary salt-carapace materials (Figure 9). This first stage is typified by mass wasting piles related to the debris rims accumulating about the salt tongue edges, as debris slides downslope across the top of a continuously resupplied salt mass. The friction along the underside of the expanding salt sheets drives overturn, contortion, and brecciation of the underlying deep seafloor bed, this ultimately creates subsalt thrust overfolds (known as gumbo zones beneath the salt allochthons of the Gulf of Mexico). The second stage of brecciation is related to the dissolution of the salt itself once the salt supply is cut off by salt withdrawal and overburden touchdown.

Because allochthons are set up in the expansion stage of salt movement across the seafloor, Stage 1 breccias tend to be more widespread at the landsurface than stage 2 breccias. Stage 2 breccias form once the mother salt supply to the salt tongue or tier is cut off, the salt tongue then dissolves and final brecciation occurs, often with significant roof collapse features in any overburden layers. Similar two-stage allochthon breccias outcrop and subcrop in salt namakier provinces across Iran (Warren 2016, Chapter 7). However, unlike Iran the HYC laminites and associated breccias accumulated in a local deeper marine anoxic sump within a dominant subaqueous normal-marine carbonate shelf setting. There are also partial analogies with salt-cored Jurassic shelf carbonates and allochthon breccias in the paleo Gulf of Mexico, or the Cretaceous mineralised and ferruginised shelf-to-slope halokinetic-cored depositional system that now outcrops in the Domes Region of North Africa (Warren, 2008; Mohr et al. 2007).

Based on the sedimentology of the HYC ore host (Figure 9), I conclude that the HYC deposit accumulated as classic DHAL deposit in a salt allochthon-floored sump. Initial ore accumulation took place as metalliferous laminites in a local salt withdrawal basin. The anoxic brine-filled DHAL sump sat atop a deflating salt allochthon sheet with one of the tiers indicted by salt dissolution breccias at the Myrtle-Mara contact.

The following observations further support this conclusion; 1) the scale and deepwater setting of the deposit, 2) the fault-bound brine-fed margin to the deposit, 3) the rapid local subsidence of the sediments in the deeper water anoxic portion that constitutes the Barney Creek Fm host (HYC Pyrite member), 4) the syndepositional nature of the inter-ore polymict mass flow breccias, 5) the presence of syndepositional barite and Mn haloes from a diagenetically imposed oxidised saline set of pore waters hosted in what were formerly normal-marine sediment pore fluids.

Salt flowing from an allochthon sheet into salt risers in the Emu-Western fault region drove fault-bound rapid subsidence that created local deeper-water anoxic brine-filled sumps in an otherwise healthy marine carbonate shelf (see Salty Matters blog, April 29, 2016, for a salt-controlled structural analogy in the Red Sea). The fault-controlled salt risers allowed brine to escape onto the seafloor at Barney Creek time and to flow across the seafloor into the large DHAL sump that is today the HYC deposit (Figure 9). With time, the salt risers evolved in salt welds and ultimately into fault welds with salt-ablation breccia textures.

The characteristic Fe-Mn and baryte haloes, along with skeletal halites, in what were porous sandstone aquifers intersected by hypersaline waters from the rising and dissolving salt mass are today indicators of the geometry of the former briny plumbing. In the Barney Creek Fm., the occurrence of the Mn and ferruginous haloes indicate the fault-conduit aquifer focus to the suprasalt brine flow and the level of hypersaline brine intersections. There are also transitions into more-typical more-oxidised marine pond and pore water masses in the upper levels atop the DHAL waters and around the edge of its brine curtain.

Williams (1978) concluded the less widespread second generation breccias in the Cooley Dolostone wedge likely formed by processes that acted only locally on the first generation breccias. I agree and would argue that a later DHAL mineralisation focus, during the creation of a later generation of breccias, was the transition from a salt feeder supplying a canopy of allochthon tongues along the Emu Fault region into a system that became first a salt weld, then a fault weld as the mother salt supply was lost (Table 2).

Williams (op. cit.) noted that the association of the two breccia generations, and the occurrence of base metal sulfide minerals and barite, presumably brought in from an outside source, in the matrix of the second generation breccias suggest that the later breccias formed by solution collapse following the introduction of mineralizing solutions into the porous, first generation breccias. I am in complete agreement with this conclusion. In addition, we now have a set of salt-related mechanisms and time-transgressive paradigms that explain the transition from one breccia generation tied to a syndepositional DHAL-related succession that we classify as the sedex brine pool stage that is forming the HYC deposit. With time and salt dissolution/source depletion, we pass to the next generation of breccias, which are linked to a fault weld, evaporite-collapse sub-economic set of MVT deposits (e.g. Cooley II Ridge II and Coxco deposits).

In my opinion, halokinesis created shallow allochthonous salt tiers at the time the normal-marine Emmerugga and Teena Dolostones. Salt withdrawal from allochthon sheets emplaced below the shallow sea floor caused it to deepen locally, this facilitated deposition of thickened intervals of deeper water, more siliceous deposits, as defined by thickness and mineralogical/ colour changes in the W-Fold shale and Barney Creek Formation (Figure 9). Where the brine accumulated in the deepened seafloor depression, which was the HYC DHAL sump, it lay atop a salt withdrawal basin, associated with flow of allochthon salt into the proto-Western Fault (now a deformed fault-weld). The stratigraphic level of the withdrawal is indicated by the allochthon collapse breccia seen at the top of the Myrtle Shale.

The salt-brine focusing time-transgressive halokinetic architecture of the mineral system allowed metal-bearing chloride rich brines circulating in the buried sediments of the basin to access and replace the reduced pyritic and bituminous laminite of the DHAL. As well as ponding in DHALs, some of the same metal-bearing brines exploited the presence of fractionally dissolved interclast calcium sulphate within diapir collapse breccias. So a similar set of redox interfaces drove discordant mineralisation in second generation breccias in the nearby Cooley, Coxco and Ridge deposits. At that time, some of the collapsing crests on the diapiric basin margin perhaps had subaerial crests. We interpret the smaller-scale currently-subeconomic Cooley, Coxco and Ridge deposits as combinations of passive infill, vein and replacement mineralisation in diapiric, dissolution and salt collapse breccias. The Pb-Zn ore, and its collapse-induced host rock, formed in a diagenetic setting much like that in suprasalt circum-diapir MVT deposits hosted in caprocks breccia and peripheral Cretaceous seafloor DHAL laminites in the Bahloul Formation of Northern Africa (see Warren 2016; Chapter 15).

The intimate relationship between breccias and mineralisation across the McArthur River region, including clasts of ore in sedimentary and diagenetic breccias, can be explained, by continual halokinetic salt movement before, during, and after the main episode of laminite Pb-Zn ore formation. This interpretation of both inter-ore “sedimentary” and Cooley Dolostone member breccias across the region reconciles what were seen as previously conflicting primary versus time-transgressive relationships (e.g., Williams 1978; Perkins & Bell 1988).

The characteristic Mn and baryte haloes, along with skeletal halites, in what were porous sandstone aquifers intersected by hypersaline waters from the rising and dissolving salt mass are today indicators of the geometry of the former briny plumbing. In the Barney Creek Fm., the occurrence of the Mn and ferruginous haloes indicate the aquifer and the level on hypersaline brine intersections with the more typical more oxidised marine water mass and pores water at levels atop the brine lake.

Williams (1978) concluded the less widespread second generation breccias in the Cooley Dolostone wedge likely formed by processes that acted only locally on the first generation breccias. I agree, and would argue that the later mineralisation focus, during the creation of the second generation of breccias, was the transition from a salt feeder supplying a canopy of allochthon tongues along the Emu Fault region into a system that became first a salt weld, then a fault weld as any ongoing mother salt supply was lost. Williams (op. cit.) in a discussion of the Ridge and Cooley deposits noted that the association of the two breccia generations, and the occurrence of base metal sulfide minerals and barite in the matrix of the second generation breccias, presumably brought in via fluids with an outside source. He suggests that later breccias formed by solution collapse following the introduction of mineralising solutions into the porous, first generation breccias. I agree also with this conclusion but would also place it in the typical saline baryte ore association seen in many salt diapir provinces such as the Walton-Magnet Cove region of Nova Scotia, or the Oraparinna Diapir in the Flinders Ranges, South Australia (see Warren 2016, Chapter 7 for detail on theses and other similar baryte deposits).

In addition, we now have a set of salt-related mechanisms and time-transgressive paradigms that explain the transition from one breccia generation tied to a syndepositional DHAL-related succession we classify as the sedex brine pool that is the HYC deposit, to the next generation of breccias that are linked to a fault weld, evaporite-collapse sub-economic set of smaller scale MVT deposits (e.g. Cooley II Ridge II and Coxco deposits).

In my opinion, halokinesis created shallow allochthonous salt tiers at the time the normal-marine Emmerugga and Teena Dolostones were deposited. Salt withdrawal below the shallow sea floor caused it to deepen locally, this facilitated deposition of thickened intervals of deeper water, more siliceous deposits defined by the W-Fold shale and Barney Creek Formation (Figure 9). Where the brine accumulated in the deepened seafloor that was the HYC DHAL sump it lay atop a salt withdrawal basin, associated with flow of allochthon salt into the proto-Western Fault (now a deformed fault- weld) with the stratigraphic level of the withdrawal indicated by the allochthon collapse breccia at the top of the Myrtle Shale.

The salt-brine focusing time-transgressive halokinetic architecture of the mineral system allowed metal-bearing chloride-rich brines circulating in the basin to access and replace the reduced pyritic and bituminous laminite of the DHAL. As well as ponding in DHALS, some of the same metal-bearing brines exploited diapir collapse breccias and drove discordant mineralisation and second generation breccias in the nearby Cooley, Coxco and Ridge deposits. At that time, some of the collapsing crests on the diapiric basin margin perhaps had subaerial crests. We interpret the smaller-scale currently-subeconomic Cooley, Coxco and Ridge deposits as combinations of passive infill, vein and replacement mineralisation in diapiric, dissolution and collapse breccias. The Pb-Zn ore, and its collapse-induced host rock, formed in a diagenetic setting much like that in suprasalt circum-diapir MVT deposits hosted in caprocks and Cretaceous seafloor laminites of the Bahloul Formation of Northern Africa (see Warren 2016 Chapter 15).

The intimate relationship between breccias and mineralisation across the McArthur River region, including clasts of ore in sedimentary and diagenetic breccias, can be explained, by continual halokinetic salt movement before, during, and after ore formation.

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Salt Dissolution (4 of 5): Anthropogenically-enhanced geohazards

John Warren - Thursday, November 30, 2017

 

Introduction

As we saw in the previous article the dissolution and collapse of nearsurface and at-surface salt is a natural and ongoing process. When salt bodies experiencing natural dissolution and alteration are penetrated by drilling or parts of the salt mass are extracted in a poorly supervised fashion, the resulting disturbance can speed up natural solution and collapse, sometimes with unexpected environmental consequences (Figure 1; Table 1). To minimise the likelihood of unexpected environmental consequences tied to enhanced rates of salt dissolution in the vicinity of engineered structures the same rule applies as is applied to safe salt mining, namely "Stay in the salt" (Warren, 2016) In this context, here is a quote from Zuber et al. (2000) in a paper dealing with flooding and collapse events in Polish salt mines.


“Catastrophic inflows to salt mines, though quite frequent, are seldom described in the literature, and consequently students of mining and mine managers remain, to a high degree, ignorant in this respect. Contrary to common opinion, inflows are seldom caused by unavoidable forces of nature. Though some errors were unavoidable in the past, modern geophysical methods are, most probably, quite sufficient to solve the majority of problems (e.g., to determine a close presence of the salt boundary). Detailed study of the recent catastrophic floods, which happened in Polish salt mines, shows that they usually occur, or have strong negative impacts, due to human errors. Most probably similar human errors caused catastrophic inflows to salt mines in other countries. It seems that a knowledge of the real history of catastrophes, better education of mine engineers and the application of modern geophysical methods could lead to the reduction of floods in salt mines.”

This article, the 4th of 5 in Salty Matters, focuses on the anthropogenic enhancement of salt dissolution, however, it discusses only a few of the many documented examples of problems that result from enhanced dissolution brought about by human activity. For a more comprehensive documentation of relevant case histories and an expanded discussion that includes mine collapse, brinefield subsidence and collapse and industrial accidents associated with salt cavity storage, the interested reader should refer to Chapters 7 and 13 in Warren (2016).


As wells as mine floods and mine losses, poorly monitored then abandoned brinefield and solution cavities can be areas with major environmental problems especially where old extraction wells are undocumented and unmonitored. Like areas of natural solution collapse; they can become zones of catastrophic ground failure. This is especially problematic if located near cities or towns (Tables 1, 2). Some of the most outstanding examples of how not to solution mine a resource and how not to control ground subsidence effects are to be found in east European countries that are still trying to deal with the environmental outcomes of being former satellite states of the Soviet Union. But caving problems have been tied also to mines and boreholes in the United States and Canada, where once again human error, greed or ignorance has created many of these problem structures (Table 1 and Figure 1).


Ocnele Mari Brinefield, Romania

SALROM, a government-owned company, solution mines Badenian (Miocene) halite in the Valcea Prefecture of Romania (Figure 2). Production from Field 2 was shut down on March 5, 1991, after significant earth vibrations were noted by SALROM workers. A subsequent sonar survey showed that poorly monitored brinefield leaching between 1971 and 1991 had created a gigantic merged cavern as salt pillars separating adjoining caverns were inadvertently dissolved (Figure 2); the upper parts of the captive boreholes 363, 364, 365, 366, 367 and 369 had merged into a common cavern. The cavern was filled with some 4.5 million cubic metres of brine, was less than a hundred metres below the landsurface, was more than 350 metres across and was overlain by loosely consolidated sandy marls (von Tryller, 2002; Zamfirescu et al., 2003). The cavern was overlain by a bowl of subsidence and, even though mining operations in the region of the cavern completely stopped in 1993, the ground above continued to subside to a maximum of 2.2 metres prior to the September 2001 collapse. In the period 1993 - 2001 the cavern continued to enlarge and the northern part of the cavity expanded by some 25-35 m (see inset in Figure 3 showing sonar surveys 1995 -2001).


The 1993 sonic survey of the cavern showed it was so large and shallow that its roof must ultimately collapse, predicting when was the unknown. The fear was that if it collapsed catastrophically, it could release a flood of at least a million cubic metres of brine. Brine could be released either over several hours (least damaging scenario) or instantly forming a wave of escaping water several metres high that would flood the nearby Sarat River valley for many kilometres downstream. There were 22 homes on top of the cavern that would be immediately affected; also at risk were the hundreds and perhaps thousands of residents in the area of the saltworks and river valley below, as well as the local ecology and the civil/industrial infrastructure. Ongoing brine delivery from other nearby operational SALROM caverns to the large Oltchem and Govora chemical plants would also be disrupted. When collapse did ultimately occur (twice in the period 2001 - 2004), each episode took place over a number of hours and so a catastrophic wave of water did not eventuate.

According to Solution Mining Research Institute, there was no good engineering solution to prevent the Ocnele cavern roof from collapsing. Prior to collapse, brine in the cavern exerted pressure against the roof helping to hold it up; removing even a small amount of brine would remove hydraulic pressure that could possibly trigger a catastrophic collapse. Nor was it practical to fill the cavern with sand and industrial wastes as suggested by the Romanian government. It would take too long and it was unsafe to place men and equipment on top of the still expanding cavern. Even if it could be done, only the areas directly below the injection wells would be filled. The best solution was to construct a dam as close to the cavern as possible, and then perhaps trigger a controlled collapse by pumping out the cavern brine.

A partial collapse of the cavern roof atop well 377 occurred on 12 September of 2001, the ensuing brine flood killed a child and injured an older man. The collapse began at 7 pm with brine spilling out of wells 365 and 367. Southward of well 377 a collapse cone some 10 metres across started to form and fill with brine. The cone continued to expand and fill with water until its southern lip was breached. Water spilled out of the cone and down the hill slope. The flow rate of the expelled flow reached a maximum of 17 m3/sec at 3 am on September 13th, with flow continuously exceeding 10 m3/sec for a period of more than 6 hours around that time. By 7 am flow was down to 4-5 m3/sec and by 7 pm, 24 hours after the flow began the flow rate was 0.4-0.5 m3/sec. Some 24 hours after the onset, roof collapse had formed a water-filled lake with an area of 2.4 ha (Figures 2, 3; Zamfirescu et al., 2003).


A second roof collapse occurred on July 13, 2004. Realitatea Romaneasca, Romania, reported that ground collapse occurred at 9:30 pm July 13, 2004, possibly triggered by heavy, recent rains. The wellhead of bore 365 was destroyed in this collapse, and many other well heads in Field 2 had been destroyed by ongoing subsidence in the period 2001-2004. This collapse was the culmination of a series of collapse-related events that began on Monday July 11. It was planned that a purpose-built earthen dam would contain the brine flood. But when the roof fall occurred at 9:30 pm, the dam was breached 30 minutes later by some 250,000 m3 of salt water. Flow rate through the breach reached a maximum of 6 m3/sec. Collapse did not occur until just after local authorities had evacuated people from 50 homes in the area around the cavern. Most of the brine forced out by the roof collapse escaped into the Sarat (Olt) River. The government attempted to dilute the effects of this flush of salt water by releasing fresh water into the river from nearby dams. Fortunately, unlike the 2001 event, there was no loss of life.


Induced collapse, Gellenoncourt saltworks, France

Gellenoncourt is one of three sites that exploit halite beds of the Lower Keuper in the eastern part of the Paris Basin. The deposits extend from Cezanne in the west to Nancy with a length of 250 km and a width of 50 to 70 km. Salt production utilising solution mining techniques is focused on three sites along the Meurthe river, which all lie immediately to the east of Nancy.

On March 4, 1998, a sinkhole more than 50 m across and 40 m deep formed atop the SG4 and SG5 brinefield caverns in the Gellenoncourt saltworks near Lorraine, France (Figure 4; Buffet, 1998). It was an induced collapse designed to prevent a possible uncontrolled future ground collapse. The problem started in 1967 with the beginning of the exploitation of Triassic salt layers in the Keuper Fm. In total the Keuper is more than 150 m thick with five salt layers at its base, passing up into variegated and poorly consolidated marls and sandstones and capped by the Dolomite du Beaumont. The top of the salt is some 220 m below the surface and divided into five beds numbered 1 to 5 from top to bottom, with the solution mining program designed to leach beds 1 through 3 in the region of SG4 and SG5 caverns (Figure 5). Five wells, SG1 to SG5, were joined using hydrofracturing in February 1967. Theoretically, the process was designed to leave a substantial salt pillow atop all the cavities and so separate the solution cavities from the overlying marls (Figure 5; 1967-1971). The SG4 and SG5 caverns unexpectedly joined in 1971. Brine injection to these two wells was stopped, but crossflowing brines flowing to the producing SG1 well continued to excavate these two caverns. By 1982 the salt cushion in the roof of the SG4-SG5 cavern had completely dissolved away, placing the cavity roof in direct contact with the base of the marls.


From 1982 until October 1992 there was no further upward growth of the cavern roof. Then a 25 m-thick section of the variegated marls in the roof broke free and fell to the cavern floor, leaving a large section of the cavity roof in direct contact with the brittle Dolomite de Beaumont. This stiff dolomite layer prevented any further immediate collapse of the roof and consequent propagation of the cavity to the surface. But continued growth of the roof span beneath the dolomite would mean a later, larger, perhaps catastrophic collapse. In 1995 the operator tried to induce a controlled collapse by placing a submerged pump in the cavity and pumping out brine to create an exposed upper face. But this didn’t work. The next approach was to further enlarge the roof span by injecting 300,000 m3 of freshwater into the cavity. Collapse occurred on March 4, 1998, forming a 50 m wide crater. To protect the surrounding countryside from any brineflood damage a dam was constructed to capture any brine overflow, but in the actual event it was not needed.

Retsof Mine, New York State, USA

The Retsof Mine collapse is perhaps one of the best documented examples of an operational mine being lost to dissolution features and consequent flooding, yet even here the exact causes of the mine loss are still argued. Was it because of the intersection of the expanding mine workings with a natural water-saturated salt dissolution and fracture system, or the insection of the mine workings with brine filled cavities formed by wild brining operations in the 1800s, or was it a result of mine roof instabilities related to a change in room and pillar sizes?

At the time it was operational, the 24 km2 area of subsurface workings in the AKZO-Retsof salt mine made it the largest underground salt mine in the USA, and the second largest salt mine in the world (Figure 6, 7).


Operational history

Retsof Mine started operating in 1885 after completion of the 3.7× 4.9-meter wide, 303.5-meter-deep Shaft #1. The mine claimed an initial 5,460-metric-tons per day hoist capacity (Goodman et al., 2009). Early main haulage ways were driven east and west while production headings were driven north (updip) for salt-tramming ease. Room heights in the 6-meter-thick salt bed were 2 to 4 meters, with salt left in both the floor and roof. Four-meter-high rooms were worked in two benches. Rooms were 9.2 meters wide and separated by 9.2-meter pillars. At that time there was no timbering, the mine was dry, mine air temperature was 17°C, and the mine was largely gas-free.

By 1958, the Retsof Mine was connected to the former Sterling Mine for ventilation and emergency escape purposes [Figure 6a; Gowan et al., 1999]. By the late 1960s, the mine had advanced beneath the modern Genesee River and Valley (Figure 6b). By the early 1970s, the Retsof Mine operators had installed an underground surge bin, fed by a new conveyor system, and the old rail-haulage system was eliminated. Mainline conveyors led to yard or panel belts feeding each mining section, where a Stamler feeder-breaker crushed salt delivered by diesel-powered Joy shuttle cars. In the early 1980s, the shuttle cars were gradually replaced by load-haul-dump vehicles (LHDs). In 1969, Netherlands-based Akzo Corporation acquired International Salt Company and operated the Retsof mine until its abandonment due to flooding in 1995.

During April 1975, an explosion occurred in the original Sterling B Shaft during efforts to control water inflow into the Retsof Mine from this abandoned and partially collapsed shaft (Goodman et al. 2009). The leaky B Shaft had not been used or maintained for years. By 1975, International Salt was concerned that freshwater inflow from the B Shaft could pose a salt dissolution, collapse, and flooding risk to the then-connected Retsof Mine. Removal of a partial shaft blockage of timber and rock debris was attempted as a means of regaining airflow needed to safely access, rehabilitate, and grout off the water inflow to the shaft and mine below. A maintenance crew attempted to dislodge the shaft obstruction by pushing a large boulder into the shaft that was to drop down and knock through the debris. A methane explosion occurred upon impact. The upward force of the explosion killed four people on the surface near the shaft collar and injured others. On November 19, 1990, a roof fall resulted in two fatalities. Deformation and fracture of roof salt can occur because of a concentration of stresses; i.e., punching of the roof by stiff pillars. After the fatalities, the mine tested smaller, yielding pillars to alleviate roof falls (Figure 6b). Positive test results led to the adoption of a yield-pillar design.

The Retsof mine was lost to water flooding in 1994-1995.Before abandonment, the mine had been in operation since 1885, exploiting the Silurian Salina Salt and prior to shut down was producing a little over 3 million tons of halite each year. At that time it supplied more than 50% of the total volume of salt used to de-ice roads across the United States.

Geology and hydrology in the vicinity of Retsof Mine

The Genesee Valley sediments preserve evidence of several complex geologic processes that include; (1) tectonic uplift of Palaeozoic sedimentary rocks and subsequent fluvial down cutting, (2) waxing and waning glacial events that drove erosion of bedrock and the subsequent deposition of as much as 750ft of glacial sediments; and (3) ongoing erosion and deposition by postglacial streams (Figure 7a; Yager, 2001; Young and Burr, 2006). The Genesee Valley spans through western New York north to south from Avon, NY to Dansville, NY, including the Canaseraga Creek up through its mergence with Genesee River. A detailed section from Palaeozoic rocks and younger have been recorded in the Genesee River Valley (Figure 7a, b); however, detailed analysis of glacial sediments and till are still somewhat scarce. The B6 salt bed (Retsof Bed) of the Vernon Formation was the salt unit extracted at the Retsof Mine (Figures 7b). Several other salt layers exist in the Salina Group both above and below the B6. These salt layers include two horizons in Unit D at the base of the Syracuse Formation approximately 50 m (160 feet) above the B6 salt level.

Quaternary-age sediment in the Genesee Valley consists mostly of unconsolidated glacial sediments ranging up to 750 feet thick. These sediments encompass gravel, sand, silt and clays that were deposited mostly during the middle and late Wisconsin deglaciation and filled the lower parts of the pre-existing glacial scour valley. End moraines consisting of glacial debris were deposited in lobes to the south of the slowly retreating glacier. As the glacier had scoured through the valley, carving out bedrock and accumulating sediment, steep-sided valley walls were cut and pro-glacial lakes formed. The glacial lake sediments are dominated by muds, but also include large boulders and cobbles carried to the lake depressions by glacial ice. Fluvial sediment from the Genesee River and Canaseraga Creek also drained into these glacial lakes. A final pro-glacial lake formed as the Fowlerville end moraine was deposited. The Fowlerville end moraine extends approximately 4.5 to 8 miles north of the Retsof collapse site (Figure 7a). The various glacial lakes and moraines disrupted the normal flow fluvial patterns of most local drainages and creeks in the valley. Alluvium is the uppermost layer of the surface and is variable in thickness throughout the valley, but normally ranges about fifty feet thick and is still being deposited across the Genesee River Valley floodplain (Yager, 2001).

The aquifer system is hosted within the glacial valley-fill and consists of three main aquifers separated by two confining layers. It is underlain by water-bearing zones in fractured Palaeozoic bedrock (Yager, 2001). The glacial aquifers are bounded laterally by the bedrock valley walls. The uppermost aquifer consists of alluvial sediments 20 to 60 ft thick (unit 1 in Figure 7b); the middle aquifer consists of glaciofluvial sand and gravel less than 10 ft thick (unit 3 in Figure 7b); and the lower aquifer consists of glaciofluvial sand and gravel about 25 ft thick overlying the bedrock valley floor (unit 5 in Figure 7b). These aquifers are separated by aquitards dominated by muds and clays (Units 2 and 4 in Figure 7b).

The now abandoned Retsof Mine lies 550 to 600 ft below the eroded valley floor (Figure 7b). Hence, the upper and middle aquifers are separated by an upper confining layer of lacustrine sediments and till as much as 250 ft thick, and the two confined aquifers are separated by a lower confining layer of undifferentiated glaciolacustrine sediments as much as 250 ft thick. The principal water-bearing zone in the bedrock overlying the mine consists of fractured carbonates and sands near the contact between the Onondaga and Bertie Limestones. The fractured aquifer that occurs at this level in the stratigraphy supplied a significant volume of the water that ultimately flooded the Retsof Mine. The glacial aquifers are hydraulically connected at the edges of the confining layers and in subcrop zones, where water-bearing zones in the bedrock intersect a fractured and karstified bedrock surface.

Ground water within the valley generally flowed northward and updip before the mine collapse (Yager, 2001). The hydraulic head distribution in the confined aquifers under natural (precollapse) conditions is assumed to have been similar to that in the upper aquifer before the collapse, but water levels in the confined aquifers were probably above the water table beneath the valley floor. Much of the ground water reservoired along the fractured Onondaga/Bertie Limestone contact also flowed northward to escape at the Bertie Limestone subcrop, now located in the valley north of the Fowlerville Moraine (Figure 7c).

Water influx tied to changes in room and pillar mining?

In 1993, ceiling falls began to occur in rooms in the deepest part of the Retsof mine near its southern boundary (Figure 6a; Yager et al., 2009). In response, the mine owner, AkzoNobel Salt Incorporated (ANSI), turned to an innovative “yielding pillar” mining technique that utilised many narrow (20 feet × 20 feet) pillars rather than few wide ones in the mined section (Figure 6b). Geotechnical analyses indicated that the resulting configuration would allow the salt pillars to slowly yield and create a “stress envelope” in the surrounding bedrock to support the entire mined room.

Closure monitoring was conducted in the yield-pillar test panels and the two full-scale panels during mining to measure panel behaviour and to see if the new design mitigated the floor and roof problems being experienced in the large pillar area of the mine (van Sambeek et al., 2000). Monitoring initially indicated that room closure rates were slightly greater than expected, but had an overall character (trend) of steadily decreasing rates, which is consistent with stable conditions. This trend changed dramatically to a rapid and unstable closure rates in the final weeks leading up to the inflow. The change in trend was initially obscured by fluctuating closure rates because salt extraction was occurring between the two yield-pillar panels as the monitored abutment pillar was isolated. Whereas the closure rates were expected to decrease after this mining was complete, they did not; in fact, they increased. This change in panel character later interpreted to indicate that a pressure surcharge existed or developed over two of the yield-pillar panels prior to the in- flow (Gowan et al., 1999).

 

Loss of roof stability and flooding

In November 1993, strain measurements in a yielding-pillar area within the mine indicated a larger than expected deformation of salt near the eastern wall of room 2 Yard South (Figure 6a, b). Mining in the area was halted as ceiling falls of salt continued during the next four months. On March 12, 1994, a magnitude 3.6 seismic event, caused by a large roof collapse, was detected by seismometers more than 300 miles away. Mine workers attempted to enter room 2 Yard South but found it was blocked by a pile of rock rubble within the formerly mined room and that saline water entering via fractures in the mine roof. Over the next several weeks, Akzo made concerted attempts to save the mine by pumping water out and drilling around the collapse area to inject cement grout so as to stabilise the collapsed room and prevent a further inflow of water. Meanwhile, unstable shale layers overlying room 2 Yard South sagged and collapsed to form a 300-foot-diameter zone of rock rubble that slowly propagated upward through overlying layers of shale (Figures 6b, 8a, b). This column of rock rubble is referred to as a rubble chimney.

The propagating rubble chimney eventually reached a layer of carbonate (limestone) rock that was strong enough to temporarily resist further collapse, stopping further the rubble chimney’s upward progression. At this point, the flow of water into the mine stabilised at about 5,500 gallons per minute. Water entering the mine was saline and probably a mixture of saline water from the shale and a prominent fracture zone aquifer within the Onondaga and Bertie Limestones (Figures 7c). By the end of March 1994, tons of cement grout had been injected into the mine and the rubble chimney through nearly 30 boreholes drilled in the collapse area, but these efforts failed to stem the rate of water flowing into the mine and the inflow was becoming increasingly less saline.


On April 6, 1994, the limestone rock layer collapsed, and 550 feet of unconsolidated sediments in the Genesee River valley quickly slumped downward into the resulting cavity, forming a sinkhole at the land surface, more than 15 feet deep and several hundred feet across (Figure 8, 9a b). The collapse of the limestone rock was like pulling the plug in a bathtub—it allowed groundwater from a fresh-water aquifer at the base of the unconsolidated glacial sediments (the lower confined aquifer (Figure10) to drain downwards through the rubble chimney and into the mine. By mid-April, a second collapse occurred in an adjacent room (11 Yard West; Figure 6b). On May 25, a drilling crew working above room 11 Yard West felt tremors and removed their drill rig, and themselves, just before this second sinkhole formed at the land surface. This one had a surface expression that was more than 50 feet deep and several hundred feet across (Figures 6a, 8a, 9a,b). The discharge from the aquifer through both rubble chimneys increased the flow of water into the mine to about 18,000 gallons per minute.

Water began to fill the southern end of the mine and then spread steadily northward, dissolving the bases of the salt pillars that supported the mine ceiling (Figure 6a). As the pillars gave way, the southern part of the mine began to collapse, causing the land surface above it to subside. The greatest subsidence (more than 15 feet) was beneath the two sinkholes, which altered the channel of Beards Creek, allowing surface water to fill the sinkholes. The surface water did not flow downward to the mine, however, because hundreds of feet of fine-grained sediments underlie the Genesee River valley. The instability also forced the closure of the U.S. Route 20A bridge over Beards Creek; the southern end of the bridge eventually subsided by 11 feet (Figure 9c, d). The bed of the Genesee River 1 mile north of the collapse areas subsided by as much as 5 feet and altered the pattern of sediment scour and deposition along a 1.5-mile reach downstream of Beards Creek.

Events indicating loss of mine

The eventual loss of the Retsof Salt Mine occurred in stages, driven first by “out of salt” roof breaches, followed by ongoing salt dissolution of the water-encased salt pillars in the flooded mine. It began in the early morning hours of March 12, 1994, with a magnitude 3.6 earthquake. The quake was caused by the catastrophic breakdown of a small mine pillar and panel section some 340 meters below the surface and was accompanied by the surface collapse of an area atop the mine that was some 180 by 180 meters across and 10 meters deep. This all occurred at the southern end of the mine near the town of Cuylerville. A month later, on April 18, an adjacent mine room collapsed to form a second collapse crater (Figure 6a, b) The initial March 12 collapse in the mine was accompanied by an inrush of brine and gas (methane) and by a sustained intense inflow of water at rates in excess of 70 m3/min, via the overlying now fractured limestone back (Gowan and Trader, 2000).

In a little more than a month, the two steep-sided circular collapse features, some 100 meters apart, had indented the landscape above the two collapsed mine rooms (Figures 6, 8, 9). The northernmost collapse feature, which was more than 200 meters across, included a central area that was about 60 meters wide and had subsided about 6 to 10 meters. The southernmost feature, which was about 270 meters in diameter, included a central area that was about 200 meters wide and had subsided about 20 meters (Figure 6b). Fractures extending up from the broken mine back created hydraulic connections between aquifers, which previously had been isolated from each and so provided new high volume flow routes for rapid migration of perched groundwaters into the mine level.

Water flooded the mine at rates that eventually exceeded 60,000 litres per minute and could not be controlled by pumping or in-mine grouting. By January 1996 the entire mine was flooded. Associated aquifer drawdown caused inadequate water supply to a number of local wells in the months following the collapse; the fall in the water table as ground waters drained into the mine in effect meant some water wells went dry (Figure 8c; Tepper et al., 1997).

Aside from the loss of the mine and its effect on the local economy, other immediate adverse effects included abandonment of four homes, damage to other homes (some as much as 1.5 kilometers from the sinkholes), the loss of a major highway and bridge, loss of water wells and prohibition of public access to the collapse area (Figure 9). Land subsidence, possibly related to compaction induced by aquifer drainage to the mine, even occurred near the town of Mt. Morris some 3 miles south-west of the collapse area. Longer term adverse effects are mostly related to increasing salinization of the lower parts of the Genessee Valley aquifer system in the vicinity of the mine (Figure 10; Yager, 2013).

 

What caused the loss of the mine?

Post-mortem examination of closure data from the two failed mine panels has been interpreted as indicating an anomalous buildup of fluid pressure above the panels in the period leading up to their collapse (Gowan et al., 1999). The initial influx of brine and gas following the first collapse coincided with the relief of this excess pressure.

Gowan and Trader (1999) argued for the existence of pre-collapse pressurised brine cavities and gas pools above the panels and related them to nineteenth-century solution mining operations. They document widespread natural gas and brine pools within Unit D of the Syracuse Formation approximately 160 ft above the mined horizon in the Retsof Mine. The satellite image also shows that collapse occurred in a pre-existing landscape low that defined the position of Beard Creek valley above the mine (Figure 6a). Brine accumulations likely formed in natural sinks, long before salt solution mining began in the valley. Salt in the shallow subsurface dissolved naturally, driven by the natural circulation and accumulation of meteoric waters along vertical discontinuities, which connected zones of dissolving salt to overlying fresh water aquifers (see Warren, 2016, Chapter 7 for a detailed documentation of this salt related hydrology and geomorphology).

Gowan and Trader (2003) argued that daylighting sinkholes had formed by the down-dropping of the bedrock and glacial sediments into pre-existing voids created by the dissolution of salt and the slaking of salt-bearing shale upon exposure to fresh water. It is likely that the extent of these brine filled voids was exacerbated by the “wild-brining” activities of salt solution miners in the 1800’s.

Nieto and Young (1998) argue that the transition to the yield pillar design was a contributing factor to the loss of mine roof integrity. Loss of mechanical integrity in the roof facilitated fracturing and the influx of water from anthropogenic “wild brine” cavities. The exact cause of the loss of roof integrity and subsequent mine flooding is still not clear. What is clear is that once the Retsof mine workings passed out of the salt mass, and into the adjacent non-salt strata, the likelihood of mine flooding greatly increased.

Even so, the loss of the Retsof salt mine to flooding was a total surprise to the operators (Van Sambeek et al., 2000). The mine had operated for 109 years with relatively minor and manageable incidents of structural instability, water inflow, and gas occurrences. A substantial database of geological information was also collected throughout the history of the mine. It was this relatively uneventful mine history and the rich technical database that provided support for pre-inflow opinions by mine staff that there was no significant potential for collapse and inundation of the mine. The Retsof collapse took place in a salt-glacial scour stratigraphy and hydrology near identical to that in the Cayuga Mine region.

 

Patience Lake Potash Mine flood

In the 1970s the Patience Lake potash mine operation, located on the eastern outskirts of Saskatoon, Canada, encountered open fractures tied to a natural collapse structure and was ultimately converted to a successful solution mining operation (Figure 11). Grouting managed to control the inflow and mining continued. Then, in January of 1986, the rate of water inflow began to increase dramatically from the same fractured interval (Figure 12; Gendzwill and Martin 1996).

At its worst, the fractures associated with the structure and cutting across the bedded ore zones were leaking 75 m3/min (680,000 bbl/day) of water into the mine. The water was traced back to the overlying Cretaceous Mannville and possibly the Duperow formations. Finally, in January 1987 the mine was abandoned. It took another six months for the mine to fill with water. Subsequent seismic shot over the offending structure suggested that the actual collapse wasn’t even penetrated; the mine had merely intersected a fracture within a marginal zone of partial collapse (Gendzwill and Martin 1996).

Part of the problem was that the water was undersaturated and quickly weakened pillars and supports, so compromising the structural integrity of the workings. The unexpected intersection of one simple fracture system resulted in the loss of a billion dollar conventional potash mine. Patience Lake mine now operates as a cryogenic solution mine by pumping warm KCl-rich brine from the flooded mine workings to the surface. Harvesting of the ponds takes place during winter after cryogenic precipitation of sylvite in at-surface potash ponds (Fig. 11).

 

Unlike the Patience Lake Mine flood, there was a similar episode of water inflow in the nearby Rocanville Potash Mine. But there a combination of grouting and bulkhead emplacement in succeeded in sealing off the inflow, thus saving the mine (see Warren 2016 for detail). Unlike Patience Lake, the brine from the breached structure in Rocanville was halite-saturated, so limiting the amount of dissolution damage in the mine walls. Different outcomes between the loss of the Patience Lake Mine and recovery from unexpected flooding in the Rocanville Mine likely reflects the difference between intersecting a natural brine-filled dissolution chimney that had made its way to the Cretaceous landsurface and is now overlain by a wide-draining set of aquifer sediments, versus crossing a blind dissolution chimney in a saline Devonian sediment surround that never broke out at the Cretaceous landsurface. Understanding the nature of the potential hydrological drainages and water source is a significant factor in controlling unexpected water during any salt mine expansion.


Lake Peigneur, Louisiana

Lake Peigneur is a natural water-filled solution doline that overlies the dissolving crest of the Jefferson Island Salt Dome Figure 13). The most recently risen part (salt spine) of the Jefferson Island stock crest, just west of the town of New Iberia, Louisiana, is now 250 m (800ft) higher than the adjacent flat-topped salt mass, which is also overlain by a cap rock. The boundary shear zone separating the spine from the less active portion of the crest contains a finer-grained “shale-rich” anomalous salt zone that had been penetrated in places by the former Jefferson Island mine workings. The known salt anomaly (BSZ) 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, in a mining scenario much like the fault shear anomaly, as mapped by Balk (1953), defined the extent of the workings at nearby Avery Island. The spine and its boundary “shear” zone are reflected in the topography of the Jefferson Island landscape, with a natural sub-circular solution lake, Lake Peigneur, created by the dissolving shallow crest of the most recently-active salt spine.

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 13; Autin, 2002). In the 12 hours following the first intersection the underlying mine had flooded, and the lake was completely drained.

Drainage and collapse of the lake began when a Texaco oil rig, 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 14a). 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. 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. In 58 years of mine life, some 23-28 million m3 of salt had been extracted. Prompt reaction to the initial flood wave by mine staff allowed all 50 personnel, who were underground at the time, to escape without anything more than a few minor injuries.


The rapid flush of lake water into the mine, probably augmented by the drainage of natural solution cavities in the caprock below the lake floor, meant landslides and mudflows developed along the perimeter of the sinkhole, and that the lake was enlarged by 28 ha. 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 tug boat and most of the Live Oak botanical gardens. 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 backward into the lake depression through a connecting bayou (Delcambre Canal, aka Carline Bayou). But, since parts of the lake bottom had slumped into the sinkhole during the collapse, the final water level in some sections was higher than before relative to previous land features. It left one former lakefront house aslant under 12 feet of water.

Of course, an anthropogenically induced disaster like this attracted the lawyers like flies to a dead dingo. On 21 November 1980, the day after the disaster, Diamond Crystal Salt filed a suit against Texaco for an unspecified amount of damage. On 25 November, Texaco filed a countersuit against Diamond Crystal. The Live Oak Gardens sued both Diamond Crystal and Texaco. Months later, the State of Louisiana was brought into the suit since the incident occurred on state land. One woman sued Texaco and Wilson Brothers (the drillers) for $1.45 million for injuries (bruised ribs and an injured back) received while escaping from the salt mine. Less than a week before the scheduled trial, an out-of-court settlement was reached between the major players. Due to human error, related to a triangulation mistake when siting the drilling barge, Texaco and Wilson Brothers agreed to pay $32 million to Diamond Crystal and $12.1 million to the Live Oak Gardens.

An ongoing environmental catastrophe that was anticipated by environmental groups at the time of the accident never materialized. The lake quickly returned to its natural freshwater state, and with it the wildlife was largely un-affected. Nine of the barges eventually popped back up like corks (the drilling rigs and tug were never to be seen again). The torrent of water helped dredge Delcambre Canal so that it was two to four feet deeper. And of course, the former 1 metre deep Lake Peigneur was now 400 metres deep in the vicinity of the borehole!

Interestingly, the filling of the mine workings with water drastically slowed the rate of land subsidence atop the mine (Figure 14b). Measurements had been carried out between 1973 and 1983, some 7 years before the accident and continued for 2 years afterward (Thoms and Gerhle, 1994). Slowing reflects the post-accident reduction in the total pressure exerted on the roof of the mine to half its pre-accident levels. Prior to the accident, there was no hydrostatic pressure to alleviate some of the lithostatic pressure exerted by the weight of the overburden and so land subsidence above the mine workings was relatively rapid.

Although this incident is not directly related to any aspect of the salt mining operation and no human lives were lost (although three dogs perished), it clearly illustrates the speed of potential leakage following a breach in a cavern roof in any shallow storage facilities filled with low-density fluids. It also illustrates the usual cause of such disasters – human error in the form of a lack of due diligence, a lack of forward planning and a lack of communication between various private and government authorities. It also illustrates that filling a solution cavity with water slows the rate of subsidence atop a large salt cavity and that waters after the disturbance will return quietly to a state of density stratification.

The incident 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. Horizontal stress on the mineshaft near the level where the Louann Salt contacts the overlying Pleistocene Prairie Complex 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 mine was completed as part of a safe closure plan.

Subsidence over the nearby Avery Island salt mine (operated by Cargill Salt) has been documented since 1986. This is oldest operating salt mine in the United States and has been in operation since the American Civil War, and after the Lake Peigneur disaster the mine underwent a major reconstruction and safety workover. Mine management and landowners did not publicly disclose the technical details of rates of subsidence, but field observations revealed the nature of the subsidence process. Subsidence along the mine edge coincided with a topographic saddle above an anomalous salt zone located inside the mined salt area, ground water had seeped into the mine, and there were a number of soil gas anomalies associated with the mine. Small bead-shaped sinkholes were initially noticed in the area in 1986, then over several years, a broad area of bowl-shaped subsidence and areas of gully erosion formed (Autin, 2002). Reconstruction has now stabilised this situation. Much of the subsidence on Avery Island was a natural process that occurs atop any shallow salt structure. Dating of middens and human artifacts around salt solution induced water-filled depressions atop the dome shows dissolution-induced subsidence is a natural process that extends back well beyond the 3,000 years of human occupation documented on the island.

Compared to the other salt domes of the Five Islands, Cote Blanche Island 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 zones, a salt overhang, and other structural complexities), and (iii) the salt is surrounded by more clayey (impervious) sediments than the other Five Islands, perhaps allowing for lower rates of crossflow and greater hydrologic stability.

Haoud Berkaoui oilfield, Algeria

Located in the Sahara, some 32 km southwest of Ouargla City, the Haoud Berkaoui oilfield is an area of subsidence where numerous exploration and development wells were completed in the 1970s. Of these, the OKN32 and OKN32BIS wells have collapsed into an expanding collapse doline. It surfaced in October 1986 when a crater, some 200 metres across and 75 metres deep, formed (Morisseau, 2000). Today the solution cavity continues to expand and is now some 230 by 600 metres across. Its outward progression is continuing at a rate around 1 metre per year. The collapse is centred on two oil wells drilled in the late 1970s. The problem began in 1978, when the OKN32 oil exploration well was abandoned because of stability problems in Triassic salt at a depth around 650 metres. The target was an Ordovician sandstone at a depth of 2500m. Because of the technical problems associated with significant caverns at the level of the salt, the well was abandoned without casing being set in the salt, probably facilitating the escape of artesian waters (Morisseau, 2000; Bouraoui et al., 2012).


When it reached the 600m level, the well had already passed through 50 metres of anhydrite (220-270m depth) along with interbedded anhydrite clay and dolomite 270m -450 m depth). These are evaporite sediments that, in their undisturbed state, can act as aquicludes or aquitards to any access by unconfined phreatic groundwaters, although at such shallow depths the evaporite beds are likely also to be variably overprinted by active-phreatic dissolution processes. Prior to drilling it was thought that the Senonian halite extended continuously to a depth of 600 m in the well and in turn was underlain by 50 m of anhydrite (600-650m depth). Below the halite-anhydrite is an artesian aquifer (Albian) with a natural hydraulic head that is larger than the surface aquifer head by 2.5 MPa (Morisseau, 2000)

In 1979 a second well, OKN32BIS, was drilled located some 80 metres from the previous well and it successfully obtained its Ordovician target. But in March 1981, the lining of this second well broke, probably because of cavity collapse at a level around 550 metres (once again the regional level of salt) and the well was lost. Five years later, on October 1 1986, a large surface crater formed, centred on these two wells. The initial at-surface diameter was 200 metres and it was 75 metres deep, today it is even larger (Figure 15). Cavern diameter below the stope breakout at that time was estimated to be 300m and water flows to be around 2000 m3/hour.

Since the initial stope breakout, leaching has become progressively less effective and expansion rates have slowed (Morisseau, 2000). This is because cavern growth and water outflow flow are thought to take place preferentially near the centre of the collapse, which is now far from the collapsing cavern walls Dissolving salt may be salinising the crossflowing groundwaters, leading to undocumented, but possible, ongoing degradation of freshwater oases in the region. Continuing expansion is evidenced by the development of fresh centripetal cracks about the expanding collapse margin. Using MT-InSAR analysis, Bouraoui et al. (2012) documented ongoing subsidence near the crater, with an average subsidence of 4 mm per year (between 1992 and 2002). The zone of current zone of subsidence is centred on the OKN32 location and is slowly migrating north east.

As in the USA (see Table 2 and Warren, 2016, Chapter 13 for examples), the loss of these wells, in this case during their active life, emphasizes the need for caution when planning well abandonment in a salt bed, especially when it is highly likely that the salt is acting as a seal, or at least an aquitard, to a regional artesian system. The fact that the first well (OKN32) was lost during drilling argues that a natural breach or cavity was already present in the salt bed and perhaps was already stoping its way to the surface. It is also possible that the inappropriate completion and cementation of casing levels, prior to the well’s abandonment, may have accelerated cavity expansion. In hindsight, the loss of the second well some 5 years later was highly likely as was catastrophic cavity collapse 5 years after that; the OKN32BIS wellhead was situated only 80 metres from OKN32 and was dealing with the same cavern-ridden salt geology.

Summary

Regarding anthropogenically-enhanced salt karst, it is important to note that a salt mass used for storage has never failed catastrophically. Weak points tend to occur wherever “the outside has access to the inside,” so problems tend to be mostly where mine expansion breaches a salt edge (Warren, 2017). Likewise, almost all the problems related to well and cavity failure are more a matter of human error, either by negligence, or a lack of understanding by on-the-ground personnel. There is the same general rule of thumb when it comes to salt cavities and salt mines, and that is, “keep it in the salt!” Most failures and breaches occur when mining or solution leaching operations allow the cavity to contact the edge of the salt. There undersaturated water crossflows can exaggerate any uncontrolled dissolution problems. Often the salt edge is irregular due to natural dissolution and assumptions of flat or gently curved shapes to a salt edge are oversimplifications.

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.

Bouraoui, S., Z. Cakir, R. Bougdal, and M. Meghraoui, 2012, MT-InSAR monitoring of ground deformation around the Haoud Berkaoui sinkhole (SE Algeria): Geophysical Research Abstracts, EGU General Assembly 2012, held 22-27 April, 2012 in Vienna, Austria, v. 14, EGU2012-3344.

Buffet, A., 1998, The collapse of Compagnie des Salins SG4 and SG5 drilling: Proc. S.M.R.I. Fall Meeting, Rome,, p. 79-105.

Gendzwill, D., and N. Martin, 1996, Flooding and loss of the Patience Lake potash mine: CIM Bulletin, v. 89, p. 62-73.

Goodman, W. M., D. B. Plumeau, J. O. Voigt, and D. J. Gnage, 2009, The History of Room and Pillar Salt Mines in New York State,” in S. Zuoliang, ed., Proceedings, 9th International Symposium on Salt, Beijing, China, September 4–6, 2009, v. 2, Gold Wall Press, Beijing, China, p. 1239–1248.

Gowan, S. W., and S. M. Trader, 1999, Mine failure associated with a pressurized brine horizon: Retsof Salt Mine, western New York: Environmental & Engineering Geoscience, v. 6, p. 57-70.

Gowan, S. W., and S. M. Trader, 2003, Mechanism of sinkhole formation in glacial sediments above Retsof Salt Mine, Western New York, in K. S. N. Johnson, J. T. , ed., Evaporite karst and engineering/environmental problems in the United States: Norman, Oklahoma Geological Survey Circular 109, p. 321-336.

Morisseau, J. M., 2000, Uncontrolled leaching of salt layer in an oil field in Algeria: Proc. S.M.R.I. Fall Meeting Technical Session, San Antonio, p. 330-333.

Nieto, A., and R. A. Young, 1998, Retsof Salt Mine Collapse and Aquifer Dewatering, Genesee Valley , Livingston County , NY, in J. Borchers, ed., Poland Symposium Volume: Land Subsidence, Spec. Pub. 8, Assoc. Engineering Geologists, p. 309-325.

Payment, K. A., 2000, Loss of the Retsof salt mine: legal analysis of liability issues, in R. M. Geertmann, ed., Proc. 8th World Salt Symp., Salt 2000, The Hague, v. 1: Amsterdam, Elsevier, p. 399-404.

Tepper, D. H., W. H. Kappel, T. S. Miller, and J. H. WilliaMS, 1997, Hydrogeologic effects of flooding in the partially collapsed Retsof salt mine, Livingston County, New York: US Geol. Survey Open File Report, v. 97-47, p. 36-37.

Thoms, R. L., 2000, Subsidence and sinkhole development over salt caverns: An introduction to the technology of solution mining; Spring 2000 Technical Class, p. 127-141.

Thoms, R. L., and R. M. Gehle, 1994, Analysis of a Solidified Waste Disposal Cavern in Gulf Coast Salt Dome: SMRI Fall Mtg. (1994) 637.

Thoms, R. L., and R. M. Gehle, 2000b, Winnfield mine flooding and collapse event of 1965: Proc. S.M.R.I. Fall Meeting Technical Session, San Antonio, p. 262-274.

Van Sambeek, L. L., S. W. Gowan, and K. A. Payment, 2000, Loss of the Retsof Mine: Engineering Analysis: Proceedings, 8th World Salt Symposium, The Hague, The Netherlands, May 7–11, R. M. Geertman (ed.), Elsevier Science Publishers B.V., Amsterdam, The Netherlands, pp. 411–416.

Von Tryller, H., 2002, The Cavern Field No. 11 in Ocnele Mari - History, Present and Future: Solution Mining Research Institute Proceedings, Spring Meeting, 28 April 1 May, 2002, Banff, Canada, p. 10 pp.

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

Warren, J. K., 2017, Salt usually seals, but sometimes leaks: Implications for mine and cavern stabilities in the short and long term: Earth-Science Reviews, v. 165, p. 302-341.

Yager, R. M., 2013, Environmental Consequences of the Retsof Salt Mine Roof Collapse, US Geological survey Open File Report 2013–1174, 10 p.

Yager, R. M., T. S. Miller, and W. M. Kappel, 2001, Simulated effects of 1994 salt-mine collapse on ground-water flow and land subsidence in a glacial aquifer system, Livingston County, New York: US Geological Survey Professional Paper, p. 1-80.

Yager, R. M., P. E. Misut, C. D. Langevin, and D. L. Parkhurst, 2009, Brine Migration from a Flooded Salt Mine in the Genesee Valley, Livingston County, New York: Geochemical Modeling and Simulation of Variable-Density Flow, USGS Professional Paper 1767, 59 p.

Young, R. A., and G. S. Burr, 2006, Middle Wisconsin glaciation in the Genesee Valley, NY: A stratigraphic record contemporaneous with Heinrich Event, H4: Geomorphology, v. 75, p. 226-247.

Zamfirescu, F., M. Mocuta, T. Constantinecu, E. Medves, and A. Danchiv, 2003, The main causes of a geomechanical accident of brine caverns at field II of Ocnele Mari - Romania: RMZ - Materials and Geoenvironment, v. 50, p. 431-434.

Zuber, A., J. Grabczak, and A. Garlicki, 2000, Catastrophic and dangerous inflows to salt mines in Poland as related to the origin of water determined by isotope methods: Environmental Geology, v. 39, p. 299-311.



 

 

Salt Dissolution (3 of 5): Natural Geohazards

John Warren - Tuesday, October 31, 2017


Introduction

Surface constructions and other anthropogenic activities atop or within evaporite karst terranes is more problematic than in subcopping carbonate terranes due to inherently higher rates of dissolution and stoping (Yilmaz et al., 2011; Cooper and Gutiérrez, 2013; Gutiérrez et al., 2014). Overburden collapse into nearsurface gypsum caves can create stoping chimneys, which break out at the surface as steep-sided dolines, often surrounded by broader subsidence hollows. Such swallow-holes, up to 20 m deep and 40 m wide, continue to appear suddenly and naturally in gypsum areas throughout the world.

Unlike the relatively slow formation of limestone karst, gypsum/halite karst develops on a human/engineering time-scale and can be enhanced by human activities (Warren, 2016, 2017). For example, in 2006, the Nanjing Gypsum mine in China broke into a phreatic cavity in a region of gypsum karst, driving complete flooding of the mine in some three days. Associated groundwater drainage caused a sharp drop in the local piezometric level of up to 90 m in a well in nearby Huashu village. Resultant ground subsidence severely damaged nearby roads and buildings (Wang et al., 2008). In Ukraine, dewatering of gypsum karst to facilitate sulphur mining substantially increased the rate of gypsum dissolution and favoured the expansion of sinkholes within an area affected by the associated cones of water-table depression (Sprynskyy et al., 2009). Natural evaporite karst enhanced by intrastructure focusing of drainage creates the various scales of problem across the Gypsum Plain of West Texas and New Mexico (Stafford et al., 2017).

Although halite is even more susceptible to rapid dissolution than gypsum, it typically is not a major urban engineering problem; large numbers of people simply do not like to live in a climate that allows halite to make it to the surface. However, in the Dead Sea region, the ongoing lowering of the water level encouraged karstic collapse in newly exposed mudflats and has damaged roads and other man-made structures (Frumkin et al. 2011; Shviro et al., 2017). Catastrophic doline collapse atop poorly managed halite/potash mines and solution brine fields is an additional anthropogenically-induced or enhanced geohazard in developed regions is discussed in detail in Warren, 2016 (Figure 1).


Gypsum karst is a documented natural hazard in many parts of Europe (Figure 2), and similar areas of shallow subcropping gypsum are common in much of the rest of the world (Table 1). For example, areas surrounding the city of Zaragoza in northern Spain are affected, as is the town of Calatayud (Gutiérrez and Cooper, 2002; Gutiérrez, 2014). Gypsum dissolution is responsible for subsidence and collapse in many urban areas around northern Paris, France (Toulemont, 1984), in urban areas in and around Stuttgart and other towns peripheral to the Harz Mountains in Germany (Garleff et al., 1997), in Pasvalys and Birzai in Lithuania (Paukstys et al., 1999), in the Muttenz-Pratteln area in northwestern Switzerland (Zechner et al., 2011), in the Perm area of Russia (Trzhtsinsky, 2002), in the Sivaz region of Turkey (Karacan and Yilmez, 1997), in the region centred on the city of Mosul in northern Iraq (Jassim et al., 1997) and in a number of areas of rapid urban development in eastern Saudi Arabia (Amin and Bankher, 1997a, b). Large subsidence depressions caused by gypsum dissolution in China have opened up in the Taiyuan and Yangquan regions of Shanxi Coalfield and the adjacent Hebei Coalfield.


Variation in the watertable level, induced by groundwater pumping or uncontrolled brine extraction, can be an anthropogenic trigger for dolines surfacing. As the watertable declines it causes a loss of buoyant support to the ground, it also increases the flow gradient and water velocity, which facilitates higher rates of crossflow and deeper aquifer recharge in subsequent floods and so reduces the geomechanical strength of the cover and washes away roof span support (Figures 1, 3). Dolines can also be associated with groundwater quality issues. Collapse dolines or sinkholes are frequently used as areas or sumps for uncontrolled dumping industrial and domestic waste. Because of the direct connection between them and the regional aquifer, uncontrolled dumping can cause rapid dispersion of chemical and bacterial pollutants in the groundwater. In the case of Riyadh region Saudi Arabia, a lake of near-raw sewage has appeared in Hit Dahl (cave) and is likely related to the increased utilisation of desalinated water for sanitation and agriculture (Warren, 2016). In the Birzai region of Lithuania numerous sinkholes developed in Devonian gypsum subcrop are in direct connection across the regional hydrology. Accordingly, the amount of agricultural fertilizer use is limited to help protect groundwater quality.

One of the problems associated with rapid surfacing of evaporite collapse features is that any assignment of sinkhole cause will typically lead to an assignment of blame, particulary when anthropogenic infrastructure has been damaged or destroyed by the collapse, or lives may have been lost. Areas of natural evaporite karst are typically areas of relatively shallow evaporites. Shallow evaporites make such regions suitable for extraction via conventional or solution mining. When a collapse does occur in a mined area, one group (generally the miners) has a vested interest in arguing for natural collapse, the others, generally the lawyers and their litigants, will argue for an anthropogenic cause. The reality is usually a combination of natural process enhanced to varying degrees by human endeavours. In the examples in this section, much of the driving process for the collapse is natural, while the cause of any unexpected karst-related disaster is typically geological ignorance combined with political/community intransigence. See Chapter 13 for a further discussion of karst and stope examples that include collapses and explosions where the anthropogenic drivers can dominate.

Problems in the Ripon area, Yorkshire, UK

The town of Ripon, North Yorkshire, and town’s surrounds experiences the worst ongoing gypsum-karst related subsidence in England (Figures 3, 4; Cooper and Waltham, 1999). Some 43 events of subsidence or collapse in the caprock over the Ripon gypsum have occurred over the last 160 years, within an area of 7 km2 (Figures 4). This gives a mean rate of one new sinkhole every 26 years in each square kilometre. Worldwide, the highest documented event rate occurs in Ukraine, in an area of thin and weak clay caprocks above interstratal gypsum karst, where new sinkholes appear at a rate of 0.01 to 3.0 per year per km2 (Waltham et al., 2005). In the Ripon area, numerous sags and small collapses also typify surrounding farmlands. Subsidence features are typically 10-30m in diameter, reach up to 20m in depth and can appear at the surface in a matter of hours to days (Figure 3). To the east of the town, one collapse sinkhole in the Sherwood Sandstone is 80 m in diameter and 30 m deep, perhaps reflecting the stronger roof beam capacity of the Sherwood Sandstone.

When a chimney breaks through, the associated surface collapse is very rapid (Figure 3 b-e). For example, one such subsidence crater, which opened up in front of a house on Ure Bank terrace on 23rd and 24th April, 1997, is documented by Cooper (1998.) as follows (Figure 3b).

“...The hole grew in size and migrated towards the house, to measure 10m in diameter and 5.5m deep by the end of Thursday. Four garages have been destroyed by the subsidence. This collapse was the largest of one of a series that have affected this site for more than 30 years; an earlier collapse had demolished two garages on the same site, and a 1856 Ordnance Survey map shows a pond on the same site. The hole is cylindrical but will ultimately fail to become a larger, but conical, depression. As it does so, it may cause collapse of the house, which is already damaged, and the adjacent road. The house and several nearby properties have been evacuated and the nearby road has been closed. The gas and other services, which run close to the hole, have also been disconnected in case of further collapse.”

Cooper (1998) found the sites of most severe subsidence in the Ripon area (including the house at Ure Terrace and in the vicinity of Magdelen's Road) are located at the sides of the buried Ure Valley, an area where the significant volumes of water seeps from the gypsum karst levels into the river gravels (Figure 4). In 1999 the Ure Terrace sinkhole was filled using a long conveyor belt that was cantilevered over the hole so that no trucked needed to back up close to the sinkhole opening. The hole was surcharged to a height of 0.5m. The hole remains unstable, but the collapse of the fill is monitored to document fill performance and the fill is periodically topped up. After the sinkhole was filled, the road adjacent to the sinkhole was re-opened and the site of the sinkhole fenced. The severely damaged Field View house remains standing next to the sinkhole. The nearby Victorian Ure lodge was not directly damaged by the 1997 sinkhole, but its western corner fell within the council-designated damage zone, and was left unoccupied. It fell into disrepair and was subsequently demolished (Figure 3b). A similar fate befell houses damaged by the surfacing of collapse sinkholes in and around Magdelen's Road, which is located a few hundred metres from Ure Terrace (Figure 3c-e). Shallow subcropping Zechstein gypsum (rehydrated anhydrite) occurs in two subcropping bedded units in this area, one is in the Permian Edlington and the other is in the Roxby Formation (Figure 4b). Together they form a subcrop belt about a kilometre wide, bound to the west by the base of the lowest gypsum unit (at the bottom of the Edlington Formation) and to the east by a downdip transition from gypsum to anhydrite in the upper gypsum-bearing unit of the Roxby Formation. The spatial distribution of subsidence features within this belt relates to joint azimuths in the Permian bedrock, with gypsum maze caves and subsidence patterns following the joint trends (Cooper, 1986). Most of the subcropping gypsum is alabastrine in the area around Ripon, while farther to the east, where the unit is thicker and deeper, the calcium sulphate phase is still anhydrite.

Fluctuations in the watertable level tied to heavy rain or long drought are thought to be the most common triggering mechanism for subsidence transitioning to sinkhole collapse. Many of the more catastrophic collapses occur after river flooding and periods of prolonged rain, which tend to wash away cavern roof span support. Subsidence is also aggravated by groundwater pumping; first, it lowers the watertable and second, it induces considerable crossflow of water in enlarged joints in the gypsum. When recharged by a later flood, the replacement water is undersaturated with respect to gypsum.


Thomson et al. (1996) recognised four hydrogeological flow units driving karst collapse in the Ripon area (Figure 4):

1) Quaternary gravels in the buried valley of the proto-River Ure

2) Sherwood Sandstone Group

3) Magnesian Limestone of the Brotherton Fm. and the overlying/adjacent gypsum of the Roxby Fm.

4) Magnesian limestone of the Cadeby Fm. plus the overlying/adjacent gypsum of the Edlington Fm.

Local hydrological base level within this stratigraphy is controlled by the River Ure, especially where the buried Pleistocene valley (proto-Ure) is filled by permeable sands and gravels, as these unconsolidated sediments, when located atop a breached roof beam, are susceptible to catastrophic stoping to base level (Figure 4). In the area around Ripon the palaeovalley cuts down more than 30 m, reaching levels well into the Cadeby Formation, so providing the seepage connections or pathways between waters in all four units wherever they intersect the palaeovalley. There is considerable groundwater outflow along this route with artesian sulphate-rich springs issuing from Permian strata in contact with Quaternary gravels of the buried valley (Cooper, 1986, 1995, 1998).

The potentiometric head comes from precipitation falling on the high ground of the Cadeby formation to the west and the Sherwood Sandstone to the east. Groundwater becomes largely confined beneath glacial till as it seeps toward the Ure Valley depression, but ultimately finds an exit into the modern river via the deeply incised sand and gravel-filled palaeovalley of the proto-Ure. Waters recharging the Ure depression pass through and enlarge joints and caverns in the gypsum units of the Edlington and Roxby Formations, so the highest density of subsidence features are found atop the sides of the palaeovalley. This region has the greatest volume of artesian discharge from aquifers immediately beneath the dissolving gypsum bed. Although created as an active karst valley, the apparent density of subsidence hollows is lower on the present Ure River floodplain than the surrounding lands as floodplain depressions are constantly filled by overbank sediments (Figure 4b).

Cooper (1998) defined 16 sinkhole variations in the gypsum subsidence belt at Ripon, all are types of entrenched, subjacent and mantled karst. Changes in karst style are caused by; the type of gypsum, the nature and thickness of the overlying deposits, presence or absence of consolidated layers overlying the gypsum and the size of voids/caverns within the gypsum.

To the west of Ripon, the gypsum of the Edlington Formation lies directly beneath glacial drift. These unconsolidated drift deposits and the loose residual marl atop the dissolving gypsum gradually subside into a pinnacle or suffusion (mantled) karst. But between Ripon town and the River Ure, the limestone of the Brotherton Formation overlies the Edlington Formation. There the karst develops as large open caverns beneath strong roof spans (entrenched karst). Ultimate collapse of the roof span creates rapid upward-stoping caverns in loosely consolidated sediment. Stopes break though to the surface as steep-sided collapse dolines or chimneys with sometimes catastrophic results. A similar entrenched situation is found east of the Ure River but there karstified gypsum units of both the Edlington and the Roxby formations are involved.


There are also thick beds of gypsum in the Permian Zechstein sequence that forms the bedrock in the Darlington area. In this area, subsidence features attributed to gypsum dissolution are typically broad shallow depressions up to 100 m in diameter, and the ponds, known as Hell Kettles, are the only recognized examples of steep-sided subsidence hollows around Darlington (Figure 5). Historical records suggest that one of the ponds formed in dramatic fashion in AD 1179 (Cooper 1995). The southern pond appears to be the most likely one to have formed at that time because it is many metres deep and is fed from below by calcareous spring water that is rich in both carbonate and sulphate. The 2D profiles have revealed evidence of foundering in the limestone of the Seaham Formation at depths of c. 50 m (Figure 5; Sargent and Goulty, 2009). The foundering is interpreted to have resulted from dissolution of gypsum in the Hartlepool Anhydrite Formation at ≈ 70 m depth. The reflection images of the gypsum itself are discontinuous, suggesting that its top surface has karstic topography. The 3D survey also acquired and interpreted by Sargent and Goulty (2009) reveals subcircular hollows in the Seaham Formation up to 20 m across, which are again attributed to foundering caused by gypsum dissolution.


Problems with Miocene gypsum, Spain

Karstification has led to problems in areas of subcropping Miocene gypsum in the Ebro and Calatayud basins, northern Spain (Figure 6). Cliff sections and road cuts indicate the widespread nature of karstification in the gypsum outcrops and subcrops in Spain (Figure 7b) Areas affected are defined by subsidence or collapse in Quaternary alluvial overburden and include; urban areas, communication routes, roads, railways, irrigation channels and agricultural fields (Figure 7a; Soriano and Simon, 1995; Elorza and Santolalla, 1998; Guerrero et al., 2013; Gutiérrez et al., 2014). In the region there can be a reciprocal interaction between anthropic activities and sinkhole generation, whereby the ground disturbance engendered by human activity accelerates, enlarges and triggers the creation of new sinkholes. Subsidence is particularly harmful to linear constructions and buildings and numerous roads, motorways and railways have been damaged (Figure 7a, b). Catastrophic collapse and rapid karst chimneying into roads and buildings can have potentially fatal consequences. For example, several buildings have been damaged around the towns of Casetas and Utebo. In the Portazgo industrial estate some factories had to be pulled down due to collapse-induced instability (Castañeda et al., 2009). A nearby gas explosion was attributed to the breakage of a gas pipe caused by subsidence. The local water supply is also disrupted by subsidence and pipe breakage so that 20,000 inhabitants periodically lose their water supply. The most striking example of subsidence affecting development comes from the village of Puilatos, in the Gallego Valley. In the 1970's this town was severely damaged by subsidence and abandoned before it could be occupied (Cooper 1996).


Collapse affects irrigation channels in the countryside with substantial economic losses (Elorza and Santolalla, 1998). In 1996 a doline collapse surfaced and cut the important Canal Imperial at Gallur village. New dolines often form near unlined irrigation canals. The ongoing supply of fresh irrigation waters to field crops can also encourage sinkhole generation in the fields. Though not directly visible, natural sinkholes also form in the submerged beds of river channels cutting regions of subcropping gypsum.

On December 19th, 1971, a bus fell from a bridge into the Ebro River at Zaragoza, near where the ‘San Lazaro well’ (a submerged gypsum sinkhole) is located (Figure 8a). Ten people lost their lives in this accident , while the remainder of the passengers were rescued, after being stranded on the bus roof in the flowing river for some hours (Figure 8b). After survivors were rescued, river waters washed the bus from the foot of the bridge supports into the nearby 'San Lazaro well (collapse sinkhole) in the water-covered floor of the river. Nine of the ten bodies in the bus were never found, although the bus was later recovered from the sinkhole. Locals suggested that bodies were carried deeper into the various interconnect phreatic sinkhole caverns fed by this losing stream.


Karstification in the Zaragoza region is characterised by the preferential intrastatal dissolution of glauberite bed, which are more soluble than the gypsum interbeds, this leads to collapse and rotation of gypsum blocks and river capture (Guerrero et al., 2013).

Sometimes even well-intentioned attempts to remediate culturally significant buildings under threat of evaporite karst collapse can exacerbate collapse problems. Gutiérrez and Cooper (2002) cite examples from the city of Calatayud, Spain. Subsidence-induced differential loading across doline edges drives the tilting of the 25-metre high tower (mudéjar) of the San Pedro de Los Francos church, which leans towards and overhangs the street by about 1.5 metres. (Figure 9) In places, the brickwork of the church indents the pre-existing tower fabric, which probably dates from the 11th Century or the beginning of the 12th Century. This indentation and the non-alignment of the church and the tower walls indicates that most of the tower tilting occurred prior to the construction of the church. In 1840, the upper 5m of the tower was removed and the lower part buttressed for the safety of the Royal family, who visited the town and stayed in the palace opposite. On 3rd June 1931, San Pedro de Los Francos church was declared a “Monument of Historical and Artistic value.” Due to its ruinous condition, the church was closed to worship in 1979. Micropiling to improve the foundation was started in 1994, but this corrective measure was interrupted when only half of the building was underpinned. Very rapid differential settlement of the building took place in the following year, causing extensive damage and aggravating the subsidence problem.


Colegiata de Santa María la Mayor was constructed between the 13th and 18th centuries, it has an outstanding Mudéjar (a 72 m high tower) and numerous Renaissance features; it is considered the foremost monument in the city of Cataluyud. As with the San Pedro de los Francos Church, recent micropiling work, applied to only one part of the cloister, has been followed by alarming differential movements that have drastically accelerated the deterioration of the building. Large blocks have fallen from the vault of the “Capitular Hall” and cracks up to 150 mm wide have opened in the brickwork of the back (NW) elevation, which has now been shored up for safety. The dated plaster tell-tales placed in these cracks to monitor the displacement demonstrate the high speed of the deformation produced by subsidence in recent years. On the afternoon of 10 September 1996, the fracture of a water supply pipe flooded the cloisters and the church with 100 mm of muddy water. Ten years earlier a similar breakage and flood had occurred. These breaks in the water pipes are most likely related to karst-induced subsidence. Once they occur, the massive input of water to the subsurface may trigger further destruction via enhanced dissolution, piping and hydrocollapse (Gutiérrez and Cooper, 2002).


Gypsum karst in Mosul, Iraq

A similar quandary of multiple areas of structural damage from gypsum-induced subsidence affects large parts or the historic section of the city of Mosul in northern Iraq (Jassim et al., 1997). The main part of its old quarter is over a century old and some buildings are a few hundred years old. Mosul lies on the northeastern flank of the Abu Saif anticline and near to its northern plunge (Figure 10a). It was built on the western bank of the Tigris River on a dip slope of Middle Miocene Fatha limestone that is directly underlain by bedded gypsum and green marl (equivalent to Lower Fars Formation). Houses in the old city were built on what seemed to be at the time a very sound rock foundation.

Water distribution in the city was done on mule back in the early part of last century and the estimated water consumption did not exceed 10 litres per person per day (Jassim et al., 1997). Discharge from households was partly to surface drainage and partly to shallow and small septic tanks. The modern piped system of water distribution did not start until the 1940s, resulting in a sudden increase in water consumption (presently around 200 litres per person per day) and it was not associated with a complementary sewer system. Increased water consumption meant larger and deeper septic tanks were dug at the perimeter of buildings (which never seemed to fill) resulting in a dramatic increase in water percolating downwards, water that was also more corrosive than previously due to the increased use of detergents and chlorination. This water passes through the permeable and fractured limestone to the underlying gypsum. On its way through the limestone it enlarges and creates new dissolution cavities, but eventually finds its way into the older gypsum karst maze, which is then further widened as water drains back into the Tigris (Figure 10b). Caverns in the gypsum enlarge until the roof span collapses. Since the 1970s more and more buildings in the old city have fractured and many are subject to sudden collapse. The problem is further intensified due to the expansion of the city in the up-dip direction (west and southwest) including the construction of industrial, water-dependent centres with integrated drainage. Water seeping/draining from these newly developed up-dip areas eventually passes under the old city before discharging in the Tigris river. The process was slightly arrested in the 1980s by the completion of a drainage system for the city, but the degradation of the old city continues.

Coping: man-made structures atop salts

The towns of Ripon in the UK and Pasvales and Birzai in Lithuania house some 45,000 people, who currently live under the ongoing threat of catastrophic subsidence, caused by natural gypsum dissolution (Paukstys et al., 1999). Special measures for construction of houses, roads, bridges and railways are needed in these areas and should include: incorporating several layers of high tensile heavy duty reinforced plastic mesh geotextile into road embankments and car parks; using sacrificial supports on bridges so that the loss of support of any one upright will not cause the deck to collapse; extending the foundations of bridge piers laterally to an amount that could span the normal size of collapses; and using ground monitoring systems to predict areas of imminent collapse (Cooper 1995, 1998).


Dams to store urban water supplies are costly structures and failure can lead to disaster, large scale mortality and financial liability (for example, Cooper and Gutiérrez, 2013). For example, at two and a half minutes before midnight on March 12, 1928, the St. Francis Dam (California) failed catastrophically and the resulting flood killed more than 400 people (Figure 11). The collapse of the St. Francis Dam is considered to be one of the worst American civil engineering disasters of the 20th century and remains the second-greatest loss of life in California’s history, after the 1906 San Francisco earthquake and fire. The collapse was partly attributed to dissolution of gypsum veins beneath the dam foundations. The Quail Creek Dam, Utah, constructed in 1984 failed in 1989, the underlying cause being an unappreciated existence of, and consequent enlargement of, cavities in the gypsum strata beneath its foundations.

Unexpected water leakage from reservoirs, via ponors, sinkholes and karst conduits, leads to costly inefficiency, or even project abandonment. Unnaturally high hydraulic gradients, induced by newly impounded water, may flush out of the sediment that previously blocked karst conduits. It can also produce rapid dissolutional enlargement of discontinuities, which can quickly reach break-through dimensions with turbulent flow. These processes may significantly increase the hydraulic permeability in the region of the dam foundation, on an engineering time scale.

Accordingly, numerous dams in regions of the USA underlain by shallow evaporites either have gypsum karst problems, or have encountered gypsum-related difficulties during construction (Johnson, 2008). Examples include; the San Fernando, Dry Canyon, Buena Vista, Olive Hills and Castaic dams in California; the Hondo, Macmillan and Avalon dams in New Mexico; Sandford Dam in Texas; Red Rock Dam in Iowa; Fontanelle Dam in Oklahoma; Horsetooth Dam and Carter Dam in Colorado and the Moses Saunders Tower Dam in New York State. Up to 13,000 tonnes of mainly gypsum and anhydrite were dissolved from beneath a dam in Iraq in only six months causing concerns about the dam stability (Figure 13). In China, leaking dams and reservoirs on gypsum include the Huoshipo Dam and others in the same area. The Bratsk Dam in eastern Siberia is leaking, and in Tajikistan the dam for the Nizhne-Kafirnigansk hydroelectric scheme was designed to cope with active gypsum dissolution occurring below the grout curtain. Gypsum karst in the foundation trenches of the Casa de Piedra Dam, Argentina and El Isiro Dam in Venezuela, caused difficult construction conditions and required design modifications.


Another illustration of the problems associated with water retaining structures and the ineptitude, or lack of oversight, by some city planners comes from the town of Spearfish, South Dakota (Davis and Rahn, 1997 ). As discussed earlier in this chapter, the Triassic Spearfish Formation contains numerous gypsum beds in which evaporite-focused karst landforms are widely documented across its extent in the Black Hills of South Dakota (Figure 12). The evaporite karst in the Spearfish Fm. has caused severe engineering problems for foundations and water retention facilities, including wastewater stabilization sites. One dramatic example of problems in water retention atop gypsum karst comes from the construction in the 1970s of now-abandoned sewage lagoons for the City of Spearfish.

Despite warnings from local ranchers, the Spearfish sewage lagoons were built in 1972 by city authorities on alluvium atop thick gypsum layers of Spearfish Formation. Ironically, at one point during lagoon construction, a scraper became stuck in a sinkhole and required four bulldozers to pull it out. Once filled with sewage, within a year the lagoons started leaking badly; the southern lagoon was abandoned after four years because of ongoing uncontrollable leaks, and the northern lagoon did not completely drain, but could not provide adequate retention time for effective sewage treatment. Attempts at repairs, including a bentonite liner, were ineffective, and poorly treated sewage discharged beneath the lagoon’s berm into a nearby surface drainage. The lagoons were abandoned completely in 1980. This was after a US $27-million lawsuit was filled in 1979 by ranchers whose land and homes were affected by leaking wastewater. A mechanical wastewater treatment plant was constructed nearby on an outcrop of the non-evaporitic Sundance Formation. The engineering firm that designed the facility without completing a knowledgeable geological site survey was reorganised following the lawsuit.

Likewise, the development of Chamshir Dam atop Gascharan Formation outcrop and subcrop in Iran is likely to create ongoing infrastructure cost and water storage problems (Torabi-Kaveh et al., 2012). The site is located in southwest of Iran, on Zuhreh River, 20 km southeast of Gachsaran city. The area is partially covered by evaporite formations of the Fars Group, especially the Gachsaran Formation. The dam axis is located on limestone beds of Mishan Formation, but nearly two-thirds of the dam reservoir is in direct contact with the evaporitic Gachsaran Formation. Strata in the vicinity of the reservoir and dam site have been brecciated and intersected by several faults, such as the Dezh Soleyman thrust and the Chamshir fault zone, which all act in concert to create karst entryways, including local zones of suffusion karst. A wide variety of karstic features typify the region surrounding the dam site and include; karrens, dissolution dolines, karstic springs and cavities. These karst features will compromise the ability of Chamshir Dam to store water, and possibly even cause breaching of the dam, via solution channels and cavities which could allow significant water flow downstream of the dam reservoir. As possible and likely partial short term solutions, Torabi-Kaveh et al. (2012) recommend the construction of a cutoff wall and/or a clay blanket floor to the reservoir

Difficulties in building hydraulic structures on soluble rocks are many, and dealing with them greatly increases project and maintenance costs. Gypsum dissolution at the Hessigheim Dam on the River Neckar in Germany has caused settlement problems in sinkholes nearby. Site investigation showed cavities up to several meters high and remedial grouting from 1986 to 1994 used 10,600 tonnes of cement. The expected life of the dam is only 30-40 years, with continuing grouting required to keep it serviceable.

Grouting costs in zones of evaporite karst can be very high and may approach 15 or 20% of the dam cost, currently reaching US$ 100 million in some cases. In karstified limestones grouting is difficult, yet in gypsum it is even more difficult due to the rapid dissolution rate of the gypsum. Karst expansion in limestone occurs on the scale of hundreds of years, in gypsum it can be on the order of a decade or less. Grouting may also alter the underground flow routes, so translating and focusing the problems to other nearby areas. In the Perm area of Russia, gypsum karst beneath the Karm hydroelectric power station dam has perhaps been successfully grouted, a least in the short term, using an oxaloaluminosilicate gel that hardens the grout, but also coats the gypsum, so slowing its dissolution. The Mont Cenis Dam, in the French Alps, is not itself affected by the dissolution of gypsum. However, the reservoir storage zone is leaking and photogrammetric study of the reservoir slopes showed ongoing doline activity over gypsum and subsidence in the adjacent land.


Probably the worst example tied to and evaporite karst hazard is the significant dam disaster waiting to happen that is the Mosul Dam in Iraq (Figure 13; Kelley et al., 2007; Sissakian and Knutsson, 2014; Milillo et al., 2016). It is ranked as the fourth largest dam in the Middle East, as measured by reserve capacity, capturing snowmelt from Turkey, some 70 miles (110 km) north. Built under the despotic regime of Saddam Hussein, completed in 1984 the Mosul Dam (formerly known as Saddam Dam) is located on the Tigris river, some 50 km NW of Mosul.

The design of the dam was done by a consortium of European consultants (Sissakian and Knutsson, 2014), namely, Swiss Consultants group, comprising: Motor Columbus; Electrowatt; Suiselectra; Societe Generale pour l’Industrie. The construction was carried out by a German-Italian consortium of international contractors, GIMOD joint venture, comprising: Hochtief; Impregilo; Zublin; Tropp; Italstrade; Cogefar. The consultants for project design and construction supervision comprised a joint venture of the above listed Swiss Consultants Group and Energo-Projekt of Yugoslavia, known as MODACON.

As originally constructed the dam is 113 m in height, 3.4 km in length, 10 m wide in its crest and has a storage capacity of 11.1 billion cubic meters (Figure 13b). It is an earth fill dam, constructed on evaporitic bedrock atop a karstified high created by an evaporite cored anticline in the Fat’ha Formation, which consists of gypsum beds alternating with marl and limestone (Figure 13a, 14). To the south, this is same formation with the same evaporite cored anticlinal association that created all the stability problems in the city of Mosul (Figures 10). The inappropriate nature of the Fat’ha Formation as a foundation for any significant engineering structure had been known for more than a half a century. Then again, absolute rulers do not need to heed scientific advice or knowledge. Or perhaps he didn’t get it from a well-paid group of Swiss-based engineering consultants. As Kelley et al. (2007) put it so succinctly....“The site was chosen for reasons other than geologic or engineering merit.”

The likely catastrophic failure of Mosul Dam will drive the following scenario (Sissakian and Knutsson, 2014); “... (dam) failure would produce a flood wave crest about 20 m deep in the City of Mosul. It is estimated that the leading edge of the failure flood wave would arrive in Mosul about 3 hours after failure of the dam, and the crest of the flood wave would arrive in Mosul about 9 hours after failure of the dam. The total population of the City of Mosul is about 3 million, and it is estimated that about 2 million people are in locations within the city that would be inundated by a 20 m deep flood wave. The City of Baghdad is located about 350 km downstream of Mosul Dam, and the dam failure flood wave will arrive after 72 hours in Baghdad and (by then) would be about 4 m deep.”



The heavily karsted Fat’ha Formation is up to 352 m thick at the dam and has an upper and lower member. The lower member is dominated by carbonate in its lower part (locally called “chalky series”) and is in turn underlain by an anhydrite bed known as the GBo. Gypsum beds typify its upper part,and the evaporite interval is capped by a limestone marker bed. The upper member, crops out as green and red claystone with gypsum relicts, around the Butmah Anticline. Thickness of individual gypsum beds below the dam foundations can attain 18 m; these upper member units are intensely karstified, even in foundation rocks, with cavities meters across documented during construction of the dam (Figure 14). Gypsum breccia layers are widespread within the Fatha Formation and have proven to be the most problematic rocks in the dam’s foundation zone. The main breccia body contains fragments or clasts of limestone, dolomite, or larger pieces of insoluble rocks of collapsed material. The upper portion of the accumulation grades upward from rubble to crackle mosaic breccia and then a virtually unaffected competent overburden. Breccia also may form without the intermediate step of an open cavity, by partial dissolution and direct formation of rubble. As groundwater moves through the rubble, soluble minerals are carried away, leaving insoluble residues of chert fragments, quartz grains, silt, and clay in a mineral matrix. These processes result in geologic layers with lateral and vertical heterogeneity on scales of micro-meters to meters.

High permeability zones in actively karsting gypsum regions can form rapidly, days to weeks, and quickly become transtratal. So predicting or controlling breakout zones via grouting and infill can be problematic (Kelley et al., 2007; Sissakian and Knutsson, 2014). For example, four sinkholes formed between 1992 and 1998 approximately 800 m downstream in the maintenance area of the dam (Figure 13a). The sinkholes appeared in a linear arrangement, approximately parallel to the dam axis. Another large sinkhole developed in February 2003, east of the emergency spillway when the pool elevation was at 325 m. The Mosul Dam staff filled the sinkhole the next day, with 1200 m3 of soil. Another sinkhole developed in July 2005 to the east of the saddle dam. Six borings were completed around the sinkhole and indicated that the sinkhole developed beneath overburden deposits and within layers of the Upper Marl Series. Another cause for concern at Mosul Dam in recent years is a potential slide area reported upstream of the dam on the west bank. The slide is most likely related to the movement of beds of the Chalky Series over the underlying GBo (anhydritic) layer.

To “cope” with ongoing active karst growth beneath and around the Mosul Dam, a continuous grouting programme was planned, even during dam construction, and continues today, on a six days per week basis. It pumps tens of thousands of tons of concrete into expanding karst features each year (Sissakian and Knutsson, 2014; Milillo et al., 2016). The dam was completed in June 1984, with a postulated operational life of 80 years. Due to insufficient grouting and sealing in and below the dam foundation, numerous karst features, as noted above, continue to enlarge in size and quantity, so causing serious problems for the ongoing stability of the dam. The increase in hydraulic gradient created by a wall of water behind the dam has accelerated the rate of karstification in the past 40 years.

Since the late 1980s, the status of the dam and its projected collapse sometime within the next few decades has created ongoing nervousness for the people of Mosul city and near surroundings. All reports on the dam since the mid 1980s have underlined the need for ongoing grouting and monitoring and effective planning of the broadcasting of a situation where collapse is imminent. For “Saddam’s dam” the question is not if, but when, the dam will collapse. To alleviate the effects of the dam collapse, Iraqi authorities have started to build another “Badush Dam” south of Mosul Dam so that it can stop or reduce the effects of the first flood wave. However this new dam has a projected cost in excess of US$ ten billion and so lies beyond the financial reach of the current Iraqi government. Problems related to the dam increased with the takeover of the region by the forces of ISIL.

Today, the Mosul dam is subsiding at a linear rate of ~15 mm/year compared to 12.5 mm/year subsidence rate in 2004–2010 (Milillo et al., 2016). Increased subsidence restarted at the end of 2013 after re-grouting operations slowed and at times stopped. The causes of the observed linear subsidence process of the dam wall can be found in the human activities that have promoted the evaporite–subsidence development, primarily in gypsum deposits and may enable, in case of continuous regrouting stop, unsaturated water to flow through or against evaporites deposits, allowing the development of small to large dissolution cavities.

Large vertical movements that typified the dam wall have resulted from the dissolution of extensive gypsum strata previously mapped beneath the Mosul dam. Increased subsidence rate over the past five years has been due to periods when there was little or no regrouting underlying the dam basement. Dam subsidence currently seems to follow a linear behavior but on can not exclude a future acceleration due to increased gypsum dissolution speed and associated catastrophic collapse of the dam (Milillo et al., 2016).

Given the existing geologic knowledge base in the 1980s, in my opinion, one must question the seeming lack of understanding in a group of well-paid consultant engineering firms as to the outcome of building such a major structure, atop what was known to be an active karstifying gypsum succession, sited in a location where failure will threaten multimillion populations in the downstream cities. The same formation that constituted the base to the Mosul dam was known at the time to be associated with ground stability problems atop similar gypsum-cored anticlines in the city of Mosul to the south. Even more concerning to the project rationale should have been the large karst cavities in highly soluble gypsum that were encountered a number of times during feasibility and construction of the dam foundations (Figure 14). Or, perhaps, as Lao Tzu observed many centuries ago, “ ...So the unwanting soul sees what’s hidden, and the ever-wanting soul sees only what it wants.”

Canals, like dams, that leak in gypsum karst areas can trigger subsidence, which can be severe enough to cause retainment failure. In Spain, the Imperial Canal in the Ebro valley, and several canals in the Cinca and Noguera Ribagorzana valleys, which irrigate parts of the Ebro basin, have on numerous occasions failed in this way. Similarly, canals in Syria have suffered from gypsum dissolution and collapse of soils into karstic cavities. Canals excavated in such ground may also alter the local groundwater flow (equivalent to losing streams) and so accelerate internal erosion, or the dissolution processes and associated collapse of cover materials. In the Lesina Lagoon, Italy, a canal was excavated to improve the water exchange between the sea and the lagoon. It was cut through loose sandy deposits and highly cavernous gypsum bedrock, but this created a new base level, so distorting the local groundwater flow. The canal has caused the rapid downward migration of the cover material into pre-existing groundwater conduits, producing a large number of sinkholes that now threaten an adjacent residential area.

Pipelines constructed across karst areas are potential pollution sources and some may pose possible explosion hazards. The utilization of geomorphological maps depicting the karst and subsidence features allied with GIS and karst databases help with the grouting and management of these structures. In some circumstances below-ground leakage {Zechner, 2011 #26} from water supply pipelines can trigger severe karstic collapse events. Where such hazards are identified, such as where a major oil and gas pipeline crosses the Sivas gypsum karst in Turkey, the maximum size of an anticipated collapse can be determined and the pipeline strength increased to cope with the possible problems.


Solving the problem?

Throughout the world, be it in the US, Canada, the UK, Spain, eastern Europe, or the Middle East, it is a fact that weathering of shallow gypsum forms rapidly expanding and stoping caverns, especially in areas of high water crossflow, unsupported roof beams, and unconsolidated overburden and in areas of artificially confined fresh water. Rapid karst formative processes and mechanism will always be commonplace and widespread (Table 2). Resultant karst-associated problems can be both natural and anthropogenically induced or enhanced. It is fact that natural solution in regions of subcropping evaporites is always rapid, and even more so in areas where it is encouraged by human activities, especially increased cycling of water via damming, groundwater pumping, burst pipes, septic systems, agricultural enhancement and uncontrolled storm and waste water runoffs to aquifers.

Typically, the best way to deal with a region of an evaporite karst hazard is to map the regional extent of the shallow evaporite solution front and avoid it (Table 3). In established areas with a karst problem the engineering solutions will need to be designed around hazards that will typically be characterised by short-term onsets, often tied to rapid ground stoping/subsidence events and quickly followed by ground collapse. If man-made buildings of historical significance are to be restored and stabilized in such settings, perhaps it is better to wait until funds are sufficient to complete the job rather than attempt partial stabilization of the worst-affected portions of the feature. Significant infrastructure (including roads, canals and dams) should be designed to avoid such areas when possible or engineered to cope with and/or survive episodes of ground collapse.

A piecemeal approach to dealing with evaporite karst can intensify and focus water crossflows rather than alleviate them. In the words of Nobel prizewinner, Shimon Peres; “If a problem has no solution, it may not be a problem, but a fact - not to be solved, but to be coped with over time.”


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Salt Dissolution (2 of 5); Salt caves

John Warren - Saturday, September 30, 2017

 

Properties of evaporite karst

At first sight, medium-scale evaporite karst surface landforms, such as dolines, polje-like depressions, subsidence bowls and collapse dolines, appear near-identical to those found in and on carbonate karst (Part 1). Likewise, all the smaller-scale intracavern features, both dissolutional (pipes and cavities) and constructive calcite speleothems), can be preserved as karst zones in and above an evaporite host, even when the formative hydrology becomes inactive and the system is buried (palaeokarst). But, the much higher rates of dissolution of halite and gypsum compared to limestone or dolomite means there are significant differences in rates of formation and cave geometries compared to carbonate caves in nonevaporitic hosts. One cannot simply take models of caves developed by studies of carbonate karst and use them to unreservedly interpret evaporite caves.

Differences reflect the inherently higher solubility and flowability of the host salts at earth surface conditions and mean most evaporite-hosted caves show peculiarities related to higher dissolution rates compared to carbonate karst. For example, one most significant in a vadose halite cave is that evaporation of a near-saturated brine is aided by airflow and means many halite stalactites tend to curve into the direction of air flow (anemolites) rather than the subvertical dripstone mechanism that controls the shape of most calcite stalactites and stalagmites. Forti (2017) is an excellent summary of evaporite speleothems and formative mechanisms in both gypsum and halite caves.

The high solubility of subsurface evaporite beds atop deeply-circulating pressurised and jointed aquifers (confined hydrologies) means deep-phreatic chimneying (vertical shafts) and deep dissolution with upward-stoping are a common deep cavity-forming mechanism in and atop buried salt beds. This, in turn, means high volumes of overburden sediment can be swallowed quickly, especially in suprasalt regions with strong roof spans covered by loose sediment (Figure 1). In such settings, large at-surface natural sinkholes and circular structures (breccia chimneys) can daylight in days or even hours and so constitute significant geohazards (see Parts 4 and 5). This type of rapid point-sourced vertical stoping atop evaporites doesn’t just happen in continental settings. Similar circular structures atop breccia-filled vertical shafts or fault-bound chimneys form beneath deep waters of the Eastern Mediterranean via substratal breaching of the Messinian evaporite, followed by upward stoping through the overlying sediments (Bertoni and Cartwright, 2005).


Worldwide, much of the dissolution action in uplifting salt beds and masses begins well below the watertable. It takes place hundreds of metres or more below the surface at bathyphreatic depths where; 1) halite is leaching, 2) anhydrite is reconverting to gypsum and 3) where caprocks and other dissolution residues are forming atop diapiric salt features. Early dissolution is greatest at contacts between the salt edges and joints or fractures in adjacent aquifers. Thus patterns of jointing or fracturing control the extent and style of caves in a salt host and typically creates high-density maze caves. With halite, the only setting where a NaCl mass makes it to the surface is as an active or recently active diapir crest or namakier. Before its dissolution, the salt matrix hosting a cave-hosting diapiric halite is largely impervious, flowing and re-annealing. Hence, most of the karst action accessible for study in halite caves is vadose and tied to perched water tables; this is most obvious near the margins of namakier salt sheets and located well away from locations of active salt fountains.

If phreatic caves ever do form near to rising salt fountain domes, these early cavities are quickly closed by the pressurised salt flow needed to bring salt to surface. Rather than being telogenetic features, many of the metre-scale blebs of finely-layered gravitationally-aligned laminar halite seen within diapirs intersected in deep salt mines, and floating in a matrix of flow foliated coarse-crystalline halite, are the result of mesogenetic salt precipitating in gas-filled open cavities within the diapir mass (see Salty Matters October 31, 2016). These laminar cavity filling growth-aligned halite textures are not telogenetic, neither are the entrained salt clasts preserving primary (relict) depositional texture of the mother salt. Instead, they indicate the existence of former or present-day gas-filled cavities (N2, CH4 or CO2), which are a hazard occasionally encountered during salt and potash mining, for example in the diapirs of NE Germany (Hedlund, 2012). Their presence and their ability to blow out mine walls show that pressurised gas pockets and associated cavern fill textures are not unusual mesogenetic features in a flowing salt mass. Away from deep mine intersections of mesogenetic gas-filled cavities, most of the shallow cavities and fill textures of a flowing salt body and studied by speleologists are telogenetic and mostly vadose.

Cave-forming processes in a telogenetic salt cave in bedded salt bodies are either vadose or phreatic (both shallow and deep), with many gypsum/anhydrite caves showing textural evidence indicating transit from one realm to the other. At any one time, the transition from phreatic to vadose landforms is tied to depth to watertable. This is seen in the gypsiferous cap karst to diapirs and allochthons of Triassic salt in the Betic Cordillera of Spain (Calaforra and Pulido-Bosch, 1999). The crests of the salt structures are dominated by collapse dolines (vadose), this passes radially out into the belt of solution dolines occupied by seasonal saline lagoons and further out into rim of saline springs that form wherever the watertable intersects the landsurface (phreatic). Some caves pass from phreatic to vadose a number of times in their histories in response to watertable fluctuations tied to varying climate and tectonics (Columbu et al., 2015.

Vadose processes characterise the uppermost part of a karst aquifer and have air in pores above a water surface or perched watertable. Water drainage is free-flowing under gravity; cave passages drain downslope and show strong gravitational orientations (numerous sub-vertical features). Because they form above the watertable, vadose cave walls are subject to surface seepage, evaporation and drying. Speleothems decorating vadose cave walls indicate varying combinations of gravity and airflow and include; stalactites, stalagmites, cave popcorn, helicites and flowstones.

Caves can form in a salt bed on its way down (eogenesis; syndepositional to early burial phreatic), on its way up (evaporite telogenesis, generally begins at bathyphreatic depths) and at the bottom of the bed’s burial history (mesogenetic; e.g., gas-driven bathykarst).


Gypsum caves

Active caves beneath gypsum karst landforms first formed as deep phreatic maze caves and are characterised by dense passage networks with numerous contemporaneously closed loops, typically in bedded evaporites. Vadose cross sections of the same system can be quite large due to the high solubility and relative homogeneity of the host, especially if formed in a thick capstone. Heimkehle cave in the Zechstein anhydrites in the Harz Mountains of Germany has an overall passage length of more than 2 km, with large rooms up to 22 metres high and 65 metres wide. It was large enough to be used in World War II to house a factory manufacturing parts for JU88 aeroplanes (Knolle et al., 2013).


Other gypsum caves, some more than 150-200 km long, are well documented in Miocene gypsum hosts in west Ukraine (Table 1; Figure 2; Klimchouk, 2000; Klimchouk and Andrejchuk, 2003; Andrejchuk and Klimchouk, 2004; Klimchouk, 2007) and the Gypsum Plain of West Texas and New Mexico (Stafford et al., 2008, 2009, 2017) and much of our current understanding of the gypsum karst process comes from these regions.

Gypsum cave walls in the vadose realm are relatively undecorated compared to carbonate caves. But calcitic speleothems can form in a gypsum cave and are well documented, as are alabastrine flowstones and selenite rinds on cave walls, for example; in northwest Texas (McGregor et al., 1963), in Alabaster Cave and others in the Blaine Gypsum in western Oklahoma (Johnson, 1996; Bozeman and Bozeman 2002), and in quarries intersecting the gypsum karst system in the Kirschberg Evaporite member near Fredericksburg, Texas (Warren et al., 1990).


Phreatic gypsum caves

Phreatic caves in bedded evaporites atop artesian aquifers typically grow as upward stoping and branching blind flow loop caverns (phreatic cupolas) and chimneys, which can begin to grow deep below the watertable as bathyphreatic or hypogenic karst. For example, deep artesian systems drive cupola karst, and blind chimney stopes in the Black Hills of Dakota, the Elk Point Basin of Canada and the Optymistychna Cave, western Ukraine (Figures 2, 4a). Deeper in a basin, where bathyphreatic caves are bathed by centripetal mesogenetic crossflows, water flow is slow and phreatic karst is influenced by the escape of H2S- and CO2-rich basinal waters, not necessarily by the confined meteoric head. Fluid flow at these greater depths is driven by pore water gradients that reflect potentiometric variations in temperature, organic maturation, pressure and salinity of basinal waters.

      

Passages in mesogenetic and telogenetic evaporite-hosted caves tend to first develop near fractures and joints in adjacent aquifers and then expand into maze cave networks. Once growing in the main evaporite mass, some phreatic maze passages can show internal upward-directed switchback gradients independent of jointing in adjacent aquifers. Cave orientations within the evaporite bed and away from the aquifer inflow level are tied to internal inhomogeneities in the host bed such as less soluble or more soluble intrasalt beds and changes in mineral proportions.

Phreatic tubes are the most common passage shape in smooth-walled blind-dissolution pockets and cupolas in bathyphreatic gypsum caves. Tube or channelway shapes range in cross section from near-circular (isotopic dissolution of a soluble host) to elliptical tubes to canyon-shaped keyholes (Figures 3, 4a, b; Klimchouk, 1992, 1996). Ornamentation is minimal where undersaturated crossflow drives the rapid dissolutional breakdown of the cave wall and the resulting passages are smooth, with local scalloped dissolution irregularities. Dense intersecting maze cave networks first form in the early stages of exhumation at the contact between a tight but soluble gypsum/anhydrite unit and a less soluble carbonate or siliciclastic aquifer (Figure 2, 3). This less soluble bed is the supply conduit for groundwater crossflows that dissolve the edges of the initially impervious anhydrite/gypsum bed. The greatest rate of water supply to the dissolving gypsum contact is along joints and fractures in the adjacent aquifer bed (especially with limestone aquifers). Accordingly, the meshwork of caves penetrates the gypsum bed and tends to follow joint and fracture patterns of the adjacent aquifer.

More stagnant phreatic portions of telogenetic CaSO4 cave systems can be saturated and so precipitate isopachous crusts and crystal rinds. Gypsum is the commonplace isopachous precipitate in voids in this setting, while anhydrite tends to dominate in cavities in the deeper mesogenetic realm (Garcia-Guinea et al., 2002). By the time phreatic maze caves are exhumed and enter a vadose (epigene) setting, where they are accessible for speleological study, much of the earlier phreatic ornamentation has already been dissolved by the increasing undersaturated throughflow and cavern interconnection. This is associated with uplift into the more hydrologically-active phreatic realm, which always precedes entry into the vadose realm. (Warren et al., 1990).

Some of the longest and most complex phreatic maze cave systems in the world are found in Miocene gypsum in west Ukraine. Optymistychna Cave, with more than 214 km of surveyed passages, is the world’s longest gypsum cave and the second or third longest cave of any type (Table 1; Figure 2; Klimchouk, 2007). The world’s longest cave, at 550 km, is the carbonate-hosted Mammoth Cave of Kentucky. The West Ukraine region contains the five longest known gypsum caves in the world, accounting for well over half of the total known length of known gypsum caves on Earth. By area and volume, the world’s largest gypsum caves are: Ozernaja (330,000 m2 and 665,000 m3; with 122 km of documented passages it is also the world’s 10th longest), Zoloushka (305,000 m2 and 712,000 m3), followed by Optimisticheskaja Cave (260,000 m2 and 520,000 m3). They are all complex joint-controlled maze caves, formed under confined aquifer conditions that existed from the Pliocene to the Early Pleistocene (that is karstification on the way up - telogenesis or gypsum exhumation, initially focused on the underside of the bed). Their growth patterns indicate upward-transverse phreatic groundwater circulation, with ultimate cavern fusion across the gypsum bed. All these west Ukrainian caves were fed by artesian crossflow in sub-gypsum and supra-gypsum aquifers that were sourced in the Carpathian Mountains.

High rates of dissolution in phreatic gypsum caves, relative to rates of water crossflow, are indicated by bevelled, faceted and “keyhole” cross sections in the Ukrainian caves along with a lack of vadose wall ornamentation (Figures 3a 4; Klimchouk, 1996; Pfeiffer and Hahn, 1976). Keyholes indicate density stratification and convectional circulation in cave-forming waters, with shapes sometimes complicated by lithological discontinuities in the gypsum bed (Figure 3b; Kempe, 1972). Convection in caves is most pronounced where sluggish artesian flow and low flow velocities dominate.

This was the case in the Pliocene to early Pleistocene history of the maze caves of the western Ukraine (Klimchouk and Andrejchuk, 2003). At that time the deeply buried gypsum dissolved via upward growing but blind, phreatic cavities, with a reflux of somewhat denser “spent” waters sinking toward the base of the cave. Spent water was replaced by less dense inflow waters supplied from the lower aquifer (Figure 3a). This set up a natural density-stratified convection, which maintained fresher (less dense) waters near the phreatic cave roof. Upward growing blind caves tend to expand more at their tops driving the transition from subcircular to keyhole caverns along the cave conduit (Figure 4b). After dissolving gypsum and increasing in density, a portion of the “spent” cave water sank all the way back into the underlying aquifer where it once again joined the regional throughflow in the lower aquifer. Once a stoping cave breached the top aquifer water flow direction in the cave was controlled by temperature, pressure and density contrasts between aquifers on either side of the gypsum bed. It seems that post-breach most of the cave water continued to rise through the cave system into the overlying aquifer (Figures 3a, 4) Similar keyhole and cupola morphologies are developed in low flow rate bathyphreatic sulphuric acid caves in carbonate hosts (e.g., early stages in the formation of the Carlsbad and Lechuguilla caverns).

Formation of keyhole passages is not an exclusively phreatic phenomenon in density-stratified gypsum caves. Keyholes in the vadose portions of many telogenetic carbonate caves indicate the transition of the cave passage from phreatic, with circular cross sections, to vadose with deepening drainage slots at the base of the passages (e.g. Calaforra and Pulido-Bosch, 2003). Phreatic stoping, followed by vadose cavern enlargement, probably explains the close correlation between caprock sinkhole distribution and position of underlying vadose passages in Permian gypsum subcrop in the Kungur Cave region in the Russian Urals, where the caprock thickness is a little as 25 m (Figure 5).


Vadose gypsum caves

A lowering of the watertable, either by uplift or climatic change tied to increasing local aridity, converts a former hypogene to meteoric phreatic cave into a vadose cave. In the latter, the walls become ornamented with gypsum and calcite speleothems. This is the recent history of the accessible portions of the gypsum maze caves worldwide, including documented examples in New Mexico, the western Ukraine and Saudi Arabia where the passage into the middle-upper Pleistocene marks the transition from phreatic to vadose in most of the accessible caves (Stafford et al., 2008). It is characteristic also of the very recent history of some natural phreatic caves where water tables were artificially lowered to allow quarrying of gypsum (Warren et al., 1990; Klimchouk, 2012).

Climate shifts or watertable fluctuations at the early end of the burial cycle (salt on the way down) can create alternating vadose-phreatic conditions in evaporite beds in the early stages of burial and so create an early watertable-associated karst level in the accumulating evaporites. That is, a gypsum bed hosting a vadose hydrology on its way down into burial, may pass through the watertable a number of times before its final burial and passage into the mesogenetic realm. This is the case today in the captured recharge playas in central Australia and about the edges of some halite-filled salars in the Andes and many Canadian salt lakes (Last, 1993; Warren, 2016), where fluctuating hydrologic conditions alternate between vadose and phreatic. Similarly, Quaternary climate shifts have variably karstified the gypsiferous sediments of many Sinic playas. For example, Yaoru and Cooper (1997) document Pleistocene lake basins in north-west China, such as in the Chaidamu Basin, where exposed gypsum beds evidence karst overprints that include: corroded flutes, fissures, small caves and associated collapse breccias and roof falls, followed by phreatic evaporite cement overprints. Watertable fluctuation is a hydrological overprint that is preserved as alternating ornamented surfaces in gypsum caves, disconformities and cave fills in many ancient lacustrine gypsum units.

Halite Caves

Because of its high solubility, halite does not make it into outcrop or shallow subcrop as easily as gypsum/anhydrite. Where halite is at the surface, it tends to be in regions of Pleistocene halite deposition (salt on its way down) or in zones of active diapirism (salt coming up very quickly). Chabert and Courbon (1997) noted caves in ancient rock salt in several regions: Algeria (in diapirs, mostly as vertical to sub-vertical shafts and short caves up to 28shafts0 m long), Chile (diapirs, with caves 250-500 m long), Israel (in Mt Sedom diapir as vadose caves and tube caves several hundred metres in length and subvertical vadose shafts that are metres across), Romania (in diapirs, with caves up to several hundred metres long), Spain (in diapirs, with caves up to 650 m long), Tajikistan (several caves 300 to 2,500 m long and up to 120 m deep), and in the namakiers of Iran and the offshore island relicts where cavern lengths range from several hundred metres to kilometres. The maintenance of landscape elevation, which can be hundreds of metres above the surrounding plains, facilitates the creation of vadose caves in the diapir crest (Table 2).


Halite in an active namakier rises and spreads rapidly, so any karst in an active salt diapir tends to be a feature associated with the immediate underside of a caprock. Karst processes cannot deeply penetrate while salt is flowing, even when the plug rises more than 300 metres above the surrounds. So, more extensive halite hosted caves are best developed about the margins of a namakier where halite’s susceptibility to rapid dissolution means the length of a cave developed below the caprock can be substantial. There is a report of a single salt halite cave (Cave 3N) on Qeshm Island, Iran, with a passage length of more than 9 km (Bruthans et al., 2002). Once the rate of diapir rise has slowed or ceased, the positive topography of the now inactive diapir controls the depth of development of doline collapse inn the diapir itself. That is, deeper collapse dolines can now form in the more central topographically higher portions of a salt structure, once the rate of rise has slowed (e.g. Calaforra and Pulido-Bosch, 1999).

The high solubility of halite means a halite cave system can form in a few hundred years rather than the thousands to tens of thousands of years needed to form carbonate karst (Bruthans et al., 2010). Not only is rate of cavern formation swift (dissolution/downcutting of 20 mm/year), these rapidly forming karst features are hosted in and atop a flowing rock mass. This means some of our notions of karst process and cave stability, related to the rate and density of cavern expansion, need to be modified when dealing with halite karst. For example, unlike gypsum and carbonate karst, jointing is not significant in controlling cavern orientation in namakier karst.

      

Dead Sea karst

Halite caves occur in the Mt. Sedom diapir, where halokinetic Miocene salt is sporadically exposed beneath a weathering and fractured gypsiferous caprock (Figure 6a,b; Frumkin, 1996; Frumkin and Ford, 1995; Frumkin 2009). Water enters the various caves in Sedom diapir through breaches in the caprock (Figure 7b). Most of the caves in the higher parts of Mt. Sedom salt are vadose inlet caves; these are meandering steeply inclined tubes and canyon slots located within or immediately below the caprock. They form where salt solution quickly carves out near vertical slots and shafts (typically < 2m broad and much deeper) that lead down from the surface, sometimes along pre-existing fractures and shears in the salt. Inlet caves in the central portions of the mountain can only be accessed through their sinks and appear to have no distinct outlet (Figure 6c). All terminate several tens of metres above the regional watertable (e.g. Karbolot Cave; Figure 6c, d; Frumkin, 1994a, b, 1996).

The lower parts of inlet caves often contain steep silt and clay banks with surge marks that indicate occurrence of low energy water ponding, with variable residence times. Silt and clay sediments settling at the bottoms of inlet caves impede infiltration, extending the residence time of pond water (Frumkin, 1994a, 1996). Three of the studied caves in northern Mount Sedom had perennial ponds throughout the period 1984-1995. The ponds are perched, without any lithologic control, tens of metres above the nearest potential outlet at the foot of the mountain (Figure 6c). The water level in each pond differs from the others by tens of metres. All pond waters are highly concentrated, up to 324 g/l, with solutes consisting mainly of sodium and chlorine. Fresh inflow waters reach halite saturation within a few hours of reaching the pond. Both dissolution and precipitation features form the pond edges, and their equivalents can be seen on cave walls wherever ponds have dried out. Dissolution is indicated by horizontal notches, which connote density stratification in the ponds when aggressive fresh flood waters are temporarily diluting the upper parts of the pond. Subsequent saturation of holomictic pond waters is indicated by the growth of cm-scale halite crystals on the bottom and sides of the ponds.


Towards the periphery of Mt Sedom, the inlet caves lead down to laterally expanding vadose cave levels that drain onto the Dead Sea plain (Figure 6c,d, 7c). Sedom Cave, is the longest laterally expanding diapir cave, with an aggregate length between two subparallel conduits of 1.8 km. Malham Cave, another large perched and laterally expanding cave, lies a few hundred metres south of Sedom Cave (Figure 6a, c, d; Frumkin, 1996). It has an aggregate passage length of more than 5.5 km and reaches to some 194 m below the landsurface. There is an upper tier of mostly inactive passages and a lower active channel level. 14C dates on fossil wood in the upper cave level shows meteoric waters began to sculpt the upper cave more than 5,500 years ago. Ongoing uplift of Mt Sedom salt means the active channel level in Malham Cave is now downcut some 10-12 metres lower than when it began. Malham Cave passages quickly developed an open outlet through which floodwater escaped directly to the Dead Sea floor, proving that during this period some 4000 years ago rock salt had already risen above region hydrological base level at the Malham outlet point within the eastern escarpment (Figure 6c). Lashelshet Cave, an inlet cave on the highest point on the diapir cross section, has an even older age of more than 7,000 years since cave initiation. Caves in the northern part of Mt Sedom did not begin to form until some 3,000 years later (Frumkin, 1996).

Based on their study of the caves of Mt Sedom, Frumkin and Ford (1995) concluded cave passages develop in two main stages: (1) an early stage characterized by inlet caves with high downcutting rates into the rock salt bed, and steep passage gradients; (2) a mature laterally expanding stage characterized by lower downcutting rates and the establishment of a wider subhorizontal perched stream bed armoured with alluvial detritus. This style of cave tends to develop toward the periphery of the diapir mound. In the mature expanding stage downcutting rates are controlled by the uplift rate of the diapir and changes of the level of the Dead Sea.

Passages may aggrade to create wide flat bevelled passages and slots with thick sediment armoured bases (Figure 7c; Frumkin, 1998). A lack of a consistent phreatic level in the blind bottoms of perched water levels and the presence of the horizontal slots in the lower levels of Sedom Cave means dissolution in both types of caves is largely restricted to times of flooding and perched or backed up freshwater in the vadose zone. This explains the tapering passages of inlet caves and the widespread alternation of armouring and bevelling as well as formation of narrow horizontal meandering slots toward parts of the top of the meandering channel that is now Sedom Cave.

Mass balance calculations in the halite caves of Mt Sedom yield downcutting rates of 0.2 mm s-1 during peak flood conditions, this is about eight orders of magnitude higher than reported rates in any limestone cave stream (Frumkin and Ford, 1995). However, floods have a low recurrence interval in the arid climate of Mount Sedom so that long-term mean downcutting rates are lower: an average rate of 8.8 mm/ year was measured for the period 1986-1991, while Frumkin (2000) estimate the average regional vadose downcutting rate in the Mt Sedom karst region to be 20 mm/year. This is still at least three orders of magnitude higher than rates established for limestone caves and more than able to cope with the rate of supply of the diapiric salt.

The highly impervious nature of halite and its resupply in actively growing diapirs means that, unlike carbonate and gypsum caves, there is no real watertable level to define maximum cave development in a rising salt stem. Rather, the inlet caves are simple dissolution tubes where rainwater has accessed halite and sank until it was saturated and then dissolution stopped until the next flood. Toward the edge of the rising stem, these inlet caves breached the edge of the salt mound and vented their perched groundwaters to the surrounding plain (Figure 6; Sedom and Malham caves). This creates a downcutting and laterally expanding cave system, which is still some metres to tens of metres above the base level of the regional watertable in the surrounding plain. The expanding cave level is dominated by mostly horizontal growth, often with a sediment-armoured floor. It has numerous benches in the walls that probably reflect changes in the hydrological base level.

Dense anastomose cave networks that characterise gypsum caves are not found in the halite caves of Mt Sedom. This reflects the ability of diapiric halite to re-anneal and the fact that all exhumed halite that makes it to the surface is diapiric, not bedded. At-surface halite is not sandwiched between jointed aquifers above and below the dissolving layer. Rather it is a growing mound subject to dissolution at its top and sides.


Away from Mt Sedom, there are active collapse sinkholes and caverns forming in the alluvial fans and clastic aprons that overlie the bedded Quaternary lacustrine halite of the Dead Sea (Figure 8a, b; salt on its way down in the burial cycle). In the sediments around the lakeshore, the pace of karst collapse has accelerated in the last 60 years due to a drastic lowering of the circum-lake watertable and the associated lakeward migration of the saline-fresh water interface (Figure 9). For example, a series of collapse dolines 2-15 m diameter and up to 7 m deep, appeared in 1990 in the New Zohar area. In January 2001 a large sinkhole, some 20 m deep and 30 m wide, cut through the asphalt surface of the main road along the western shore of the Dead Sea. It was opened by the passage of a busload of tourists on their way from Ein Gedi to the Mineral Beach solarium. Existing tourist facilities, such as the Ein Gedi beachside parking, were shut down after the road was damaged and several buildings have since collapsed into sinkholes. Sinkholes have since developed in other areas about the Dead Sea Margin including Qalia, Ein Samar and Ein Gedi. The process began in the southern part of the Dead Sea coast and slowly spread northward along the Israeli coast. Collapse is more localised in the northern and southern regions on the Jordanian side, and across the region, continues to increase in frequency as the sea level falls (Ezersky and Frumkin, 2013).

Three main types of sinkhole or doline fill have been recognized atop the dissolving Holocene salt beds of Mt Sedom; 1) Gravel holes in alluvial fans, 2) mud holes in the intervening bays of laminated clay deposits between fans, and 3) a combination of both types at the front of young alluvial fans where they overlap mud flats. Fossil, relict sinkholes have been observed in the wadi channels cutting into some old alluvial fans, showing this is a natural and ongoing process. While lake levels continue to fall (Figure 9b,c), the potential for subsidence hazards related to karst collapse is ongoing.

Sinkholes and related subsidence have been the focus of much geological study of the halite caves, but Closson et al. (2010) pointed out that an even more significant and ultimately damaging environmental effect of the ongoing water level lowering is the hectometre and larger scale landslides along the retreating shorezone. In the 1990s, international builders created major tourist resorts and industrial plants along the Jordanian and Israeli shore while, during the same period, geological hazards triggered by the level lowering spread out. From the beginning of the year 2000, sinkholes, subsidence, landslides, and river erosion damaged infrastructures more and more frequently: dykes, bridges, roads, houses, factories, pipes, crops, etc. all suffered as a result.

There is evidence of an older set of widespread ground collapses, sinkholes and caves that are tied to an earlier substantial fall in the Dead Sea water level some 4 ka. It may even be that the events described in Genesis 14 in the Christian Bible took place at the time of a substantially lowered sea level. The described battle, which occurred prior to the fall of Sodom and Gomorrah, perhaps took place on the subaerially exposed flats of the Southern Basin of the Dead Sea. The “pits of slime” described in the fall perhaps were solution collapse sinkholes activated by the 4000 ka fall in the Dead Sea water level (Frumkin and Elitzur, 2002).


Halite karst in diapiric Hormuz salt, Middle East

Namakier outcrops in and about the Arabian Gulf range from structures actively extruding salt to those in ruins where salt has not flowed for tens of thousands of years (Figure 9). Likewise the halite caves developed in the namakiers of Iran, or their offshore island counterparts, show a broad range of ages and styles of salt cave development tied to the time since cessation of salt flow (Filippi et al., 2011; Bruthans et al., 2010; Talbot et al., 2009).

Surfaces of actively flowing namakiers on the Iranian mainland are characterised by karren flutes and pinnacles, with numerous small-medium dolines, collapse structures, swallow holes and small caves at their base. As at Mt Sedom, caves tend to be sediment-armoured meandering tube caves or subvertical canyon slots that are centred on joints in the salt beneath a thin suffusion mantle. In contrast, the halite caves in the diapiric cores of the many islands in the Arabian Gulf have a more mature bevelled meandering style with thicker sediment armouring on the cavern floor. Many of these caves breach the retreating edges of former namakiers and salt fountains.

Salt movement in the various diapiric cores of these islands is inactive, or is greatly reduced, compared to the Miocene when these structures were active namakiers. For example, Dragon Breath Cave on Hormuz Island is a linear meander tube cave fed by an ephemeral stream in a shallow valley filling with alluvium (Figure 11; Bosak et al., 1999; Filippi et al., 2011). The surrounding landscape is classic salt karst with numerous depressions, blind valleys, ponors and subrosion sinks. Together they form a highly pockmarked centripetally-ringed topography, which outlines those central parts of the island underlain by shallow subcropping Hormuz salt. The cave itself is one of a number of tube caves exiting about the edge of the zone of diapiric salt. Hosted in steeply dipping diapiric salt, it is around 100 m long with its main passage created by a minor ephemeral stream. Its near flat roof, with an average inclination of 5.4%, is a notable feature and reflects either joint-related dissolutional spalling of the roof or an earlier watertable slot related to backup of a freshened water body.

The current cave passage has cut down a metre or more into earlier cave floor sediments (sediment armour), which contain clasts up to 50 cm in diameter. The cave formed by initial ingress along a linear joint, which was then widened by salt dissolution, so allowing meandering of the stream trace within the salt. It is a cave system very similar to the mature stages of laterally expanding caves in Mt Sedom. The base level of the cave correlates with the surface of a widespread marine terrace, which is now uplifted some 20 metres above sealevel and defines much of the periphery of Hormuz Island. The raising of the marine terrace is related to the ongoing raising of the island via salt flow.


Bruthans et al. (2000, 2010) show that the style of karst landform developed in dissolving diapiric salt in the Arabian Gulf Islands reflects the thickness of the carapace that caps the dissolving salt core (Figure 10). They distinguished four classes of diapir cap, each with a particular association of superficial and underground karst forms, namely: 1) outcropping salt, 2) thin capping (0.5-2 m), 3) capping with moderate thickness (5-30 m), 4) capping with greater thickness (more than 30 m). Cap thickness controls or reflects: 1) the density of recharge points, with high densities of recharge points in the thinner caps; 2) the amount of concentrated recharge which occurs at each recharge point, with suffusion karst characterising thinner caps; 3) the rate of lowering the ground surface atop the salt, with the faster rates of lowering occurring beneath thinner caps, and 4) the amount of load transported by underground flood-streams into cave systems. The volume of sediment load tends to be locally higher and focused beneath the thicker caps, particularly when inflow streams abut the edges of a dissolving salt dome. The thickness of caps atop expanding halite caves does not appear to influence the shape or style of the cave developed within the salt mass; more important seems to be the thickness of cap in the recharge area of the cave and the type of recharge into the salt environment. That is, how much water is passing into the salt and is its flow ongoing or ephemeral?


Halite caves in the relatively mature salt stems of the various islands of the Arabian Gulf, unlike carbonate systems, can swallow and store huge volumes of clastic sediment, volumes that would clog the entrance to a carbonate system. The extreme solubility of halite enables the pace of dissolution/corrosion enlargement in a salt cave to keep pace with large amounts of sediment carried into the cave by external inflows (Figure 11b). Stream sediments arriving at the cave entrance, including boulders, move inside and are trapped within the salt itself. Sediment is not dumped outside the cave entrance, which is the typical situation in blind valley river mouths at carbonate caves (Bruthans et al., 2003). For example, coarse-grained sediment fractions are carried hundreds of metres into the cave by two large intermittent streams entering the upper part of the Ponor Cave (Hormuz Island). The clasts in the resulting intra-cave alluvial fan conglomerates range from cobble of several centimetres up to 1 m diameter boulders; only sand-sized particles make it to the lower part of the same cave.

Caves capable of storing such coarse alluvium within the cavern itself are halite-specific with no equivalent in a carbonate karst terrain. There the boulder-size fraction in the cave itself is the result of roof fall, and almost all stream-borne coarse alluvium is deposited outside the cave.

 

Evaporite Speleothems

Cave walls in zones of less intense dissolution and stream crossflow are decorated with halite, gypsum and anhydrite speleothems (Figure 12). Halite has a much higher potential to form macro- or mono-crystalline speleothems than calcite and gypsum (Forti 2017). Therefore, in most of the studied halite caves around the world, relatively large euhedral or hopper halite crystals have been observed as in the Iranian and Atacama caves (Forti, 2017; De Waele et al., 2009). The preferred location for these crystals are the pools in the cave entrances, where evaporation is sufficiently low to allow the development of euhedral crystals up to 10 cm in size (Figure 12, 13; Fillipi et al., 2011).

In the Iranian caves halite macrocrystals normally form also along streams, whereas in the Atacama Desert they are completely lacking. This is because they need time to develop and therefore the stream must remain active for at least a few days (Figure 12a; Filippi et al., 2011).

Monocrystalline stalactites are widespread in the Iranian caves; the most common of which are the skeletal forms (Figure 12a; Filippi et al., 2011). An idealised skeletal stalactite normally consists of a central rounded “stalactite” from which, at different heights, three smaller and shorter rounded branches develops, being equally spaced at an angle of 120° (Figure 12b). Moreover, each branch and the central stalactite form an angle of ~70°. These values show that the directions of the central columns and the side twigs correspond to that of the four cube diagonals. Finally, at the end of each twig, there is a small halite crystal with one of its diagonals perfectly coincident with the twig (Figure 12b).

It is therefore evident that the entire structure of these peculiar stalactites consists of a single crystal lattice, albeit with a fractal appearance, and this fact is also confirmed by the presence along the rounded column of evident crystal facets oriented in the same direction (Figure 12b). Air currents and other local perturbing factors may cause a deflection from the theoretical direction of both the main column and of the side twigs (anemolites; Figure 12a). Finally, the rounded structure of the central column and of the external twigs is normally covered by glazy halite suggesting that cycles of deposition and dissolution alternate, while no inner feeding tube is present within the central column.


By these observations, the genesis of these skeletal monocrystalline stalactites is induced as driven by solutions mainly coming from brines and sprays, that then flow via gravity and capillarity only on the external surface of the stalactites. The amount and the composition of these solutions must change in time, becoming sometimes slightly undersaturated, probably during rain falls. The location of these speleothems close to waterfalls, where sprays are easily formed support that interpretation (Figure 13; Filippi et al., 2011; Forti 2017).

Compared to the growth rates of calcite speleothems in carbonate caves the growth rates of evaporite speleothems are phenomenal. Halite stalactites, several metres long and curving into the direction of airflow, formed in the mouth of Dragon Breath cave in a few years rather than millennia needed for carbonate counterparts (Bokacs et al., 1999). In August 1997, a network of numerous halite stalagmites and stalactites blocked the entrance to Dragon Breath Cave. In March 1998 there were no remains of the speleothem meshwork, while in February 1999, the stalagmites had reappeared (Bosak et al., 1999). Similar halite structures occur in the caverns of Mt Sedom and on the wet roofs of some salt mines.

Telogenetic halite deposits forming in namakiers encompass a range of mechanisms and speleothem textures (Figure 13; Filippi et al., 2011): i) via crystallization in/on streams and pools, ii) from dripping, splashing and aerosol water, iii) from evaporation of seepage and capillary water, and iv) other types of evaporative deposits. The following examples of halite textures are distinguished in each of the above-mentioned groups: i) euhedral crystals, floating rafts (raft cones), thin brine surface crusts and films; ii) straw stalactites, macrocrystalline skeletal and hyaline deposits, aerosol deposits; iii) microcrystalline forms (crusts, stalactites and stalagmites, helictites); iv) macrocrystalline helictites, halite bottom fibres and spiders, crystals in fluvial sediments, euhedral halite crystals in rock salt, combined or transient forms and biologically induced deposits. The occurrence of particular forms depends strongly on the environment, in particular on the type of brine occurrence (pool, drip, splashing brine, microscopic capillary brine, etc.), flow rate and its variation, atmospheric humidity, evaporation rate and, in some cases, on the air flow direction. Combined or transitional secondary deposits can be observed if the conditions changed during the deposition. Euhedral halite crystals originate solely below the brine surface of supersaturated streams and lakes.

Macrocrystalline skeletal deposits occur at places with abundant irregular dripping and splashing (i.e., waterfalls, or places with strong dripping from the cave ceilings, etc.). Microcrystalline (fine-grained) deposits are generated by evaporation of capillary brine at places where brine is not present in a macroscopically visible form. Straw stalactites form at places where dripping is concentrated in small spots and is frequently sufficient to assure that the tip of the stalactite will not be overgrown by halite precipitates. If the tip is blocked by halite precipitates, the brine remaining in the straw will seep through the walls and helictites start to grow in some places.

Macrocrystalline skeletal halite deposits and straw stalactites usually grow after a major rain event when dripping is strong, while microcrystalline speleothems are formed continuously during much longer periods and ultimately (usually) overgrow the other types of speleothems during dry periods. The rate of secondary halite deposition is much faster compared to the carbonate karst. Some forms increase more than 0.5 m during the first year after a strong rain event; however, the age of speleothems is difficult to estimate, as they are often combinations of segments of various ages and growth periods alternate with long intervals of inactivity.

Anhydrite forms speleothems in preference to gypsum in those rare parts of a cave with very high salinity waters, But overall gypsum speleothems dominate Some of these gypsum speleothems can be quite large, up to a few metres long.

Unlike halite and gypsum caves, which are rich in halite and gypsum formations respectively, anhydrite caves do not host anhydrite speleothems at all. This is a direct consequence of the CaSO4 – CaSO4.2H2O solubility disequilibrium, which makes the hydrated mineral (gypsum) less soluble than the anhydrous one (anhydrite) at normal cave temperatures, thus totally hindering the development of secondary anhydrite formations. Most of the gypsum produced by hydration replaces anhydrite within the rock structure, and therefore anhydrite does not form any speleothems. Nonetheless, a minor part of this secondary gypsum may develop some small deposits. In the caves of the Upper Secchia Valley, small gypsum crusts and flowstones were observed where condensation water, after dissolving anhydrite, flows over the gallery roof or walls where air currents induce evaporation (Chiesi & Forti, 1988). In the same caves, when per ascensum capillary flow and evaporation are possible, euhedral aggregates of small gypsum crystals may develop on top of rock. The interested reader is referred to a comprehensive paper by Forti, 2017 dealing the great variety of halite and gypsum speleothems.

Anhydrite karst is known from several countries of the world, but in almost all cases it is located at depths that make direct exploration almost impossible (Ford & Williams, 2007). This is the reason why until present, only caves from two locations (South Harz in Germany and Upper Secchia Valley in Italy) were explored and their speleothems studied.

The Upper Secchia Valley anhydrite caves and their chemical deposits were already known when the first monograph on speleothems in gypsum caves was published (Forti, 1996). However, at that time these caves were incorrectly considered as formed in gypsum and therefore their deposits were described along with those hosted in classical gypsum karst.

The genesis and evolution of the German and Italian anhydrite caves are completely different; in fact, the first are hypogenic caves (Kempe, 2014) and lack any natural entrance, whereas the second ones are epigenic and often develop very close to the surface (Malavolti, 1949). Therefore, chemical deposits are different in the two locations and restricted to the peculiar environment that controlled the evolution of the caves.

Leaving aside the widespread secondary gypsum produced by the hydration of the host rock, anhydrite caves are extremely poor in chemical deposits. The lack of minerals in the hypogenic caves is because they were filled with near-stagnant water for most of the time during their development. In the epigenic caves, instead, the absence of cave minerals is mainly attributed to the strong increase in volume caused by hydration of anhydrite (that turns into gypsum), which makes the wall and the ceiling of these cavities extremely fractured. In this latter setting, the rather continuous breakdown normally inhibits the development of even small chemical deposits, which, in any case, are easily washed away by the frequent floods that characterise the Upper Secchia Valley. Despite all these restrictions, the anhydrite caves proved to be interesting not only from a mineralogical point of view, as they host one cave mineral (clinochlore, Chiesi & Forti, 1985) restricted to this environment, but also for the presence of a unique gypsum/anhydrite speleothem, i.e., the huge “leather like sheets” of Barbarossa Cave (Figure 14).

      

Implications

Part 1 and Part 2 of this set of articles dealing with evaporite dissolution emphasise the importance of rapid rates of volume loss in creating a unique set of karst landforms and speleothems. This rapidity creates cavities in a hydrological milieu of contrasting brine salinity and temperature interfaces and permeability contrasts.This inherent association of voids in a setting with abrupt chemical interfaces facilitates the enrichment levels of economic commodities (part 6) and drives rapid bed stoping and foundering that forms zones of significant natural and anthropogenically enhanced geohazards in the landscape (parts 4 and 5).

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Dissolving salt (1 of 5): Landforms

John Warren - Thursday, August 31, 2017

 

Introduction

This series of five articles discuss effects of nearsurface evaporite dissolution at various scales and times, in terms of; 1) landform expression, 2) evaporite-hosted caves and speleothems 3) natural evaporite karst geohazards 4)Anthropogenically enhanced karst features, 5) economic associations tied to evaporite paleokarst (hydrocarbon and metal).

As salt dissolves in the subsurface, it creates void spaces of various sizes and shapes into which overlying strata can drape or brecciate, so creating characteristic landforms at both local (m to km) and regional (km to tens of km) scales (Figure 1). Water sources producing these features can be shallow unconfined meteoric or marine, confined shallow or deeply circulating meteoric, or hypogene, basinal and hydrothermal. Solution-related evaporitic landforms can be either active karst or palaeokarst features. Paleokarst refers to ancient karstic features that are no longer active and tied to ancient basin flushing hydrologies and now buried land surfaces. Active karst is responding to the modern ambient hydrology and typifies landscape elements discussed in the following section.


Evaporite solution karst landforms overprint nearby carbonate or evaporite strata, with many descriptive features and terms are common to textures and geometries across both host lithologies. Evaporite-related karst includes many varieties of karren, sinkholes (dolines and caves), blind valleys, poljes and subsidence basins (Ford and Williams, 2007; Warren, 2016). Bedrock exposures of gypsum or salt at a finer scale are sculpted into irregular curved and grooved surfaces called rillenkarren or other varieties of gravity-oriented solution flutes (Macaluso and Sauro, 1996; Stenson and Ford, 1993).

Local-scale karst geometries that define evaporite-related subsidence range in lateral extent from the metre-scale cones of suffosion karst to collapse cones to kilometre-scale doline depressions (Figure 1). In regions of soil-mantled outcrops of bedded evaporites, these features tend to occur at higher densities, within or near broad subsidence valleys. In the context of the lateral extents of evaporite-related landforms, it is also useful to separate landforms associated with the dissolution of bedded evaporite versus dissolution landforms and namakier residues that occur above variably active salt diapir crests. Features related to dissolving diapirs and namakiers are discussed in an earlier article (Salty Matters, March 10, 2015).

The geometry of the karst set-up hydrology is distinct in sediments above diapiric versus bedded evaporite substrate, and this article focuses on landforms related to a dissolving substrate made up of shallow dipping salt beds. The next section outlines characteristics feature of evaporite karst at the metre to kilometre scale, as defined in Figure 1. Then we will compartmentalise these features using the classification of Gutiérrez et al., 2008, and finally look at three of the more distinctive regions responding to the dissolution of subcropping evaporite beds.

Evaporite solution dolines

Dolines are closed circular to elliptical hollows or depressions, often funnel shaped, with diameters ranging from a few metres to a few kilometres and depths from a metre or so, to hundreds of metres (Figure 1; Ford and Williams, 2007). They indicate subsidence and collapse in underlying salt or carbonate units. Valley sides in larger subsidence dolines can be steep and expose karstified gypsum, or can be gentler slopes covered by soil.

Larger dolines may enclose one or more smaller sinkholes and can be further subdivided into, suffusion, subsidence sag, and collapse dolines (Figure 1). A collapse sinkhole is initiated by collapse of a roof span into an underlying solution cavity. A subsidence sag doline indicates a more diffuse broader and gentler lowering of the ceiling above the dissolving bed. Suffosion dolines are small-width high-density karst features (1-2m diameter) covered and filled with soil and debris that has washed or fallen into closely-spaced fissures cutting into the evaporite bed. They typically indicate that a dissolving salt mass is very close to the landsurface.

Dissolution, collapse and suffosion processes are more active, more rapid, more frequent and more noticeable in regions of shallow evaporite units, compared to carbonate terranes at the same depths. For shallow weathering evaporites versus carbonates, at roughly the same depths in the Perm and Bashkir regions of Russia, Gorbunova (1979) reported doline densities of 32 and 10/km2 respectively.


Suffosion dolines

Suffosion in a karst terrane describes the downward migration of unconsolidated cover deposits through voids. High levels of dispersed impurities (mostly clays and muds, often dolomitic and other dissolution residues) in a rapidly dissolving evaporite mass, means a soil or carapace of insoluble residues quickly covers the subcropping top and edges of a dissolving salt mass (Figures 1a, 2). Downwash transport into growing fissures and voids in the shallow underlying evaporite is via downwashing of fine particles carried by percolating waters, cohesionless granular flows, viscous sediment gravity flows (non-Newtonian), freefall of particles, and sediment-laden water. The carapace is continually undermined by ongoing rapid solution, as residual debris is continually and rapidly washed into the doline crevices into the growing cavities in the evaporite. This creates a unique soil-covered dimpled landscape, which is typified by a high-density doline terrain where the dissolving evaporite is only a metre or so below the landsurface.


Densities of up to 1000 suffosion dolines/km2 occur in many evaporite-cored fold axes or at subcrop contacts of evaporites with other outcropping lithologies. Densities of 1100-1500/km2 have been documented in evaporite karst in the Italian Alps and in regions to the west of Sivas, Turkey (Figure 3; Belloni et al., 1972; Kacaroglu et al., 1997). The inherently high solubility of evaporite salts explains similar densely-packed schlotten depressions and large karren shafts seen in gypsum karst of Antigonish County, Nova Scotia, where extrapolated doline densities in the latter case range up to 10,000/km2 (Martinez and Boeher, 1997). Extrapolated, because such dense networks do not extend over any more than a square kilometre or two and are typically found near retreating escarpment edges underlain by shallow subcropping salts. Such high densities tend to occur in more humid rather than arid areas, with high hydraulic gradients so that the resulting suffosion dolines have diameters around 5 m or less.


Subsidence Dolines

Solution (subsidence) dolines, as described in much of the literature tend to be larger down-warped doline craters or bowl-like subsidence depressions (Figure 1c). In contrast to the numerous small steep-sided suffosion and collapse dolines formed atop shallow subcropping shallow evaporites, these larger, bowl-like dolines can have diameters of 100-500 m and depression depths of 10-20 m or more. Most subsidence sag dolines have a well-developed soil cover and a thick sediment fill, with dissolving salt units found depths measured in tens to hundreds of metres below the land surface (Figure 4). Compared to collapse and suffusion dolines, subsidence or sag dolines have lower angle to near flat slopes into the deeper parts of the doline hollow, the doline walls at the outcrop level are usually hosted in non-evaporites, with the dissolving salt bed lying some distance below the landsurface. Compared to suffusion karst, subsidence dolines occur in regions of more deeply buried evaporites (tens of metres) including; Italy (Belloni et at. 1972; Burri 1986; Ferrarase et al., 2002), Spain (Gutiérrez, 1996) and the Pecos Valley of West Texas and New Mexico (Gustavson et al., 1982; Davies, 1984a, b; Quinlan et al., 1986). At the larger end of the scale of subsidence doline development, a subsidence doline merges into a subsidence basin (Figure 1d). The latter is a large solution hollow that had created enough accommodation space to be considered a small sedimentary basin, generally filling with varying combinations of fluvial lacustrine and other continental sediments.


Collapse dolines

Collapse dolines are steep-sided sinkholes, often defined by cave entrances that contain large blocks of roof material (Figure 1c). They form when solution of an underlying evaporite bed creates a roof span that can no longer be supported by the overlying lithology. Collapse doline walls are frequently asymmetrical; one wall is steep, and the other one is gentle (Figure 5). Soil covered doline floors, when not water covered, tend to display either concave-up or flat geometries. Apart from blocks of the collapsed roof span, active doline floors can be veneered by thin collapse breccias in a matrix of insoluble residues. The high density of dolines in areas of evaporite subcrop and the ubiquity of the associated breccias indicates the inherently higher solubility of the evaporite salts compared with interbedded and overlying carbonates. Active collapse sinkholes atop shallow bedded evaporites typify terminations of dry arroyos in many deserts and may also line up along active or former river courses, as in Bottomless Lake State Park, New Mexico (Figure 6).


Sinkhole Classification

Following the definitions of Waltham et al. (2005), Gutiérrez et al. (2008) constructed a nongenetic classification of subsidence dolines or sinkholes (Figure 7). It is based on two observations that refer to the material affected by downward gravitational movements (cover, bedrock or caprock) and the primary type of process involved (collapse, suffosion or sagging). The classification also applies to both evaporite and carbonate karst. The term “cover” refers to allogenic unconsolidated deposits or residual soil material, bedrock to karst rocks and caprock to overlying non-karst rocks. Collapse indicates the brittle deformation of soil or rock material either by brecciation or the downward migration deposits through conduits and its progressive settling, while sagging is the ductile flexure (bending) of sediments caused by the lack of basal support. In practice, more than one material type and several processes can be involved in the generation of sinkholes. These more complex sinkholes can be described using combinations of the proposed terms with the dominant material and/or process followed by the secondary one (e.g. cover and bedrock collapse sinkhole, bedrock sagging and collapse sinkhole).


The cover material may be affected by any of three subsidence mechanisms. The progressive corrosional lowering of the rockhead may cause the gradual settling of the overlying deposits by sagging, producing cover sagging sinkholes. An important applied aspect is that the generation of these sinkholes does not require the formation of cavities These depressions arc commonly shallow, have poorly defined edges and may reach several hundred metres across.

Cover deposits may migrate downward into fissures and conduits developed in the rockhead by action of a wide range of processes collectively designated as suffusion: downwashing of particles by percolating waters, cohesion-less granular flows, viscous gravity flows (non-Newtonian), fall of particles, and sediment-laden water flows. The downward transport of the cover material through pipes and fissures may produce two main types of sinkholes depending on the rheological behaviour of the mantling deposits. Where the cover behaves as a ductile or loose granular material, it may settle gradually as undermining by suffosion progresses. This creates sags and slumps in the overburden materials. Where cover behaves in a more brittle manner, collapse breccias form.

Sinkholes that result from the combination of several subsidence processes and affect more than one type of material are described by combinations of the different terms with the dominant material or process followed by the secondary one (e.g. bedrock sagging and collapse sinkhole). The mechanism of collapse includes any brittle gravitational deformation of cover and bedrock material, such as upward stoping of cavities by roof failure, development of well-defined failure planes and rock brecciation. They define suffosion as the downward migration of cover deposits through dissolutional conduits accompanied by ductile settling. Sagging is the ductile flexure of sediments caused by differential corrosional lowering of the rockhead or interstratal karstification of the soluble bedrock. Sag plays a major role in the generation of sinkholes across broad areas underlain by shallow dissolving evaporites (not included in previous genetic classifications mostly based on carbonate karst). Likewise, collapse processes are more significant in extent and rate in areas underlain by evaporites than in carbonate karst, primarily due to the order of magnitude greater solubility of the evaporites and the lower mechanical strength and ductile rheology of gypsum and salt rocks (Warren, 2016).


Broader-scale evaporite landforms

There are some unique aspects to regional landform and doline density associations above dissolving evaporites because; 1) halite, and to a lesser extent anhydrite, units are many times more soluble than carbonate or siliciclastic strata, and 2) matrix in most salt units away from its retreating edges tends to remain impervious even as salt masses approach the landsurface (Figure 8; melting block of ice analogy, see Warren, 2016). Such features form at the basin-scale in regions where an ancient salt bed is moving from the mesogenetic to the telogenetic realm, especially in regions where mountains are growing adjacent to the uplifted evaporite basin. The broader-scale landscape features not seen in carbonate karst in similar regional settings are; 1) laterally migrating subsidence basins (Figure 8a) and 2) regions of breccia chimneys (Figure 8b).

Subsidence basins

A subsidence basin is a large (tens of kilometres width) solution hollow that creates enough accommodation space to be considered a small sedimentary basin with widths measured in tens of kilometres (Figure 8a). In a continental setting, the basin is filling with varying combinations of fluvial and lacustrine sediments.

Subsidence troughs are large-scale elongate depositional depressions created by interstratal solution along the dissolving edge of shallow dipping ancient uplifted salt bodies. The largest solution-induced depositional basins tend to occur along the margins of the great interstratal halite deposits, creating a solution form that may be represented by a shallow retreating salt slope at the surface, with a laterally migrating monoclinal drape of beds in the vicinity of a salt scarp, which is defined the dissolving edge of the underlying salt bed or beds. Subsidence basins are filled or partly filled by clastic sediments (Olive, 1957; Quinlan, 1978; Simpson, 1988). If the subsidence zones atop a retreating salt edge lack a significant volume of sediment fill, it called is a subsidence trough, as seen in the vicinity of the outcropping edge of the Jurassic Hith Anhydrite in Saudi Arabia (see part 2).

Salt-edge leaching of shallow dipping salt underscores a set of self-perpetuating processes. Fractures created by the collapse of overburden and intrasalt beds contribute to an expanding accommodation space generated by salt cavities and the lateral retreat of the salt scarp. The ongoing salt dissolution provides new fissures and sinks that act as additional conduits for further percolation of meteoric or upwelling of undersaturated basinal waters into the salt. This, in turn, instigates yet more salt solution in the vicinity of the collapse and typically, so creating an elongate corridor of subsidence at the surface, which can extend for many kilometres parallel to the dissolving salt edge

Hence, large solution-induced depositional basins and monoclines outline the edges of the great interstratal salt deposits of the world. Depression corridors some 5-500 km long, 5-250 km wide, with up to 100 to 500 m of subsidence induced relief and sediment fill define saline karst plains along the edges of bedded and dissolving saline giants in the Devonian evaporite subcrop of Canada (Figure 9b; De Mille et al., 1964; Tsui and Cruden 1984; Christiansen and Sauer, 2002; Tozer et al., 2014; Broughton et al., 2017), the Permian Basin of New Mexico, Oklahoma and Texas (Figure 9a; Anderson, 1981; Davies, 1984a, b; Bachman, 1984), the Perm region of Russia (Gorbunova, 1979) and the Jurassic Hith region of Saudi Arabia (Amin and Bankher, 1997a, b;). Associated regolith and sediment infill typically reduces the regional solution edge landscape to a few tens of metres of relief.

Salt Breccia Chimneys

Undersaturated waters not only influence the dissolving edges of a shallow-dipping salt bed, but can also cut up through the salt as a series of salt chimneys, located well out in the salt basin, with diameters between 20 and 250 m. Most are plumb vertical, with reported heights ranging from tens of metres to kilometers. Higher examples have usually stoped upward through one or more cover formations, which may include siliciclastics, coal measures, extrusive volcanics, etc. (Figure 9b). Breccia pipes in an evaporite mass may exhibit one of four possible states (Figure 8b): a) Active, and propagating upwards towards the surface, with fault and collapse bounded edges, but not yet expressed at the surface - a blind chimney; b) Active or inactive, expressed at the surface as a closed depression or a depression with a surface outflow channel; c) Inactive, and buried by later strata - paleokarst chimney; d) Inactive and standing up as a positive relief feature on the landsurface, because the breccia (generally cemented) is more resistant to weathering than the upper cover strata - so forming a residual pipe with positive relief. Some upstanding features are firmly cemented and resistant structure, such as the castiles above an erosional plain in the Delaware Basin of Texas (Hill, 1990)

Downward flexure in the uppermost strata atop a growing pipe tends to form a bowl of subsidence and a sag doline structure.

In some intrasalt pipes, the halite or gypsum is entirely removed, leaving only an accumulation of insolubles and collapsed breccia blocks. In others, portions of the less soluble salts can remain in the pipe at the level of the mother salt. Mature pipes are typically filled by a jumble of intrasalt and suprasalt breccia blocks. In some, a lower brecciated zone is succeeded by an upper zone in which the overlying strata failed as a coherent block, settling downwards via a cylindrical pattern of steep to vertical faults with downthrows of up to 200 m. Such variably filled and zoned salt breccia chimneys have created one type of barren zone in the potash ores of the Prairie Evaporite in Canada, and we shall discuss this in detail in part 5(economic associations).

Most breccia pipes originate via a point-source dissolution breakthrough of a halite bed (occasionally gypsum or anhydrite) and are located above a fracture junction, an anticlinal crest, or a buried reef that channels groundwater into a local area in the salt (Figure 8b). Hence, salt breccia chimneys and pipes form best where there is an artesian head to create subsalt pressures. In the Western Canada Sedimentary Basin the head comes from the adjacent uplifted Rocky Mountains, while the Delaware and Guadalupe Mountains play a similar role in the Delaware Basin. A pipe creates anomalous chemical interfaces with adjacent and overlying strata and so may be targets for the later precipitation of economic ores, including uranium and base metals (part 5).


Delaware Basin, West Texas

The Delaware Basin of West Texas and southeast New Mexico, in the southwest portion of the Permian Basin, contains bedded evaporites of the Late Permian Castile, Salado and Rustler Formations (Figure 9a). Outcrops and subcrops of these three formations constitute an area of widespread subsidence troughs, collapse sinks, dolines and breccia chimneys, all created by ongoing removal of underlying bedded Permian salts. The Delaware is an eastward-dipping basin, mainly surrounded by the Capitan Reef and its equivalents (Figure 9a). The original extent of the evaporites in the basin was much greater than today due to ongoing salt dissolution.

The area called the “Gypsum Plain” of Texas lies to the west of the Pecos River and comprises about 2,600 square kilometres of subcropping gypsum of the Castile Formation (Olive, 1957; Quinlan, 1978; Kirkland and Evans, 1980; Anderson, 1981; Stafford et al., 2008a,b; Holt and Powers, 2010). South of Carlsbad and Carlsbad Caverns, the plain exhibits many small examples of solution subsidence troughs, typically 0.7–15 km in length, 100–1500 m wide, but no more than 5–10 m deep (Figure 9a: Stafford et al., 2008a). The plain is also where the “Castile” landforms are found and indicate an overprint of bacterial and thermochemical sulphate reduction (Kirkland and Evans, 1976). Additional gypsum outcrops are present to the east in the Rustler Hills and to the north in Reeves County, Texas. Permian strata in region of the "Gypsum Plain" once contained significant volumes of halite that were dissolved well before the evaporite succession reached the landsurface.


In the centre of the Delaware Basin, the thick evaporites of the Castile and Salado Formations retain their halite beds as do the thinner evaporites of the overlying Rustler Formation. All are underlain by relatively permeable carbonates and siliciclastics, including some prolific hydrocarbon reservoirs in Permian backreef of the Central Basin Platform (Figure 10). Reef mounds, fractures and faults in these underlying sediments have provided focused conduits for upward stoping breccia chimneys through the buried evaporites as well as the subsurface formation of now-exhumed Castiles to the west. An eastward-flowing deeply-circulating regional artesian hydrology, in combination with centripetal escape of buoyant hydrocarbon-rich basinal waters, drives the formation of these chimneys, with their surface expressions occurring in areas such as the Wink Sink (chain of breached chimneys) and the Gypsum Plain (Castiles).

The upper sides of the shallow dipping salt beds are also affected by the hydrologies of the zone of active phreatic circulation. The Pecos River has migrated back and forth across the top of subcropping evaporites for much of the Tertiary. Its ancestral positions drove substantial salt dissolution, now evidenced by large sediment-filled subsidence basins and troughs in the centre of the Delaware Basin (Figures 9b, 10). The regional eastward dip of the Delaware basin sediments means first halite, and then gypsum has disappeared along the updip eastern edge of the basin. Relatively undisturbed salt remains along the more deeply buried western side of the basin that abuts and covers the Central Basin Platform. Bachman and Johnson (1973) estimate the horizontal migration rate of the dissolution front across the basin as high as 10-12 km per million years so that more than 50% of the original halite is gone. Multiple smaller examples of sediment-filled subsidence troughs occur at the edge of the gypsum plain of the Delaware Basin south of Carlsbad Caverns, where depositional troughs, 0.7 to 15 km in length, 100 to 1,500 m wide and no more than 5-10 m deep, are well documented, as are subsidence sinks within the subsidence swales, such as Bottomless Lake, which is a region where the watertable intersects a collapse chimney (Figure 10c; Quinlan et al. 1986).

The San Simon swale is a 25 km2 depression defining a residual karst feature atop the Capitan Reef on the northeastern margin of the Basin (Figures 9a, 10a, b). San Simon Sink sits atop a subsidence chimney within the San Simon swale; it is the lowest point in the depression and is some 30 m deep and 1 km2 in area. It, in turn, encloses a secondary collapse sink some a few hundred metres across and 10 metres deep (Figure 10b). During a storm in 1918, the San Simon sinkhole formed as a gaping hole about 25 metres across and 20 metres deep in the lower part of the sink. In one night, nearly 23,000 cubic metres of soil and bedrock disappeared into the collapse cavern. Annular rings that cut the surface around the San Simon sinkhole today suggest ongoing subsidence and readjustment of the sinkhole is still occurring in response to earlier collapses. The position of the San Simon sink over the Permian reef crest led Lambert (1983) and many others to suggest that the sinkhole originated as a groundwater cavity breakthrough, atop a series of stoping reef-focused breccia pipes or chimneys. Sinkhole breakouts, which can emerge in a matter of hours, continue to form across this dissolution basin, in some case aided by poorly-monitored brine extraction operations and improperly-cased water wells. But the majority of the sinkholes in the Delaware Basin are natural, not anthropogenically enhanced (Land, 2013).


Nash Draw is a southwesterly trending depression or swale, some 25 km long and 5-15 km wide, at the northern end of the Delaware Basin with its sump in a salt lake (Poker Lake) at the southern end of the draw (Figure 11). The underlying evaporitic Rustler Formation and parts of the Salado Formation have largely dissolved so that more than 100 caves, sinks, fractures, swallow-holes, and tunnels make up a complex local karst topography in the Draw, which is still active today (Figure 11b, c; Bachman, 1981, Powers et al., 2006; Goodbar and Goodbar, 2014).

An extensive drilling program conducted for the nearby WIPP site (now a low-level radioactive waste repository) showed that natural dissolution of halite in the Rustler and upper Salado formations is responsible for the subsidence and overall formation of Nash Draw (Lambert, 1983; Holt and Powers, 2010). To the immediate west of Nash Draw, the WIPP/DOE drilling program defined the formation of a solution trough in the Dewey Lake Redbed; it was created by preferential leaching of halite beds in the Rustler Formation, with interstratal anhydrite and breccia residuals (Figure 12). This dissolution occurred at a level some 400 metres above the salt-encased storage level of the WIPP waste isolation facility and so is not considered a significant risk factor in terms of longterm site stability (Holt and Powers, 2010). A heated scientific (and at times not so scientific!) debate of just how deep surface karst penetrates into the bedded halite of the Salado Formation in the vicinity of the WIPP site continues today.


Further north, near the subcropping western edge of the Northwest Permian Shelf, is Bottomless Lakes State Park, located some 20 kilometres southeast of Roswell. Encircling the lakes are the gypsum, halite and dolomitic redbeds of the Artesia Group and San Andres Formation. Away from the lakes are numerous other sinkholes and collapse dolines, most of which are circular, steep walled or vertical holes, 50-100 metres across and 30-60 metres deep, with the greatest density of features aligned along the eastern side of the Pecos River floodplain (Figure 6). Water in the various sinks that make up the Bottomless Lakes is crystal clear and brackish to saline (6,000-23,000 ppm), attesting to its passage through subsurface layers of gypsum and salt. Although the lakes are around 30 metres deep, dark-green moss and algae coat the bottoms giving the impression of great depth and hence the name of the park (Lea Lake; Figure 10d.). To the west, many of the playas in depressions near Amarillo in the High Plains of Texas have a similar genesis as solution depressions atop dissolving Permian salt, but most do not intersect the regional watertable and so do not hold permanent surface water (Paine, 1994). Further south, near Carlsbad Caverns, there are other subsidence chimneys that form lakes where they intersect the watertable and so are also locally known as “Bottomless” (Figure 9a).


Western Canadian karst

The distribution and timing of chimneys and subsidence troughs created by the subsurface dissolution of the Prairie evaporite are well known and tied to the distribution of oil sands and the quality of potash ores (part 5). Dissolution drape features are more pronounced nearer the retreating edges of the thick multilayered subsurface Devonian salt succession while breccia chimneys occur above the buried salt successions (Figure 9b). For example, the Rosetown Low and the Regina Hummingbird Trough accumulated more than 100 metres of depression trough and drape sediment during interstratal dissolution of the underlying Prairie salt (Devonian) in southern Saskatchewan, especially in Cretaceous time (Figure 13a,b; DeMille et al., 1964; Simpson, 1988). This evaporite-related Cretaceous subsidence is also the principal control on the distribution of the Athabasca oil sands in subsidence basins along the eastern side of the evaporite extent (Tozer et al., 2014). Uplift of the ancestral Rocky Mountains likely created the potentiometric head that drove much of the subsalt aquifer flow. As in the Delaware Basin, the positions of sub-salt reefs and pinnacles focused many of the upwells of deeply circulated meteoric water that ultimately created solution breccia layers and breccia chimneys (Figure 8b).

In various circum-salt subsidence troughs, now filled or partially filled with late Mesozoic and Tertiary sediment thicks, the concurrence of a supra-unconformity thick, adjacent to the sub-unconformity feather-edge of a bedded salt sequence, is at a scale that is easily recognised in seismic and constitutes one of the classic signatures of a salt collapse-induced hydrocarbon trap (Figure 8a). For now, we will focus on the influence of dissolving evaporites on the modern Canadian landscape in regions of active karst, but we will return to the topic of economic hydrocarbon associations with paleokarst in this basin in part 5.

Evaporite karst domes, and laterally extensive solution breccia units, tied to a waxing and waning cover of glacial ice and permafrost, were first documented in Canada in northern Alberta and the Northwest Territories (e.g. De Mille et al., 1964). Karst domes remain as surface features of positive relief once the surrounding evaporite mass has completely dissolved and are outlined by megabreccia with caverns. Unlike breccia chimneys, the cores of evaporitic karst domes can expose blocks from below, as well as above the original bedded and folded evaporite level. The evaporitic karst domes in western Canada are related to the stratigraphic level of former bedded salts; elsewhere in the world others are the remains of now dissolved salt thrusts, diapirs and allochthons (see diapiric breccias and rauhwacke discussion in Warren, 2016). Hence, domes and residual units are dramatic landscape features in the gypsiferous terrain of northern Alberta, Canada (Wigley et al., 1973; Tsui and Cruden, 1984). Similar features, tied to the dissolution of bedded evaporites, typify the Arkhangelsk gypsum-residue karst region inland of the Barents Sea coast of Russia (Korotkov, 1974).

Canadian karst domes range from 10 to 1000 m or more in length or diameter, and can rise to 25 m above the surrounding land surface. Many domes are highly fractured, with individual overburden blocks displaced by heaving and sliding, with the residual gypsum showing well-developed flow foliation.

At the extreme end of disturbance and dissolution range, the domes breccias are megabreccias; positive relief features made up of collapsed or even upthrust jumbles of large blocks, some the size of houses. The largest reported Canadian megabreccia example is in a steep-limbed anticline that extends along the shore of Slave Lake for a distance of 30 km. It is up to 175m in height with a brecciated crestal zone that is marked by a ‘chaotic structure and trench-like lineaments’ (Aitken and Cook 1969).

These brecciated landforms develop upon a distinctive geological association in the central Mackenzie Valley region, NWT, at a stratigraphic level equivalent to anhydrites and halites in the deeper subsurface (Hamilton and Ford, 2002 ). In widespread outcrops the breccia unit is formally named the Bear Rock Formation, when covered by consolidated strata it is termed the Fort Norman Fm (Meijer Drees, 1993; Law, 1971). It is centred in Late Silurian-Early Devonian strata and defines an outcrop and subcrop belt more than 50,000 km2 in extent. In core, the Fort Norman Formation is 250-350 m or more in thickness. It consists of a thin upper limestone (Landry Member), a central Brecciated Member and, in some cores, undisturbed lower sequences of inter­bedded dolostones and anhydrite remnants. It is conformably overlain by 90-150 m of limestones and calcareous shales (Hume Formation-Eifelian). This evaporite dissolution breccia is possibly derived from the dissolution of not one, but several subcropping Devonian salt layers. That is, although not much discussed in the literature, the mother level may not be tied to a single stratigraphic layer.

Typical till-covered collapse and subsidence karst of the Bear Lake Formation can develop through the Hume and higher formations as a consequence of interstratal dissolution of the salt layers in the Fort Norman and equivalents, where meteoric groundwater circulation and sulphate dissolution have been recorded at core depths as great as 900 m (Figure, 14, 15).


In the NWT the Bear Rock Formation is considered to host to much of this widespread solution breccia, which up to 250 m thick. If present, the Landry Member is a brecciated limestone no more than 20 m thick. The main breccia level forms a visually set of outcrops made up of chaotic, vuggy mixtures of limestone, dolomite, and dedolomite clasts, variably cemented by later calcites, with small residual clasts and secondary encrustations of gypsum. Pack breccias in the Bear Creek and its equivalents displaying rubble fabric (predominant), crackle or mosaic fabric, are common and cliff-forming (Stanton, 1966; James & Choquette, 1988). Float breccias are rarer and tend to be recessive. Meijer Drees (1985) classifies the Bear Rock Formation as a late diagenetic solution breccia created by meteoric waters. Hamilton (1995) shows that calcite, dolomite and sulphate dissolution, plus dedolomitisation with calcite precipitation, are continuing today. They are re-working older breccia fabrics, creating new ones, as well as forming a suite of surficial karstic depressions and subsidence troughs ranging from metres to several km in scale.

The Canadian example constitutes an important set of observations that also relate to many other regions with widespread, basin-scale evaporite dissolution breccias, namely; evaporite solution breccias are multistage and can encompass significant time intervals. Similar-appearance evaporite solution breccias can form at different times in a basin's burial history, from different superimposed undersaturated hydrologies. Individual breccias in any single sample are typically responses to multistage, multi-time diagenetic-fluid overprints. This is also why potash ores in the Prairie evaporite preserve evidence of multiple times of potash mobilisation and mineralogy (part 5; Warren, 2016).

Detailed evaporite karst landform studies have focused on Bear Rock Mountain (type area) and the Mackay Range, which are outlying highlands on the east and west banks respectively of the Mackenzie River, and terrain between Carcajou Canyon and Dodo Canyon in the Mackenzie Mountains (Hamilton, 1995). These sites were covered by the Laurentide Ice Sheet (Wisconsinan) but were close to its western, sluggish margin. They are at the boundary between widespread and continuous permafrost in the ground today and can display some year-round groundwater circulation via taliks. Thaw/freeze and solifluction processes compete with ongoing evaporite dissolution to mould the topography.


Regionally, the principal karst landforms hosted in and above the Bear Rock Formation are varieties of sinkholes, blind valleys, solution-subsidence troughs and fault-bounded depressions (e.g. Figures 14, 15). Sinkholes range from single colla­pse features a few metres in diameter to merged or compound dolines up to 1.5 km2 and 100 m deep. Smaller individual sinkholes and collapse dolines may retain seasonal meltwater ponds, and there are permanent lakes draining to marginal ponors in some of the larger subsidence troughs. Blind valleys have developed where modern surface streams flow for several km into the evaporite karst zone from adjoining rocks. There are many relict, wholly dry valleys that may have been created by glacial meltwaters. Subsidence troughs have developed at the surface along contacts between the Bear Rock breccias and underlying, massive dolostones that are typically gently dipping. In the centre of the Mackay Range and at Bear Rock itself are solution depressions formed where the breccias make-up hanging-wall strata on steeply inclined fault planes on the collapse or pipe edge. The Mackay example is 3.2 km long, 1.0 km wide and-160 m deep (Ford, 1998).

Where patches of the Landry Member survive above the main breccia level(s), they are often broken into large, separate slabs that tilt into adjoining depressions in sharply differing directions, creating a very distinctive topography of dissolution draping and block rotation (see also Dahl Hit in Warren, 2016). Some slabs are rotated through 80-90°. Ridges (inter-sinkhole divides) that are wholly within the main breccias often display stronger cementation, represented by pinnacles as much as 30 m high The many sinking streams pass through the permafrost via taliks and emerge as sub-permafrost springs at stratigraphic contacts or topographic low points.

According to Hamilton (1995), the variety and intensity of karst landform assemblages on the Bear Rock Formation are like no other in Canada, and he notes he had not seen or read of very similar intensive karst topographies elsewhere. He attributes their distinctiveness to repeated evaporite dissolution and brecciation, with dedolomitisation and local case hardening, throughout the Tertiary and Quaternary, with these processes occurring multiple times in mountainous terrains subject to episodes of glaciation, permafrost formation and decay, and to vigorous periglacial action.

The spectacular karst domes that typify dolines and glacially associated surfaces atop an evaporite subcrop and are particularly obvious in regions of anticlinal salt-cored structures within regions of widespread permafrost. The accumulation of ground ice in initial fractures in the evaporite layer probably contributes to the heaving, folding and other displacement of breccia fragments in the dome. Tsui and Cruden (1984) attributed the examples that they studied in the salt plains of Wood Buffalo National Park Canada (Lat. 59-60°N) to hydration processes operating on subcropping bedded gypsum during the postglacial period. Ford and Williams (2007) argued such features indicate local injection of gypsum residuals during times of rapidly changing glacial ice loading. Whether they are created by glacial unloading/reworking or are a type of gravitationally-displaced dissolution breccia in a permafrost region is debatable; that in Canada they are a widespread type of evaporite karst residue is not. They characterise those parts of the permafrost-influenced region defined by dissolution of shallow subcropping folded Devonian salts in Northern Alberta, where the evaporites are typically exposed in anticlinal crests are today still retain fractionated gypsum residuals.

And so, in addition to widespread breccias hoisting karst domes, there are numerous natural collapse dolines atop dissolving shallow salt beds. A classic example is the water filled doline some km NE of Norman Wells (Figure 15).


Holbrook Anticline, Arizona

Subsidence driven by natural salt solution at depth generates regional-scale drape or monoclinal folding in strata atop the retreating salt edge. This is the corollary of the formation of a subsidence trough (Figure 16a). The 70-km long dissolution front in the Permian Supai salt of Arizona is defined by more than 500 sinkholes, fissures, chimneys and subsiding depressions some 40-50m across and 20-30m deep (Neal, 1995; Johnson, 2005). Away from the main dissolution zone the northeasterly-dipping Schnebly Hill Formation (aka Supai Salt) is composed of up to 150 metres of bedded halite, with local areas of sylvite along its northern extent (Figure 16a). Atop the solution front, there are a number of topographic depressions with playa lakes that in total cover some 300 km2. Salt-dissolution induced features to include areas known as; The Sinks, Dry Lake Valley, and the McCauley Sinks (Neal et al. 1998; Martinez et al. 1998). The Sinks region includes more than 20 steep-sided caprock sinkholes, with some that are more than 100 metres across and up to 30 metres deep (Figure 16b). The McCauley Sinks are the likely surface expression of compound breccia pipes and chimneys (Neal and Johnson, 2002).

Regional expanses of Supai Salt removal produce a regional gravity depression, largely coincident with a surface topographic depression, where there is as much as 100 metres of collapse and topographic displacement. The solution front is essentially coincident with the updip end of the Holbrook Anticline, a flexure defined by dip reversal in an otherwise northeasterly dipping succession (Figure 16a).

Rather than orogenically driven, the Holbrook Anticline is a subsidence-induced monoclinal flexure created by the northeasterly migrating dissolution front (Figure 16a). It may be the largest single solution-collapse fold structure in the world (Neal, 1995). The reverse dip of the flexure directly overlies the salt-dissolution front and marks the location of two major collapse depressions known as The Sinks and Dry Lake Valley, both occur where salt is within 300 m of the surface (Peirce, 1981). Although it has periodically held surface water, reports of several hundred acre-feet of flood water in the Dry Lake Valley playa draining overnight supports the notion of active fissure and cavern formation related to salt removal at depth. Major surface drainage events took place in 1963, 1979, 1984 and 1995, with more than 50 new sinkholes forming in the valley during that period. The continuing rapid appearance of new sinkholes testifies to the ongoing nature of dissolution in the underlying evaporites. According to Neal (1995), dissolution front features began forming in the landscape in the Pliocene and continue to form today.

The caprock collapse sinkholes centre in the Coconino Sandstone bed that overlie the salt, which is located 200-300 metres below the at-surface features. These caprock sinkholes define regions of focused breccia pipe development and discreet upward cavity migration (chimney stoping). The underlying salt is bedded and there are no indications of halokinesis anywhere in the basin. Sinkhole regions lie just ahead of a dissolution front that is migrating downdip, driven by the widespread dissolution of halite.

Other karst features attributed to evaporite dissolution in the Holbrook basin are; pull-apart fissures, graben sinks (downward-dropped blocks), breccia pipes and plugs, and numerous small depressions with and without sinkholes (Neal et al., 1998, 2001). Interestingly, many of the “karst” features occur in sandstones, not limestones; such caprock collapse sinkholes can form wherever pervasive dissolution has removed the underlying salt, independent of caprock lithology. The presence of the more than 500 karst features in the Holbrook basin, some of which formed in days, evokes practical karst hazard and infrastructure concerns, even in such a sparsely populated region (Martinez et al., 1998).

Implications

Evaporite-related karst landforms are in many ways similar carbonate karst features. But, the much higher solubilities of halite and other evaporite salts compared to limestone and dolomite means there are additional features unique to regions of salt dissolution.

In the subsurface, a bedded evaporite can be composed of thick impervious halite with intrasalt beds composed of anhydrite dolomite and calcite. The edges of this halite bed can dissolve even in the deep mesogenetic (burial) realm, wherever the edge of the salt is in contact with undersaturated waters. Insoluble residues bands start to form and can take the form of a basal anhydrite or a fractionated caprock. If the rise of undersaturated water is focused, a breccia cavern can form, and once it breaches the salt, the cavity will contain collapse blocks of the less soluble intrasalt beds. The cavity can then stope to the surface, forming a breccia pipe. A transtratal breccia pipe can rise through kilometres of overburden before attaining the surface. There is no equivalent process-response in the carbonate realm.Subsidence basins, troughs and megabreccia plains are also features that owe their origins to the rapid dissolution rates inherent to salt bed-fresher water contacts

In a similar fashion, the rapid dissolution of halite in the shallower parts of a salt basin means the evaporite karst features will transition from mesogenetic or bathyphreatic cavern formation to meteoric phreatic to vadose effects. This is the emphasis of the second article in this series, which will discuss processes forming vadose and phreatic caves in evaporites.

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Lithium in saline settings

John Warren - Sunday, July 30, 2017

 

Introduction

Historically, until the electronic revolution, lithium located near the top of the periodic table, was of little economic interest. With an atomic number of 3 and an atomic weight of 6.9, lithium is exceptionally small and light, with a high charge/radius ratio. Estimates of the average crustal abundance of lithium vary, but it is likely to be approximately 17–20 parts per million (ppm). In igneous rocks, the abundance is typically 28–30 ppm, but in sedimentary rocks, it can be as high as 53–60 ppm (Evans, 2014; Kunasz, 2006).Lithium-prone hypersaline brines can attain values as high as 6000 ppm, but such high levels are unusual.

Lithium compounds are consumed in the production of ceramics, lubricants, glass, and primary aluminium (Figure 1). Its high specific heat capacity makes lithium ideal in heat transfer technology where it is used in welding and metallurgical applications. Its light weight and its high electrochemical potential (it is most electronegative metal known) and its high electrical conductivity make it amenable to battery applications requiring low weight and high storage potential.

Compared to nickel metal hydride batteries, the type of battery currently powering most hybrid electric vehicles, Li-ion batteries are lighter, less bulky, and more energy efficient. Lithium batteries have three times the energy of nickel hydride at one-third of the weight, and they operate at very low temperatures with a longer battery life. The use of Li-ion batteries in electric cars and electronic devices has increased the global demand for lithium, a trend that is likely to continue. Currently, lithium‐rich saline brines are the most economically recoverable Li source on the planet. (Kesler et al., 2012; Grosjean et al., 2012).


Economically recoverable forms of lithium occur in four types of deposits (Table 1; Figure 1)

(1) Pegmatites,

(2) Continental saline brines

(3) Hydrothermally altered saline lacustrine clays,

(4) Geothermal and basinal brines.

Figure 2 plots know occurrences of lithium in saline deposits. There are four main clusters of hypersaline brine-related lithium occurrences, 1) Andean Altiplano, 2) Tibetan Highlands, 3) Qaidam Basin and 4) Playa brines in the Basin and Range of the south-west USA. Then there are the lesser volumes, as yet un-economic lithium accumulations, associated with lacustrine clays near Hector, California and in the Jadar Valley, Serbia. Basinal (oilfield) brines are known to entrain elevated levels of lithium in the Smackover Fm, USA and the Fox Creek region of Canada. Geothermal brines below the Salton Sea can also contain elevated levels of lithium.


Lithium‐rich continental brine sources account for about three‐fourths of the current world lithium production with the remainder from pegmatites (Figure 3; U.S. Geological Survey, 2017). Only the geology of hypersaline brine sources and associated saline sediment hosts, basin brines and clay replacements are discussed in detail in this article. For information on pegmatites and oilfield brines the interested reader is referred to Garrett (2004), Kesler et al.(2012) and Evans (2014). All three articles contain a broader discussion of occurrences of lithium raw materials and their processing.

In natural brines, most lithium salts are highly soluble and tend to stay in solution until lithium concentrations approach and exceed 6000 ppm. Lithium can be to be absorbed by saline playa clays at lower concentrations, as in the hectorite beds in Clayton Playa, Nevada or hydrothermally from hypersaline saline waters as formed jadarite nodules in the Jadar Valley, Serbia. Actual lithium carbonate precipitates are highly soluble and so very rare in sedimentary basins; lithium carbonate (zabuyelite) is a natural precipitate from the high-altitude hypersaline waters of Lake Zabuye on the Tibetan Plateau.


Lithium carbonate brines

Production from hypersaline pore brines in South American salars dominates current world lithium brine production, with Chile and Argentina producing some two-thirds of the documented world brine production (Figure 2). Chile has emerged as the largest lithium carbonate producer from a lake brine, largely through the exploitation of brines in Salar de Atacama. China and Argentina are the other main producers of lithium from saline lake brines (Figure 3).

Lithium in salar brines of the Andes

Salar de Atacama lies on the Tropic of Capricorn at an altitude of 2,300 m in the Desierto de Atacama, some 200 km inland from Antofagasta. In its more central portions this salt-encrusted playa contains a massive halite unit (nucleus) that is more than 900 m thick, with an area ≈ 1,100 km2. Fringing saline muds, with an area ≈ 2,000 km2, surround the nucleus (Figure 4a, b).

The current salt crust atop this halite nucleus contains a sodium chloride interstitial brine that is rich in Mg, K, Li, and B ((Figure 4c; Figure 5; Alonso and Risacher, 1996). Lithium contents of the pore brine range from 200-300 ppm in the marginal zone, some 500-1,600 ppm in the intermediate zone and 1,510-6,400 ppm in the salt nucleus). The nucleus zone averages 4,000 ppm lithium and is asymmetric with respect to the salar centre due to the sump offset via ongoing faulting. Main inflows to the salar drain volcanic formations of the Andean Highlands located to the east of the basin.


Salts dissolved in inflow waters have a double origin. Weathering of volcanic rocks supplies K, Li, Mg, B and, to a lesser extent, Na and Ca. Leaching of ancient halokinetic evaporites sourced in a mother salt layer beneath and piercing the volcanic formations provides additional amounts of Na, Ca, Cl, SO4 to the most saline inflow waters. The mass-balance of the upper halite nucleus in the salar shows a strong excess of NaCl with respect to the bittern solutes Mg, K, Li, B. According to Alonso and Risacher (1996), this suggests that the nucleus did not originate from evaporation of inflow waters similar to the present groundwaters. Rather, the excess of NaCl is due to NaCl-rich inflow waters that formerly drained the Cordillera de la Sal, a Tertiary-age evaporitic ridge along the western rim of the present-day salar (Figure 4b).


Although annual salt accretion rates in a salar salt nucleus facies in an Andean salar can be as high as 5-6 cm/year (Ruch et al., 2012), the average sedimentation rate of halite in the Atacama lake centre is ≈ 0.1 mm/year, based on the age of an ignimbrite interbedded with the salt. This slow aggradation rate implies a climatic setting of long dry periods and inactivity alternating with short wet periods during which large amounts of water, and so large amounts of salt are first recycled and then accumulate in the halite crusts of the basin sump. The lack of peripheral lacustrine deposits and the high purity of the Atacama salt also suggest that the main salt unit is not the remnant of an ancient deep saline lake, but originated mostly from evaporation of waters supplied by subsurface and subterranean saline seeps.


Once Li-rich lake brines are pumped to the surfac,e they flow into a series of evaporation pans where three main economic products (halite, potash and lithium salts) are recovered. To achieve this controlled salt-series production, the brines are first pumped from 30 metre deep boreholes that penetrate the porous salt nucleus layer into a series of solar evaporation ponds (Figure 6a). Over the successive passage through the concentrator pons, liquors are concentrated by a factor of 25, generating a final brine strength of 4.3% Li (Figure 6b). During evaporation and processing for production of halite, potassium chloride and potassium sulphate from lithium precursor brines, the ion ratios are continuously monitored and adjusted to avoid the precipitation of a lithium potassium sulphate salt. This combination of solar concentration and brine processing, proceeds as follows:

1) Sodium chloride (common salt) precipitates first. If required, this salt can be scrape-harvested as a by-product.

2) At the appropriate level of concentration, the brine is transferred to a second set of ponds in which a mixture of sodium chloride (salt) and potassium chloride (potash, in the form of sylvinite) is precipitated. These salts usually are harvested and the two components separated in a flotation plant.

3) The remaining brine is piped to another set of evaporation ponds where it remains until the concentration increases to 6000 ppm Li (essentially the saturation point of lithium chloride - saturated brines typically show a green colour as visible in Figure 6b). Ripe brine is then transferred to a recovery plant where impurities such as magnesium and boron are removed. When soda ash (sodium carbonate) is added to the ripe brines, lithium carbonate drops out. Brine with low magnesium levels is the preferred feed brine as this makes for simpler processing.

The high initial lithium content of the Atacama brines and the extremely arid setting (3200 mm pan evaporation and <15 mm precipitation) means that only 90 hectares of evaporation ponds are required in one of the current brine operations on the salar, this is only 5% of the area required at Clayton Valley, Nevada with its milder climate and lower Li concentration in the feeder brine (Figure 7). Borate (as perborate) is recovered at levels of 0.84 g/l during lithium extraction at Atacama. Increasing volumes of lithium are also produced by new salar brine processing facilities in nearby Salar de Hombre Muerto, Argentina and Salar de Uyuni, Bolivia. All these salars have lower levels of Li in the primary brine feed than Atacama.

Lithium brines in the USA

Clayton Valley is host to the only commercially producing lithium project in North America, Albemarle’s Silver Peak brine evaporation pond project (Figure 7). Historically, the Clayton Valley playa produced about one-third of the US lithium requirements, but its economic viability suffered from fierce market competition, especially from South America, and a largely depleted brine supply. Originally, the central valley area contained 100–800 ppm Li, and the discovery well at 229 m depth contained 678 ppm when pumped at 450 gpm (Garrett, 2004). The average brine analysis when commercial production of lithium carbonate began in 1966 was about 400 ppm (Figure 7). Since that time the feed concentration of lithium has been slowly declining, and in 1998 the concentration was about 100–300 ppm Li (averaging 160 ppm, Harben and Edwards, 1998).


The Silver Peak Playa has an area of 50 km2 and an elevation of 1300- 1400 m (Figure 7). It lies in the rain-shadow of the Sierra Nevada, with an annual rainfall ≈130 mm and an evaporation rate of ≈1380 mm. Near-surface sediments consist of a mixture of clays (smectite, illite, chlorite, kaolin) and salts (halite and gypsum) and widespread pedogenic calcite. Lithium in the brines is derived from weathering and leaching of volcaniclastics in the Tertiary Esmeralda Formation and Quaternary ash-fall tuffs (Davis et al., 1986). Lithium content is highest on the eastern side of the playa adjacent to the outcropping marls of the Esmeralda Fm. Before it is leached, lithium is held in the clay fraction of the playa sediments and is probably part of the clay structure (hectorite is a widespread but minor component in the Clayton Valley clays - see later)).

Lithium-rich brine feed to the plant averages 0.023% (230 ppm) lithium in a background NaCl concentration of 200,000 ppm, is pumped from depths of 100-300m in the Clayton Valley (Silver Peak) playa via a number of gravel-packed wells. The lithium (and potassium) in the deposit probably originated from hot springs along the Silver Peak Fault, with the current brine composition being a blend of evaporated water from these springs and surface and ground water that drains into the basin (Garrett, 2004). Modern saline spring outflows contain 9280–10,000 ppm Na, 786–826 ppm K and 24–43 ppm Li. Unusually high brine temperatures in some areas of the deposit (up to 44°C at fairly shallow depths ≈ 25m) tend to support a volcanic/geothermal origin for the lithium. Some of the brine feeder wells show elevated levels of radon gas.

Pumped brine progresses through a series of fractionating evaporation ponds (Figure 7; Zampirro, 2004). Lithium concentration in the liquor increases to 6,000 ppm over the course of 12 to 18 months in the solar evaporation pans. When the lithium chloride level reaches optimum concentration, the liquor is pumped to a recovery plant and treated with soda ash to precipitate lithium carbonate, which is then removed by filtration, dried, and shipped.

Lithium from brine, when the Clayton Valley first produced product in the 1970s, was considered a unique deposit. Its operations established the technology and economic viability of lithium recovery from saline brine, which led to the development of brine production from the salars of South America that now dominate world production of lithium from brine.

Lithium brine in Chinese salt lakes: Zabuye (Zhabei) and Qaidam basins

The lithium brine resource of China is mostly stored in two saline lake regions in high altitude zones, Lake Zabuye region in the Alpine tundra climate zone on the Tibetan Plateau and four salt lakes in the cold arid steppe climate region of the Qaidam Basin on the Mongolian Plateau (Figure 2). Something like 80% of brine lithium resource found in China is contained in the four salt lakes of the Qiadam: Bieletan, DongTaijinaier, XiTaijinaier, and Yiliping (Figure 8; Yu et al., 2013). Zabuye lake on the Tibetan Plateau is probably the most geologically interesting as the Li content of the lake waters are so elevated that it is the only known lacustrine location where lithium carbonate, zabuyelite, is a natural brine precipitate (Figure 9. Nie et al., 2009; Gao et al., 2012).


Qaidam Lakes

Detailed sedimentological and hydrological work in the Qaidam by Yu et al. (2013) has shown that: (1) Some 748.8 tonnes of lithium is transported annually into the lower catchment of the four salt lakes via the Hongshui-Nalinggele river (H-N river in Figure 8), which is the largest river draining into the Qaidam Basin, (2) Li-rich brines are formed only in those salt lakes in the Qaidam that are associated with inflowing rivers with Li concentrations greater than 0.4 mg/l, and (3) the water's Li concentration is positively correlated with elevated levels in both the inflowing river and the associated subsurface brine. Their findings show that long-term input of Li+ from the Hongshui-Nalinggele river controls the formation of lithium brine deposits. They conclude that the source of the lithium in the lake brines is ultimately from hydrothermal fields, where two active faults converge in the upper reach of the Hongshui River. These hydrothermal fields are associated with a magmatic heat source, as suggested by the high Li+ and As3+ content water in geysers in the geothermal field. Based on the assumption of a constant rate of lithium influx, they estimate that the total reserves of lithium in the Qaidam were likely formed since the postglacial period.

Field mapping and coring indicate that lithium reserves in each of the four salt lakes depend on the influx of Li+-bearing water from the H-N River. The data also suggest that during the progradation of the alluvial Fan I, the Hongshui-Nalinggele drained mostly into the Bieletan salt lake, until the Taijinaier River shifted its watercourse to the north and began to feed the salt lakes of the DongTaijinaier, XiTaijinaier and Yiliping salt lakes, while also driving Fan II progradation (Figure 8).

One of the You et al. (2013) major findings in terms of lithium enrichment models is the importance of the contrasting hydroclimatic conditions between the high mountains containing ice caps and the terminal salt lakes. The greater than 4000 m of relief in the watershed enables a massive amount of ions, such as K+, to be weathered and transported, together with detrital material from the extensive, relatively wet alpine regions to the concentration sumps in hyperarid terminal salt lakes, where intense evaporation rapidly enriches the lake water, resulting in evaporite deposition and associated K- and Li-rich brines. It is no surprise that a saline lake at the foot of the nearby Golmud River fan is one of the few places in the modern world where carnallitite is found (Casas and Lowenstein, 1992).


Lake Zabuye

Lake Zabuye is located some 1000 km west of Lhasa, the Tibetan capital, and lies in the ET Köppen high altitude climate zone of the Tibetan Plateau (Figures 2, 9). The lake is perennial, and water levels can vary by metres each year; in 2008 the water level was some 4422 m above sea level. At this level, the lake’s area is approximately 247 km2. Salinity varies from 360 to 440 ‰, depending on seasonal differences in water input and evaporation rate. The volume of lithium product at the lake is currently limited by the sulphate-rich nature of the primary lake brine, prior to concentration in solar pans (Gao et al., 2012).

When concentrated, the crystallisation sequence of salts from highly concentrated Zabuye lake brine at 25°C is (Figure 10a; Nie et al., 2009):

halite (NaCl) --> aphthitalite (3K2SO4•Na2SO4) --> zabuyelite (Li2CO3) --> sylvite (KCl) --> trona (Na2CO3•NaHCO3•2H2O) and thermonatrite (Na2CO3•H2O)

The lake’s brine is naturally supersaturated with NaCl and other salts, so millions of metric tons of halite, potash, trona, and other minerals have accumulated on the bottom of the lake in the past few thousand years (Zheng and Liu, 2010). Lithium carbonate and sylvite precipitate, via a combination of brine concentration and cooling, and higher levels of lithium carbonate precipitation in the end brine can be induced by the addition of soda ash, as is done in the South American salars (Figure 10b, c).


The problem with the natural lake chemistry of the Zabuye salt lake is that a lithium sulphate salt Li2SO4.3Na2SO4. 12H2O precipitates naturally in the early stages of the low-temperature evaporation process, so reducing the levels of lithium carbonate in the end-stage brines. If the brine concentration series in the pans can be artificially held at mirabilite concentration, then the amount of lithium lost to the sulphate salt is reduced, so levels of lithium in the end-stage brines improve (Gao et al., 2012).

Zabuye Lake is of significant economic value as it is a new type of exploited saline lacustrine deposit (compared to the salars of South America) in that contains it precipitates lithium and borate salts in addition to significant volumes of potash, halite, natron and Glauber’s salt. Lake waters also retain elevated levels of caesium, rubidium and bromine.

Lithium in minerals soaked in saline brines

Two saline minerals in sedimentary basins known to have significant lithium contents are hectorite and jadarite. Hectorite [Na0.33(Mg,Li)3Si4O10(F,OH)2] is a clay mineral of the smectite group, where the replacement of aluminium by lithium and magnesium is essentially complete. It has a lithium content of more than 1%, a hardness of 1–2 on Mohs scale, and a density of 2–3 kg/m3. To date, an economically viable technology for extracting lithium from hectorite, rather than from brines that enclose some of these clay deposits, has yet to be developed (Evans, 2014). Jadarite [LiNaB3SiO7(OH)], is a newly recognised mineral with up to 5.7% Li and 14.7% B. Jadarite is a white porcellanous borosilicate mineral with a Moh hardness of 4-5, and a density of 2.45 gm/cc. It is associated with borate salts such as colemanite in the Oligocene-Pliocene lacustrine host sediments in its type area in the Jadar Valley in Serbia (Stanley 2007). Hectorite is probably associated with crossflows of moderate salinity hydrothermal waters, while jadarite requires a bath of hypersaline hydrothermal waters to form.

Hectorite

Hectorite has a soft, greasy texture, a candlewax-like appearance and feels like modelling clay when squeezed between the fingers. As a colloid, hectorite’s unique thixotropic properties for emulsion stabilising, gelling, suspending, binding, bodying and disintegrating, means it sells for more than US$2,000 a ton, generally as a lubricant to the oil and gas industry. Associated authigenic clays include stevensite and saponite, and in its type area at Hector California lies adjacent to a colemanite deposit.

Hectorite is mined periodically (not as a lithium source) in its type area, the Hector Mine, near Barstow, California. There, hectorite is the main clay mineral in a sequence of altered volcanic ash beds that are interbedded with lake sediments and travertines along an 8 km fault zone (Figure 11; Ames et al., 1958). The hectorite is thought to have formed through hydrothermal alteration of the ash by saline fluids moving up the fault zone (Sweet, 1980). Lithium-bearing volcanic rocks that probably formed in the same way have also been described from Arizona, and the Clayton and King Valleys in Nevada (Brenner-Tourtelot and Glanzman, 1978; Kesler et al., 2012). Hectorite is not considered to be a prime lithium resource in any of these occurrences. It is, however, considered of co-indicator of the former, or current, presence of Li-rich saline brines and as such is considered a pointer mineral to a possible lithium brine resource.


Hectorite is thought to be a result of the combination of three distinct geological processes: 1) the alteration of volcanic ash or glass; 2) precipitation of authigenic phases from saline lacustrine pore waters; and/or 3) the incorporation of lithium into existing smectite clay deposits (Asher-Bolinder, 1991). To form hectorite, all three processes require an arid environment and are associated with lithium-enriched saline alkaline waters, volcanic rocks and hot springs that can also co-precipitate travertines and fine-grained amorphous silica (Zientek & Orris, 2005).

The same association of processes explains the lithium-rich hectorite clays in King Valley (Nevada Lithium prospects) Nevada. There, layers of hectorite occur in a sequence of sedimentary and tuffaceous rocks in moat sediments along the western side of the McDermitt caldera (Figure 12; Kesler et al., 2012). Volcanic activity at the McDermitt caldera complex has yielded precise 40Ar/39Ar ages of 16.5 to 16.1 million years ago and was characterised by extrusion of early metaluminous and peralkaline rhyolite, followed by the eruption of a voluminous ignimbrite with peralkaline rhyolite to metaluminous dacite compositions (Carew and Rossi, 2016). After collapse, the central part of the caldera complex was the site of resurgence, and a moat-like lake formed between this resurgent dome and the caldera walls. The lake was the site of deposition of volcaniclastic sediments that now form a nearly continuous ring within the caldera and host the various hectorite lenses(Figure 11).


Hectorite layers ranging from 1 to 90 m in thickness and have been recognised over a length of about 20 km. Individual layers or groups of layers extending for several km and are annotated as stage 1-5 lenses. The Stage 1 lens of the Lithium Nevada deposit (informally known as the King Valley deposit) has proven and probable reserves of 50 million tonnes, averaging 0.312% Li (Carew and Rossi, 2016). As in the type area in the Hector Mine in California, hectorite in the various lenses is the main lithium-bearing clay mineral in a sequence of altered volcanic ash beds. These ash beds are interbedded with saline lake sediments and travertines, and are hosted in the sedimentary moat facies adjacent to an 8 km fault zone. That is the hectorite formed through hydrothermal alteration of volcaniclastic ash in regions where moderately saline hydrothermal fluids moved up a fault zone.

Hectorite clays are also found in the Sonora Lithium Project, 11 km south of Bacadehuachi in the Sonora state of north-west Mexico. The resource statement, in an April 2016 report, lists 839,000 tonnes of contained lithium in the indicated category and a further 515,000 tonnes in the inferred category, within two distinct lacustrine clay units situated below basaltic caprocks (Pittuck and Lepley, 2016). A pre-feasibility study has been completed, which proposes a two-phase open-pit mine with lithium carbonate processing facility and a mine life of 20 years. A pilot plant has also been constructed, and discussions have commenced regarding possible off-take agreements.

None of these hectorite occurrences are currently mined as a lithium resource.

Jadarite

Jadarite was discovered in 2007 by Rio Tinto and the Jadar deposit, near the town of Loznica, and at that time was estimated to contain an inferred resource of 125.3 million tonnes at a weighted average of 1.8% Li2O, in addition to an inferred resource of boron minerals. Jadarite has so far only been identified in significant amounts within the 20-km long Jadar Basin of Serbia. The Jadar Basin entrains oil shales, dolomicrites, pyroclastic sediments and evaporites which are believed to have accumulated in an intermontane lacustrine environment.


The jadarite occurs both in massive form, several metres thick, and also as small nodules within a fine-grained carbonate matrix (Stanley, et al., 2007). At the main Jadar deposit, a layer containing nodular colemanite (Ca2B6O115H2O) overlies three separate layers or lenses containing jadarite LiNaB3SiO7(OH). Jadarite likely formed via a hydrothermally-facilitated interaction between saline brine and clastic/evaporitic sediment, either in a tuffaceous or clay host (Kesler et al., 2012).

In May 2017, Rio Tinto announced that the Jadar area contains one of the largest lithium deposits in the world, lifting their estimate for Lower Jadar's deposits to 138 million tonnes. Extraction is scheduled to begin in 2023, with a projected underground exploitability of 50 years. As of June 2017, construction of a mine has not begun. A jadarite processing plant is also planned next to the mines, that plant will process the ore into lithium carbonate and boric acid.


Summary

Characteristics that appear to be essential to define a potential lithium carbonate brine resource are; i) an arid climate and, ii) a closed, tectonically active basin, with significant elevation and tectonically activity, which can entrain brines with elevated lithium contents (Figure 15; Bradley et al., 2013; Yu et al., 2013; Warren, 2010, 2016). Sources of lithium can be deeply circulated magmatic or recycled basinal fluids. Magnesium levels in the brine should not be too high as this complicates brine processing during lithium carbonate extraction. A co-occurring potash resource, extractable from the same brine, if present can reduce processing costs.

Another possible requirement—or at least a favourable characteristic—is elevated heat flow, as evident from young volcanoes or hot springs and the associated increase in Li-rich juvenile waters flushing the surrounding drainage basin, as is occurring beneath the Andean Altiplano. Volcanogenic source rocks in the lake drainage, such as felsic, vitric tuffs that have abundant and readily leached lithium are favourable, but perhaps not essential, since lithium is present in most crustal rocks at tens of parts per million. Worldwide all the exploited salt lakes have lithium levels in their lake brines that are well above typical (Figure 14).

Another possible favourable indication of a lithium brine is the existence of hectorite or jadarite in associated clays in the bajada rims.

All known and potential lithium brine deposits are located in arid tectonically active areas, typically in subduction or collision belts with deep-faulted suture systems (Figure 2). At the world scale, lithium-prone saline deposits are latitudinally restricted to cool arid Koeppen climate belts within endorheic brine sumps surrounded by high altitude drainage basins (Figures 2, 15b; Warren, 2010). Borates as evaporite salts are generally tied to the same setting (Warren, 2010).


Active faulting appears to be involved in forming a suitable spring-fed hydrology for all known economic lithium brine basins. Fault-related subsidence also creates accommodation space, without which only a thin veneer of arid basin sediments and brines can accumulate. Thus, a thick basin fill is needed to provide an aquifer of sufficient volume to hold a viable lithium brine resource.

In contrast, saline lakes atop shallow, superficial basins in intracratonic regions such as the Sahara Desert and most inland Australian deserts largely lack active fault control and associated rapid subsidence, and are not known to be prospective for lithium brines.

Salt fills in some lithium basin lacustrine sumps are cut by active intrabasinal faults (known from boreholes and seismic) but have no surface expression due to rapid infill and levelling of the accommodation space by salt precipitation. Significantly, the brine pools in Clayton Valley, Salar de Atacama and the Qaidam sumps are localized along active intrabasinal faults, which also control the distribution of aquifers and influence groundwater movement patterns, as well as the position of maximum stacking of concentrates and brines in the halite nucleus, along with porosity retention levels in the subsurface halite host (Zampirro, 2004; Jordan and others, 2002).

Porosity levels in a host halite aquifer are a major constraint on the potential economics of any salar or salt lake lithium brine resource. Most halite units lose their effective porosity and permeability by depths of 50-60 metres (Warren, 2016; Chapter 1). Thus, most Quaternary lithium brine operations hosted in a halite bed/aquifer will have an economic basement to brine recovery at around this depth. It is unlikely that recovery operations in Salar de Atacama and planned projects in Salar de Uyuni can recover economic brine volumes at much greater depths. There may be a 900m thick halite-dominant succession infilling Salar de Atacama and a number of halite beds to a depth of 120m in Salar de Uyuni, but economic porosities in the halite will likely only be present in the upper portions of the halite fill in both salars.

Postulating likely lithium resources, in a salar of a salt lake, to depths greater than 50-60m should only be done after salar-hydrology-aware drilling has established the presence of economic permeabilities in the hosting halite aquifer and this is likely related to the presence of active faults-. Such measurements require drilling and sampling equipment that facilitates reliable “in-situ” determinations of porosity and permeability in the halite mass and Li measurements in the brine that are related to actual content at the level of measurement, with minimal contamination by waters from outside the measured horizon.

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Aeolian Gypsum and Saline Pans - an indicator of climate change

John Warren - Friday, June 30, 2017

Introduction

Evaporites deposited as aeolian dunes are not commonplace in Quaternary successions and not yet documented in any pre-Quaternary succession (Table 1). These eolian deposits are deposited above the water table in a vadose setting, generally in a degrading playa or salt lake hydrology. Consequently, there is an inherent low preservation potential for this style of evaporite; most documented examples are less than a few tens of thousands of years old.

 

Even though relatively rare as an evaporite type, the presence of eolian evaporites, usually as gypsum dunes or lunettes with associated soils and saline mudflats, does indicate particular climatic and hydrological conditions. Eolian gypsum deposits may have possible counterparts in the Martian landscape (Szynkiewicz et el., 2010).

Over the Quaternary and across the Australian continental interior, increased aridity is expressed by episodes of dune reactivation, lake basin deflation with eroded sediment accumulating downwind in transverse dunes or lunettes (Bowler, 1973; Fitzsimmons et al., 2007), Deposition is tied to increased dust mobility (Hesse and McTainsh, 2003) and reduced river discharge and channel size (Nanson et al., 1995). Such responses to increasing landscape aridity in saline groundwater sumps are seen in most arid to semi-arid regions of the world where water tables are falling, usually driven by increasing aridity.

This article focuses on eroded subaerial evaporites as a response to increasing aridity, especially the formation of gypsum dunes and lunettes (Table 1; Figure 1).


Gypsum dune styles and saline pans

Figure 1 and Table 1 plot documented occurrences of eolian gypsum across the world, overlain on a Koeppen climate base (Figure 1a). Figure 1b plots the latitudinal occurrences of documented gypsum dunes versus elevation and Koppen climate type. Figures 1c and 1d plot the detail of these same occurrences for the USA and Australia, where individual deposits are better documented. At the worldscale, there is an obvious tie to the world's desert belts with occurrences consistently situated in regions of the cool dry descending cells of northern and southern hemisphere Hadley cells (positions indicated by light blue rectangles in Figure 1b - See also Salty Matters article from Jan. 31, 2017). Many occurrences are also situated in Late Pleistocene to Holocene climate transition zones, marked by aridification at the transition from Late Pleistocene to Holocene climates, and in many case tied to transitions from perennial saline lakes and mega-lakes to continental saltflats to dunes and interdunal pans, An example of a quartz sand erg association (downwind of a gypsiferous strandzone) is seen in the transition area into the southern Kallakoopah Pans from the northern margin of Lake Eyre, Australia and its megalake precursor (Figure 2).


At the local scale, gypsum dunes generally occur downwind or atop a saline pan or playa that is, or was, recently subject to a lowering of its lacustrine watertable. In many situations the elongation of individual pan shapes line up in an orthogonal direction to the dominant wind and so also show an eolian control, like the associated gypsum dune position and alignment (Figure 3). Wind-aligned lakes and sumps and oriented-pans are much more numerous with a broader climatic range than gypsum dunes (Goudie and Wells, 1995; Goudie et al., 2016). When present, eolian bedforms associated with oriented pans lacking evaporites are dominated by clay pellets or quartz sand.


Many of the pan edge dunes show crescent shapes and so are termed lunettes. (Figure 3; Bowler, 1973). Lunette sediments range in composition from quartz-rich to sandy clay, gypsiferous clay to nearly pure gypsum. Pure quartz dune lunettes likely formed under lake-full conditions, and so show a distinct hydrology from that of the clay pellet or gypsum-rich varieties, which form by deflation of subaerially-exposed adjacent lake floors. The flocculation of suspended clays into pellets requires some degree of salinity but is less than that required to precipitate gypsum.

Lunette sediments range in composition from quartz-rich, sandy clay, through gypseous clay to nearly pure gypsum. Pure quartz dunes formed under lake-full conditions and are distinct from that of the clay and gypsum-rich varieties, which formed by flocculation and deflation from adjacent subaerially exposed lake floors. (Bowler, 1986). Gypsum and pelleted clay dunes (lunettes) line the edges of many salt lakes and playas in southeastern, southern and southwestern Australia; Prungle Lakes and Lake Fowler (gypsum lunettes), Lake Tyrell (clay lunette with occassional gypsum enrichment) and Lake Mungo (quartz sand lunette). All these lunettes are lake or pan-edge relicts from the Late Pleistocene deflationary period, when the lacustrine hydrology changed from perennial water-filled lakes to desiccated mudflats. Likewise, there are gypsum dunes in deflationary depressions in Salt Flat Playa and the Bonneville/Great Salt Lake region of Utah (Figure 4; Table 1).


Internal sedimentary structures in many of these lake-edge gypsum dunes or lunettes show tabular cross beds with consistent bedform orientation. Many lack abundant trough or festoon cross beds, suggesting consistent wind directions (Jones 1953; Bowler, 1973, 1983). Grain constituents clearly indicate deflation of former lake sediments, which were mostly vadose prior to deflation and passage into the dunes (Figure 4).

Gypsum dunes are part of a much broader lake-edge eolian sandflat association with the lakes often supplying large volumes of quartzose eolian sediment into adjacent sand seas or ergs (Figure 2; Warren, 2016). As mentioned pan-edge dunes described as ‘lunettes’ have a characteristic crescentic shape, other lake edge dunes may show more linear or longitudinal outlines, sometimes with parts of large sand seas or ergs being fed by the deflation of the salt lake or pan as at the southern edge of the Simpson Desert in Australia where it is in contact with the expanding and contracting edge of (Lake Eyre Figure 2).

Hydrological transitions from downwind evaporite dunes and lunettes

The role of salts, groundwater oscillations and the associated lake water levels/watertables are critical in creating eolian evaporites. Typically, once seasonal drying of an increasing arid lake floor sump begins, remaining surface waters with suspended clay become saline enough for the clay to flocculate and sink to the bottom of the desiccating water mass. If surface water concentration continues and the water surface sinks into the sediments to become a saline water table, then secondary gypsum prisms and nodules grow within the capillary zone of already-deposited sediment. In waters that are increasingly saline but not saturated with gypsum or halite, pelletization can continue to occur in the capillary fringe of clayey surface sediment (Figure 5).


Ongoing seasonal aridity further lowers the watertable in a saline mudflat, so the upper part of the vadose sediment column leaves the top of the capillary zone. It then deflates, leading to an accumulation of sand-sized sediment in adjacent eolian lunettes. If there is a prevailing wind direction, this builds significant volumes of dune sediment in a particular wind-aligned quadrant of the saline pan edge. Whether clay pellets or gypsum crystals are the dominant lunette component depends on the humidity inherent to the pan climate. In hyperarid situations, halite can be an eolian component in the lake hydrology (Salar de Uyuni; Svendsen, 2003).

In some lunettes, the mineralogy changes according to climate-driven changes in the hydrogeochemistry of the lake waters sourcing the lunette. For example in the Lake Tyrell lunette in semi-arid southwest Australia, the sediments in a layer range from clay pellets (75%) and dolomite (25%) in somewhat humid times of deflation to layers, with gypsum making up >90%, indicative of a more arid hydrochemistry. Lunettes associated with the shrinkage and deflation of Late Pleistocene Estancia megalake (New Mexico, USA) show similar variations in the proportions of clay pellet and gypsum sands in lake margin deposits around the edges of up to 120 blowout depressions. These blowouts define the former extent of the shrinking megalake and encompass both shoreline and lunette sands (Allen and Anderson, 2000)

Thus, the presence of an active gypsum lunette-field at a saline pan or playa edge is tied to landscape instability and a change from more humid to more arid conditions. To form a lunette requires a change in climate and an associated change in pan or playa hydrology and it hydrological base level and lake edge water table level, over time frames typically measured in hundreds to thousands of years.

 

Not just sand and dust-sized particles

Coarser than sand-sized gypsum crystals are transported in in lake margin mounds under hyperarid windy conditions that typify ephemeral pans and saline mudflats in parts of the Andean Altiplano and even higher elevations in the alpine tundra climatic zones. Salar Gorbea is a type example for this type of coarse-grained eolian transport (Figure 6; Benison, 2017). Whirlwinds, dry convective helical vortices, can move large gypsum crystals in their passage over the saline muflat. The transported gravel-sized crystals are entrained on the saline pan surface, after they first grew subaqueously in shallow surface brine pools. Once the pools dry up the crystal clusters disaggrate and then are transported as much as 5 km to be deposited in large dune-like mounds.

The dune gravel is cemented relatively quickly by gypsum cement precipitating from near-surface saline groundwater, resulting in a gypsum breccia. This documentation marks the first occurrence of gravel-sized evaporite grains being moved efficiently in air by suspension and provides a new possible interpretation for some ancient breccias and conglomerates, and improves understanding of limits of extremity of Earth surface environments.

 

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