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Evaporite interactions with magma Part 3 of 3: On-site evaporite and major extinction events?

John Warren - Saturday, April 13, 2019

 

Introduction

The previous two articles in this series dealt with heating evaporites, volatiles expelled into the atmosphere, and major biotal extinction events. I argued that short-term heating of a megaevaporite mass during emplacement of a Large Igneous Province (LIP) or heating of evaporities at the site of a large bolide impact, will move vast volumes of sulphurous and halocarbon volatiles, as well as solids, CO2 and CH4 into the earth's upper atmosphere (Figure 1a). The resulting catastrophic climatic effects link in time and probable causes to earth-scale major extinction horizons. (Figure 1b). In this article shall examine how three of the five major Phanerozoic extinction events have an evaporite association, starting with the most intense extinction event of the Phanerozoic; the end-Permian and its link to LIP emplacement into two separate sequences of massive bedded evaporite (Cambrian or Devonian mega-salts) in the Tunguska Basin, Siberia.


End-Permian - Saline interactions during emplacement of Siberian Traps

The Siberian Traps LIP is of significant size (~7 × 106 km2) and total volume (~4 × 106 km3) (Ivanov et al., 2013 and references therein). It is, however, smaller than the Late Cretaceous Deccan Traps and has a volume that is about a half of the Late Triassic Central Atlantic Magmatic Province (CAMP). All three of these continental LIPs are dwarfed by the Early Cretaceous marine Ontong-Java LIP (≈20 × 106 km3). So, it seems that the volume of igneous material in a LIP does not directly relate to the intensity of the extinction event (Figure 1b).

The Siberian Traps include ultramafic alkaline, mafic and felsic rocks that erupted in different proportions within a vast region extending over several thousands of square kilometres across Western and Eastern Siberia (Figure 2a). The Siberian Traps are considered have been emplaced atop a hotspot in a relatively short time frame (≈1 million years), when a large volume of deep mantle-derived igneous material was intruded and erupted at the Permo-Triassic boundary (Burgess et al., 2017).


Trap geology

Near Noril'sk, lava outflows reach thicknesses of over 3 km, while further to the northeast in the Maymecha-Kotuy region, half of the total lava pile is composed of ultramafic rocks including magnesian rich meimechites (Figure 2a). The very high MgO contents (8-40 wt %) of the meimechites in such low-degree melts indicates that the site of initial melting was very deep, as much as 200 km, and either in the lowermost continental lithosphere or in the underlying asthenosphere (Arndt et al., 1995). Melting probably was linked with the arrival of a mantle plume that was in its turn the source of the Siberian basaltic flood volcanism.

Thickness of volcaniclastic material in the Siberian Traps ranges from intercalated layers less than a meter thick on the Putorana Plateau to hundreds of meters near the base of the volcanic sections in the Angara and the Maymecha-Kotuy areas (Figure 2a). The total volume of mafic volcaniclastic material has been estimated at >200,000 km3 or >5% of the total volume of the Siberian Traps (Black et al., 2015). Volcanic rocks of this age are also present in drillcore in the West Siberian Basin (Ivanov et al., 2013).

Magma-sediment and magma-water interactions active during emplacement of the Siberian Traps in the upper lithosphere encompass a variety of heated evaporite interactions: batholith metal-evaporite interactions, lava-water interactions and intense phreatomagmatic explosions via vents and breccia pipes that formed saline-igneous volatile fountains reaching the upper atmosphere. The positions of these fountains are perhaps indicated by vent-related iron-rich diatremes (Figure 2a; Svensen et al., 2009). All these interactions are critical inputs to the End-Permian extinction event that links vast volumes of altered evaporites with the heating mechanisms inherent to Siberian Trap geology.

 

Evaporite basins (Devonian and Cambrian)

The Siberian Traps region is not only significant because of its vast extent and its deep nickel-prone mantle source, but also in that the immense volumes of igneous rocks that making up the traps were emplaced into two chemically prone saline giants with differing dominant mineralogies and ages; 1) Cambrian mega-halite sediments in the south, with interlayers of hydrated potash salts (mostly carnallitite) and 2) Devonian megasulphates in the north, containing two 50-100m beds of anhydrite (Figures 2b, 5). The interactions with the two types of salt basins, one halite-dominant, the other anhydrite-dominant, gives rise to two distinct meta-evaporite indicator associations. In the North, the interaction of picritic magmas with bedded thick anhydrites formed the supergiant Noril'sk nickel deposit, while in the south the LIP emplacement formed numerous magnetite-rich explosive breccia pipes, sourced at the stratigraphic level of the Cambrian salts (Figure 2b).


Norils'k region & Devonian evaporites

In the northern part of the Tunguska Basin the evaporite sediments hosting the intrusives of the Siberian Traps are a combination of Devonian anhydrites and carbonates, with overlying Carboniferous coals. Trap basalts, now cover this sedimentary sequence (Figure 4a), while sill-like tholeiitic intrusions, varying in composition from subalkaline dolerite to gabbro-dolerite are emplaced in the sediment pile and were part of the feeder system to the flood basalts (Figures 4b, 5, 6).


The region of Devonian evaporites contains the Noril'sk-Talnakh ore deposit, the largest Phanerozoic nickel deposit in the world (Figures 3, 4; Naldrett 2004). In the mine area, ore-bearing gabbroic-dolerites are differentiated, whereby picrite and picritic dolerite are overlain by more felsic differentiates. The Cu-Ni-platinoid mineralisation at Noril'sk forms relatively persistent stratabound horizons of massive sulphides in the lower portions of the three mineralised intrusions (Noril'sk, Talnakh, Kharaelakh), which are made up of segregations and accumulations of pyrrhotite, pentlandite and chalcopyrite (Figures 5, 6).


At the world-scale, the supergiant Permian Noril'sk-Talnakh deposit is an unusual Cu-Ni deposit. It did not form in the Precambrian, and so is unlike almost all the world's other supergiant magmatic nickel-sulphide deposits (Figure 3). It formed at the end of the Palaeozoic and straddles the Permo-Triassic boundary (Black et al., 2014a). Magmatic nickel ores at Noril'sk crystallised outside the influence of the reducing planetary atmosphere that typifies Archaean Ni flood basalt deposits and is not tied to greenstone terranes and the athenospheric transition to more sialic plate-scale conditions. (Figure 3). The high temperatures and near complete assimilation of Devonian sulphate evaporite blocks within the Noril’sk magma mean that this is one of the more enigmatic (“salt is elsewhere”) styles of evaporite-related high-temperature ore deposits (Warren, 2016, Chapter 16). Notions of evaporite assimilation for ore deposits tied to igneous-evaporite interactions are usually only one of multiple possible explanations of a magmatic ore but, in my opinion, for Noril’sk this is the most likely scenario. So, I emphasise the evaporite connection for the Noril’sk-Talnakh deposit in this article. Alternate non-evaporitic orthomagmatic explanations can be found in papers such as Wooden et al., (1992); Lightfoot et al. (1997), and Krivolutskaya (2016). Independent of the mode of nickel-ore fixation, most authors working in the Tunguska Basin agree that the emplacement of the trap intrusives drove the escape of a huge pulse of sediment-derived volatiles into the Earth's atmosphere.


Regional structure of the Noril’sk district is dominated by NNE-NE Permo-Triassic block faulting, which was coeval with magmatic activity. Individual faults may be over 500 km in length with throws of up to a kilometre (Figure 4b; Naldrett, 1997). Mineralised intrusions radiate outward and upward from intrusive centres and penetrate all levels of the overlying sedimentary sequence. Most intrusive centres are associated with prominent block faulting and fault intersections. The main Noril’sk-Kharaelakh fault occurs within the Siberian Platform, but is parallel to the main fault system that defines the boundary between the platform and the nearby Yenisei Trough. The Kharaelakh-Noril’sk fault guided the main upwelling magma body (Figures 4b, 6). Individual sills splay off this fault control and are interlayered with sulphate evaporite beds to can attain lateral lengths of 12 km, widths of 2 km and thicknesses of 30 to 350 m.

Mineralogical compositions of the Devonian sediments interlayered with these sills are of great importance in understanding the geological responses to heating by intrusive igneous sills in the Noril'sk-Talnakh area (Figures 5, 6). Based on their lithological features and paleontological character, the intruded Devonian succession is subdivided into the Yampakhtinsky, Khrebtovsky, Zubovsky, Kureysky, and Razvedochninsky Formations (Lower Devonian), the Manturovsky and Yuktinsky Formations (Middle Devonian), and the Nakohozsky, Kalargonsky and Fokinsky Formations (Upper Devonian) (Figure 5; Krivolutskaya, 2016; Naldrett, 2004). The two main evaporite levels are the Middle Devonian and Lower Devonian anhydrite-dominant successions, both deposited in a subsealevel transitioning rift (Figure 5, 6; Naldrett, 2005; Warren, 2016).

The Yampaktinsky and Khrebtovsky Formations consist of Lower Devonian carbonates interbedded with abundant gypsum (in outcrop) and anhydrite (subsurface), along with some of the oldest lenses of celestite in the area (Figure 5). The total thicknesses of these two CaSO4 units are around 100 and 80 m, respectively. The Lower Devonian Zubovsky Formation is composed of grey-colored dolomitic marls interbedded with argillaceous dolomites, mudstones, and anhydrite with a total thickness of 110–140 m. The Zubovsky Formation unconformably overlies the Lower Devonian Khrebtovsky Formation in the Noril’sk region. The Lower Devonian Kureysky Formation consists of mottled dolomite and calcareous mudstones and marls with rare siltstone and limestone. The thicknesses of all units in the outcrop section remain stratiform and vary within 50–60 m. The contacts with the overlying and underlying formations are conformable.

The Lower Devonian Razvedochninsky Formation is dominated by siltstones, sandstones, and conglomeratic sandstones with a thickness that regionally does not exceed 110–150 m, but reaches 150–235 m in troughs, and decreases sharply to the south until fully wedging out.

The Middle Devonian Manturovsky Formation overlies the eroded Razvedochninsky Formation and consists of a terrigenous-carbonate section with abundant salt-bearing strata, most of which consist of rock salt or brecciated equivalents. This formation’s thickness is 100-210 m but ranges up to 500 m (Figure 6). The Middle Devonian Yuktinsky section is dominated by clastic–carbonate sediments ranging from 12 to 40 m thick, while in the troughs the thickness of interlayered sulphate rocks reaches 55 m. The contacts with the underlying and overlying Middle Devonian Manturovsky deposits are considered comformable. The Upper Devonian Nakokhozsky Formation consists of folded calcium-sulphate-rich variegated shale–carbonate rocks with a thickness of 2–60 m that increases in the troughs to 80–130 m (Figure 5). The Upper Devonian Kalargonsky Formation is characterised by a grey-colored terrigenous-carbonate section that includes dolomites, dolomitic marl, dolomite–limestone, and anhydrite dominate in the basins. This formation’s thickness is 170–270 m. The Kalargonsky Formation unconformably overlies the Middle Devonian Nakokhozsky sediments and the contact is typically a breccia (Figure 5).

The Middle Devonian Fokinsky Formation (as distinct from the mineralised Fokinsky intrusions) consists of evaporite sulphate-rich clastic–carbonate sequences, primarily within the troughs, and anhydrite, dolomitic marls interbedded with limestone lenses of rock salt, and clay–carbonate breccias (Krivolutskaya, 2016). The thickness of this formation is 220–420 m (approximately 500 m in the western part of the Vologochansky Trough).

The Fokinsky Formation is not recognised by all authors working in the region. This disparity in stratigraphic recognition across the region underlines a problem inherent in the litho-stratigraphic descriptions of many bedded evaporite regions worldwide, where it is assumed that a layer-cake stratigraphy/correlation is present pre- and post-intrusion. Thereby the effects of evaporite collapse dissolution, bed wedge-out and possible salt flow are not quantified. In my opinion, sedimentary breccias in such regions are more likely to be diagenetic and laterally discontinuous (see Warren, 2016; Chapter 7).

In summary, the Devonian stratigraphy in the vicinity of the Noril'sk Mine retains significant thicknesses (50-100m) with variations centred on transitions in and out of bedded anhydrite. There is a strong likelihood that the current outcrop geology interpretations under-illustrate former thicknesses of bedded evaporites during to ongoing dissolution, collapse and possible flowage.

The anomalous Phanerozoic age of the Noril’sk-Talnakh ore deposits, compared with the Precambrian ages of other magmatic Ni-Cu deposits, and its relative enrichment in Ni, Cu, Pt and Pd compared with Sudbury and Jinchuan (Figure 3), is thought to reflect the anomalously high volumes of sulphur in the parent magma. Additional sulphur entered the evolving magma chamber via intrusion and assimilation of CaSO4 blocks and associated hydrothermal solutions altering and dissolving adjacent thick-bedded anhydrite successions (Figure 7; Naldrett 1981, 1993, 1997; Pang et al., 2013). Noril'sk-Talnakh's rich sulphur supply contrasts with that of the komatiitic Archaean Cu-Ni deposits, where the sedimentary sulphur supply came from more ubiquitous, less-focused sulphur sources sometimes entrained in widespread sedimentary pyrite (Figure 3). Such pyrite characterises a significant portion of fine-grained sediments accumulated under an anoxic reducing Archaean to Palaeoproterozoic atmosphere.


Abundant crystals of magmatic anhydrite today typify the olivine-bearing (picritic) gabbros in the Kharaelakh intrusion, which is located in the basin stratigraphy at the level of the Devonian anhydrites (Figure 6; Li et al., 2009 Spiridonov, 2010). Along with disseminated sulphides, the anhydrite crystals are characterised by planar boundaries with co-associated olivine and augite. Dihedral angles of ~120°, characteristic of simultaneous crystallisation, are common throughout the anhydrite-augite assemblages. Inclusions of anhydrite in augite and vice-versa are also typical.

Rounded and subrounded sulphide inclusions composed of pyrrhotite, pentlandite, and chalcopyrite, that crystallised from immiscible sulphide liquid droplets in the magma, are commonplace within the magmatic anhydrite crystals and in the contact aureoles (Figure 7). Visual estimates by Li et al. (2009), based on five polished thin sections, indicate that the ratio of anhydrite to sulphide in mineralised samples varies from 0.05 to 0. The observation of abundant wollastonite in contact aureole rocks at this stratigraphic level suggests that reactions such as CaSO4 + SiO2 + H2O = CaSiO3 + H2S + 2O2 occurred, and that sulphate was likely reduced to sulphide before incorporation into the magma (Ripley et al., 2007).

Picritic magmas in mantle plumes can have melt temperatures as high as 1600°C (Hezberg et al., 2007). Assimilation of anhydrite via partial melting of a cooler basaltic magma at shallower depths can be more difficult, owing to the high melting point of pure anhydrite (melt temperatures typically rang between 1360 and 1450°C, although this is significantly lowered in the presence of organics and water). Rather than only melting anhydrite enclosed by picritic magma, additional fluxing mechanisms likely move additional anhydrite-derived sulphur into the melt, either by hydrothermal leaching of sulphate followed by partial reduction, or via a process involving the dissolution of anhydrite during thermochemical sulphate reduction (TSR; Warren, 2016; Chapter 9). The latter process requires heat, anhydrite and organics (generally in the form of hydrocarbons or kerogen).

Some authors use the euhedral outline of anhydrite in mineralised sills, as seen in Figure 7, to argue blocks anhydrite country rock was not assimilated. This is a specious argument as this type of anhydrite was precipitated during cooling of an already sulphur-saturated magma, the euhedral spary outline does not relate to the source of the sulphur, which is more clearly indicated by its sulphur-isotope signature (Figure 8a - also Warren, 2016; Chapter 8).


Isotopic analysis of δ34S in the magmatic anhydrites and associated metal sulphides in the Kharaelakh intrusives require the assimilation of externally-derived high-δ34S sulphur from the adjacent country rock (Figure 8: Ripley et al., 2007, 2010). Where complete sulphate reduction occurred, the δ34S values require mixtures of some 60% anhydrite-derived evaporitic marine sulphur (δ34S values near 20‰), with 40% mantle-derived sulphide (δ34S of 0‰) to produce the required measured magmatic sulphide values ≈12‰ (Figures 8a, b).

The sulphur isotope data and the nature of the sampled contact aureoles suggest intense intracontinental rifting in the Noril’sk region brought deeply-sourced mafic magmas into contact with supracrustal sulphur from evaporitic sulphates at the level of the Kharaelakh intrusion. Sulphur isotope data show the mineralised intervals at Noril’sk are anomalously heavy in δ34S (Figure 8a, b). These data are inconsistent with sulphur derived from mixing of the mantle magma sulphur (δ34Svalues near zero) with sulphur from an evaporitic sulphate source (Godlevski and Grinenko, 1963; Grinenko, 1985; Li et al., 2009; Pang et al., 2012; Black et al., 2014a).

Sulphur isotope values from Paleozoic evaporites vary between +10 and +35‰ (Figure 8b; Claypool et al., 1980). Cambrian evaporites, including the major Irkutsk basin salts in Siberia, are the most 34S-enriched evaporites in the Phanerozoic, with mean δ34SVCDT = +30‰ (Claypool et al., 1980; Black et al., 2014a). Two-member mixing curves between meimechite and anhydrite sulphur (with δ34S = +20 to +35‰) convincingly reproduce the observed δ34S trends for the Noril'sk ores (Figure 8b; Black et al., 2014a).


As the magma rose through the sedimentary cover, it penetrated and assimilated sulphur from extensive Devonian anhydrite layers (Figure 9). Sulphur in calcium sulphate was reduced to sulphide, CaO entered the magma, and iron from the magma reacted with reduced sulphur so that the end result was droplets of immiscible iron sulphide dispersed through the melt (Naldrett and Macdonald, 1980). These droplets acted as collectors for Ni, Cu and the platinum group elements, which are now so enriched in the Noril’sk ores.

Naldrett (1991, 1997, 2005) concluded that prehnite + biotite + anhydrite + carbonate + zeolite + chlorite ± sulphide globules, which typify chromite agglomerations in the picrite of the Noril’sk intrusions, represent remnants of partially assimilated sulphate-rich country rock. Assimilation of anhydrite-rich rocks, coupled with the reduction of sulphate to sulphide, would have introduced considerable oxygen into the silicate melt, which then drove precipitation of chrome-spinel minerals (chromite - FeCr2O4; mangnesiochromite - MgCr2O4). Inclusions of anhydrite-rich material, floating in the magma, would have served as loci for chromite crystallisation, thus giving rise to the association between the agglomerations and the globules. Tarasov (1970) pointed out that Middle Carboniferous coal measures were also assimilated and may have supplied organics that assisted in the reduction of sulphur in the magmas (Figures 6, 7).

Evidence of the assimilation of large volumes of anhydrite and coaly organics into the magma mass has implications beyond the formation of the Noril’sk-Talnakh ore deposits. Li et at. (2009) identified magmatic anhydrite-sulphide assemblages in a subvolcanic intrusion associated with the Siberian Traps. The δ34S values of anhydrite and coexisting sulphide crystals analysed by ion probing are 18‰–22‰ and 9‰–11‰, respectively, are much higher than the anhydrite-contaminated ore values shown in Figure 8). To obtain this level of fractionation means more than 50% of the total sulphur in the intrusion was derived from marine evaporites in the footwall strata. The contaminated magma was highly oxidised and able to dissolve up to one order of magnitude more sulphur than pure mantle-derived basaltic magma. Such sulphur-contaminated magma, when erupted, would have released vast volumes of SO2 into the atmosphere (Black et al., 2012, 2014b). That is, the eruption of the anhydrite-contaminated magma that is the Siberian Traps in the northern Tunguska Basin can help explain the intensity of the end-Permian extinction.

In summary, such igneous - sulphate sediment interaction explains, at least in part: (1) the vast amount of sulphide melt in the Noril’sk-Talnakh ore field; (2) the heavy quasi-anhydrite isotopic composition of sulphur in sideronitic and massive nickel ores; (3) the reduced contents of noble metals in these ores (compared with the drop sulphides that occur toward base of the intrusions and have a likely mantle sulphide source); and (4) the high contents of radiogenic (crustal) osmium in sideronitic and massive ores (Spiridonov, 2010; Walker et al., 1994). In summary, the reserves of the world-class Ni-PGE deposit at Noril’sk-Talnakh, with its anomalous Phanerozoic age, likely reflect a fortuitous occurrence of thick Devoninn anhydrites (ultimate sulphide source) atop an active later set of deep mantle-tapping rift grabens that drove the LIP outlined by the Siberian Traps. Wherever these magmas vented into the Earth's atmosphere they carried significant volumes of sulphurous volatiles.


Cambrian evaporites, potash & breccia pipes

Salt deposits of late Vendian to Early Cambrian age in East Siberia cover an extensive area (ca. 2 million km2) located to the north-west of Lake Baikal with an extent showing it extends across much of the Permian Siberian Traps (Figure 2b). The thickness of this upper Vendian-Lower Cambrian evaporite succession is 2.0–2.5 km in the southern, western, and central parts of the basin, and 1.3–1.5 km in the NE part (Nepa-Vilyui). This saline giant (total volume of upper Vendian–Lower Cambrian evaporites is 785,000 km3; Zharkov, 1984) is characterised by the occurrence of fourteen regional marker carbonate units and 15 salt units (Figure 10; Zharkov, 1984, with references therein). Five major phases of salt deposition are distinguished, namely the late Vendian (Danilovo) and Early Cambrian (Usolye, Belsk, Angara, and Litvintsevo) salt basins (Figure 10a; Zharkov (1984), Kuznetsov and Suchy. (1992).

Average thicknesses of the Cambrian evaporite deposits decreases with time (Figure 10b) as does the area (Figure 10a). The area of the oldest Cambrian basin, the Usolye salt basin is almost 2 million km2, and the average thickness of deposited salt around 200 m (Zharkov, 1984), while the area of the youngest, Litvintsevo salt basin is 0.5 million km2 and the average thickness of its evaporite bed (rock salt and anhydrite) is 50 m (Figure 10; Zharkov, 1984).

Most of the petroleum reservoirs in the region are located in the Cambrian carbonates. The post-Cambrian stratigraphy contains major erosional breaks. As we saw in the Noril'sk discussion, Devonian evaporites are rare in the south but abundant in the north, whereas Ordovician rocks (limestones, marls) are locally abundant in the central parts of the basin. Cambrian salt deposition is interpreted as mostly taking place in a deeper water basin: Petrichenko (1988) concluded that at the termination of halite deposition the final brine depth was 50–260 m, and at the onset of potash deposition it was ≈10–50 m.


Lower Cambrian Angara evaporites host the largest known bedded potash deposit in Russia, which is not yet produced (Figure 11; Garrett, 1995; Warren 2016, Chapter 11). Potash salts occur at the base of the Angara Formation in what is called the sixth halite series (Table 1). This intracratonic potash basin is one of the larger potash-entraining salt sumps in the world, it is several times larger than the Permian Upper Kama deposit and approaches the Prairie Evaporite in aerial extent, but not in lateral continuity, thickness or purity (Figure 11) due in large part to the effects of igneous disburbance.

Plans were made in 1986 under the old Soviet regime to initiate a mining program in a section of this basin called the Nepskoye deposit but were never fully implemented, although some ore was extracted in the mid 1980s (Andreev et al., 1986). The proposed potash development region is located near the towns of Nepa and Ust-Kut (300 km apart) in Irkutsk State. Regionally, the dominant potash mineral is carnallite, but high-grade sylvinite is intersected at depths of 600-1,000 m in beds some 1.5-5 m thick over an area ≈ 1,000 km2 (Garrett, 1995). The lower Bur or K1 bedded potash horizon lies at a depth of 750 - 960 m and is 2-18 m thick (4-6 m in the central area (Figure 11a; Table 1). Two sylvinite zones in this horizon were mapped, with the central one being 16-26 km long and 6-8 km wide (Figure 11a). In the lower horizon (K1) the sylvinite was 1.5-3 m thick, and averaged 15-50% KCl, 0.05-0.5% MgCl2,with 0.5% insolubles. The overlying K2 potash zone (Tunguaka) also entrains several sylvinite beds and is some 679-880 m deep and 2.5-20 m thick. It has a 15-45% KCl content and comparatively low MgCl2 and insoluble contents. This zone represents the major potash reserves of the deposit. In the upper potash beds (K2) the sylvinite strata become more discontinuous, but some reasonably thick, high grade and extensive zones exist (Andreev et al. 1986). The sylvite ore sits in a more regional potash succession composed of a combination of carnallitite and sylvinite (Figure 11b). The broader Nepa potash region as generally mapped in Figure 2 has two interesting characteristics; 1) The igneous trap rocks as defined in the drill-controlled cross sections of Malykh and Geletii (1988) sit below the potash level (Figure 11b), 2) There is a paucity of magnetitic explosive breccia pipes in the Nepa potash region (Figure 2b).

To the south and west, between Irkutsk and Taseyevo some 400 km to the west, other large potash occurrences have been reported in the same general but poorly delineated evaporite basins. For instance, in the Kanak-Taseyevo basin, potash beds (sylvite-carnallite containing 3-24% K2O) have been intersected at depths of 1,240-1,415 m (Garrett, 1995). Potash beds at these depths would require a solution mining methodology, but the at-surface climate would mean either cryogenic pan processing or evaporators, making recovery more difficult and expensive (Warren, 2016; Chapter 11).

Basaltic breccia pipes, Tunguska Basin Siberia

Basalt pipes form a rim to the main basalt body of the Siberian Traps and are genetically linked to trap emplacement (Figure 2b; Polozov et al., 2016). The pipes pierce through all sedimentary strata, even dolerite sills higher in the Permo-Carboniferous portion of the basin stratigraphy, and are considered to be a type of diatreme. Importantly, the basalt pipes with magnetite cores tend to occur across the southern Tunguska Basin, while unmineralised basalt pipes are more widespread (Figure 2b). Some of the basalt pipes bearing magnetite mineralisation are of commercial grade and are mined for their iron ore.

Regionally, it is difficult to estimate the total number of pipes (both “barren” basalt and magnetite-enriched) because repeated glaciations have flattened relief, while thick taiga forest covers significant parts of Siberia. Thus, many pipes are hidden by swampy coniferous forests and so are difficult to map. However, conservative estimates based on prospecting surveys for iron mineralization in the southeern portion of Tunguska Basin, and geological mapping elsewhere, suggest there are more than three hundred magnetite-bearing basalt pipes. This includes 6 large (>100 Mt of iron ores), 14 medium (20–100 Mt) and 19 small <20 Mt) sized iron deposits. All other mineralised basalt pipes are currently of sub-economic grade or underexplored (Polozov et al., 2016).

The magnetite deposits are consistently located in the Tunguska Basin region underlain by Cambrian evaporites and mainly defined by subvertical and cylindrical breccia bodies with magnesio-ferrite and magnetite as the primary ore minerals (Figure 2b). In many ways these deposits are similar to iron oxide, gold and copper (IOGC) deposits worldwide, but are classified in the Russian literature as Angara–Ilim type deposits, named after the two rivers where a large number of iron- mineralised basalt pipes crop out (Soloviev, 2010; Warren 2016, Chapter 16).

Korshunovsky (Korshunovskoye) region, Siberia

This region, in the Irkutsk district, is the eighth largest iron ore producer in Russia, with an annual output of 5 Mt of iron ore concentrate. Across the region, the pipes are sourced in the Cambrian evaporite part of the basin stratigraphy and pierce younger Paleozoic sediments composed of argillites, limestones, marls, siltstones, sandstones and clays of the late Cambrian Lena, Ust'kut, Mamyr and Ordovician Bratsk groups and overlying Early Carboniferous limestone.


We shall focus one of the largest magnesite deposits in the region, the Korshunovshoe (Korshunovsky) magnetite breccia pipe, with an estimated reserve of 1.5 Gt of ore to a depth of 1700 m (Soloviev, 2010; Polozov et al., 2016). It is mined (open pit) and so the interior structures and relationships are well documented (Figure 12). The currently mined pipe is adjacent to another explosion pipe to the immediate south-east, with the mineralised breccias sourced mainly at the level of the Cambrian evaporites (halite, potash and anhydrite; Mazurov et al., 2007). At outcrop and in the pit little evidence, other than secondary textures (dissolution-collapse and brecciation), remains of the primary minerals of the mother saline layer, although remnant, recrystallised evaporite clasts (including halite and anhydrite) typify the mineralised breccia in the lower parts of the pipe (Mazurov et al., 2007, 2018). Textures at the evaporite level in the diatremes are not unlike those seen in regions of Eocene sill interaction with hydrated salts in the Zechstein potash mines of East Germany (Schofield et al., 2014; Warren 2016 and part 1 in this series of Salty Matters articles).

The Korshunovshoe pipe is filled with tuff breccias and fragmentals composed of the surrounding saline country rocks which have undergone considerable metasomatic alteration. They incorporate fragments and larger blocks of sedimentary (60 to 80 vol.%; sandstones, siltstones, limestones, evaporite residues and argillites) and igneous (10 to 40 vol.%; gabbro-dolerites, dolerites and basalts) rocks, cemented by essentially chloritic material as well as by fine-grained carbonate (Figure 13). The central part of the magnetitic diatreme characterised by intense multiple brecciation, with rock fragments in the breccias represented mostly by variably-altered dolerites. They are cemented by a finely-dispersed matrix, entirely replaced by skarn, post-skarn alteration assemblages and iron oxides.


Outside of this zone, intense fracturing has occurred, locally with brecciation in altered sedimentary rocks. The fractures are filled with magnetite, accompanied by chlorite and calcite. Finally, the outermost zone is characterised by weak, predominantly sub-horizontal fractures within sedimentary host rocks, locally replaced by skarns. Steeply-dipping dykes of gabbro-dolerite, dolerite, dolerite-porphyry, and basalt-porphyry are present, both within and outside the breccia pipes, while sub-horizontal dolerite sills occur at depth (Soloviev, 2010; Mazurov et al., 2007, 2018).

Magnetite pipe orebodies at Korshunovshoe are texturally and mineralogically complex (Figures 12, 14) and are composed of: i) Banded masses of metasomatic magnetite that are within, and conformable to saline to calcareous members of the host sedimentary wall rocks (dominantly in dolomitic limestones, marls, calcareous argillites and sandstones with a calcareous or limy matrix, but only to a minor degree in sediments without a saline carbonate component) at a depth of some 700 to 1500 m from the surface; ii) Stock-like, lensoid, layered and columnar bodies of magnetite within the altered pyroclastics of the breccia pipe; and iii) Steeply dipping vein-like masses in zones of intense brecciation and replacement by skarns.

Together these mineralisation styles form two large continuous bodies in the Korshunovshoe pipe (Figure 12). The main deposit has the form of a sub-vertical breccia pipe with plan dimensions of approximately 2400 x 700 m. Mineralisation has been traced by drilling to a depth of 1200 m, and by geophysical data to at least 3 km below the surface (Soloviev, 2010).

The bulk of the ore is associated with brecciation and occurs within sediments, tuffs and igneous rocks and are demonstrably due to the partial replacement and alteration of the host. Massive and banded ores are less well developed. The mineralisation is mostly magnetite (≈82% of iron resources), with minor magno-magnetite, hematite and martite. The main orebody comprises vertically overlapping zones, with variable amounts of hematite and martite in the upper layers, calcite and magnetite in middle layers, and halite and magnetite in lower layers. The magnetite of the upper to middle zone is accompanied by pyroxene, chlorite and minor epidote with lesser amphibole, serpentine, calcite and garnet, and rare quartz, apatite and sphene and occurs as oolites, druses, masses and disseminations. Calcite increases downwards to 20 to 30%. In the lower part of the deposit, halite, amphibole and Mn-magnetite are more abundant. Pyrite, chalcopyrite and pyrrhotite are found throughout. Much of the magnetite is magno-magnetite which contains up to 6% MgO.

Across the region of magnetite breccia pipes, ore is extracted from magmatic diatremes that completely penetrated the highly evaporitic lower Phanerozoic succession (Figure 14; Mazurov et al., 2007, Polozov et al., 2016). Early work on this intrusive magnetite style, which surrounds brecciated diatreme-like pipes, classified it as a skarn association, forming a halo around a set of explosive pipes that accompanied regional trap magmatism (Ivashchenko and Korabel’nikova, 1960).

Characteristic spinel-forsterite magnesian skarns are confined to the overdome parts of large doleritic bodies and are the result of interactions of massive evaporitic and petroliferous dolomites with fluids released from liquid magma (Mazurov et al., 2007). Magnesian skarns of the postmagmatic stage are localised in the marginal parts and on the front (outwedged portions) of doleritic sills, apophyses, and the branches of intrusive bodies hosted at the level of the Cambrian carbonate-evaporite successions (Figure 14). The skarns penetrating the evaporite levels have a banded or layered structure and resemble gravel conglomerates, with carbonate cements. The round fragments (metasomatic pseudo-conglomerates) are composed of globules of disintegrated doleritic porphyrite, completely or partially substituted by zonal magnesian skarns. Their mesostasis is cryptocrystalline, and early phenocrysts of olivine, plagioclase, and pyroxene have undergone dispersion and substitution. Unaltered cores of the metasomatic ‘conglomerate’ are in contact with a fassaite zone, which passes outward into a spinel-fassaite zone and then into a forsterite-magnetite and calciphyre zone.


The geometry of pipe emplacement is broken down into three related styles; i) Root zone, ii) Diatreme zone, iii) Crater zone (Figure 14). The upper crater zone is sometimes complicated by the presence of reworked crater-lacustrine deposits (Polozov et al., 2016). The root zone is typically brecciated with pseudoconglomerated and other saline volatilisation textures described in the previous paragraphs. The root zone can be traced out from the pipe stem as disturbed zones with considerable lateral extents at the level of the Cambrian evaporite beds. Subhorizontal brecciated dolerite “sills” of the Kapaevsk iron deposit were cemented with calcite, magnetite and halite in various ratios and traced down to deep levels close to the root zones in some basalt pipes In the Korshunovsk iron-ore deposit, such a brecciated body extends from the main diatreme pipe some 5 km to the west and 9 km to the south-west (Von der Flaass and Nikulin, 2000).

Although not discussed in terms of a volatilisation mechanism in the published literature, I would argue that the lateral apophsyes are indicative of the former presence of hydrated salt layers, probably carnallitite beds showing similar responses to those seen in the potash mines of East Germany (Shofield et al., 2014; or part 1 in this current series of Salty Matters articles).

The diatreme chimney atop the root zone indicates the rapid rise of a overpressured and upward flowing gas-charged rock mass. Basalt magma served as the ultimate source of iron for the magnetite in the breccia pipes. Extraction of iron from the melt and its transition and accumulation took place in the presence of chlorine-rich fluids, which were formed in the course of thermal decomposition of halite-hosted hydrated salt beds (carnallite). In the later stages of ore formation, some chlorine was fixed in scapolites, while sodium was fixed in albitites and scapolitites (dipyres). In the tuffs of a number of diatremes and paleovolcanoes of the Siberian Platform, native iron can form metal balls in association with moissanites and diamonds (Goryainov et al. 1976). The occurrence of such phases, as well as bitumen in calderas and carbonaceous matter in pisolite tuffs, points to the migration of hydrocarbon fluids through the volcano-tectonic structures (Ryabov et al., 2014).


Hydrocarbons are abundant in the Cambrian and Ordovician sections of the Tunguska Basin, while coals are widespread in the Permo-Carboniferous Tunguska Series sediments (Figure 15). The juxtaposition of a vast volcanic province with its dykes, sills and diatremes interacting with extensive intracratonic saline Cambrian beds containing evaporites sealing substantial oil accumulations and interacting with coal-bearing deposits, likely produced massive quantities of halocarbons along with methane and CO2. Notably, contact metamorphism with hydrothermal systems rich in chlorine, created during pressure dissolution and dehydration of the surrounding evaporites, potentially synthesized large amounts of the organohalogens methyl chloride (CH3Cl) and methyl bromide (CH3Br) (Beerling et al., 2007; Visscher et al., 2004; Svensen et al., 2018).

In terms of rapid transfer of volatiles to the atmosphere, the phreatomagmatic-sediment pipes (diatremes) generated tall, explosive volatile-rich eruption columns, which at times reached the stratosphere (Svensen et al., 2009). Such features simultaneously promote removal of highly soluble volcanic gases, such as HCl and SO2, and potentially deliver large volumes of sulphur, halocarbons water, methane and CO2 to the upper atmosphere (Black et al., 2015).

Timing of trap emplacement

Siberian Traps magmatic activity at the end-Permain is segmented into three distinct emplacement stages (Figure 16; Burgess et al., 2017). Stage 1, beginning just before 252.24±0.1 Ma, was characterised by initial pyroclastic eruptions followed by lava effusion. During this stage, an estimated two-thirds of the total volume of Siberian Traps lavas were emplaced (>1×106 km3). Stage 2 began at 251.907±0.067 Ma, and was characterised by cessation of extrusion and the onset of widespread sill-complex formation. These sills are exposed over a >1.5 × 106 km2 area and form arguably the most aerially extensive continental sill complex on Earth. Intrusive magmatism continued throughout stage 2 with no apparent hiatus. Stage 2 ended at 251.483±0.088 Ma, when extrusion of lavas resumed after an ~420 ka hiatus, marking the beginning of stage 3. Both extrusive and intrusive magmatism continued during stage 3, which lasted until at least 251.354 ± 0.088 Ma, an age defined by the youngest sill dated in the province. A maximum date for the end of stage 3 is estimated at 250.2 ± 0.3 Ma.


Integration of LIP stages with the record of mass extinction and carbon cycle at the Permian-Triassic Global Stratotype Section and Point (GSSP) shows three important relationships (Burgess et al., 2017). (1) Extrusive eruption during stage 1 of Siberian LIP magmatism occurs over the ~300 kyr before the onset of mass extinction at 251.941 ± 0.037 Ma. During this interval, the biosphere and the carbon cycle show little evidence of instability. (2) The onset of stage 2, marked by the oldest Siberian Traps sill, and cessation of lava extrusion, coincides with the beginning of mass extinction and the abrupt (2–18 kyr) negative δ13CPDB excursion immediately preceding the extinction event (Figure 16a). The remainder of LIP stage 2, which is characterised by continued sill emplacement, coincides with broadly declining δ13CPDB values following the mass extinction. (3) Stage 3 in the LIP begins at the inflexion point in δ13CPDB composition, after which the carbon reservoir trends positive, toward pre-extinction values.

Explosive volcanism in the Siberian Traps can be classified in three distinct groups: 1) deep-rooted sediment–magma interactions and pipe eruption where feeder sills are emplaced in evaporites (Cambrian and Devonian country rock), 2) shallower magma-water interactions in areas with abundant groundwater or hydrated salts, and 3) lava flows and lava fountaining during the main stage of effusive volcanism (Jerram et al., 2016a,b). Each stage has a differing set of expressions in terms of the interacting evaporites and the landscape expression of these interactions.

Outcomes of the end-Permian igneous evaporite interplay

A unusual aspect of the Siberian trap eruption compared to many but not all LIPs is the saline and kerogen-rich nature of regional geology in the Siberian platform that interacted with the LIP magmas. The main lithologies of the region are large volumes of Devonian anhydrites in the north, Cambrian halite and hydrated-potash salts in the south, hydrocarbon source rocks and evaporite-sealed hydrocarbons, and coals in the Permo-Carboniferous portions of the stratigraphy sitting directly below the basaltic otflows. Notably, contact metamorphism and the development of hydrothermal systems rich in chlorine (produced from the pressure dissolution and volatilisation of the surrounding evaporites, kerogens, coals and hydrocarbons with evaporite seals) potentially synthesized large amounts of the organohalogens methyl chloride (CH3Cl) and methyl bromide (CH3Br) along with vast volumes of sulphurous gases, CH4 and CO2 (Figure 17).


End-Triassic extinction event - Saline interactions with CAMP magmas

The Central Atlantic Magmatic Province (CAMP) was emplaced at the end of the Triassic (≈201 Ma) in a region created by the tectonic unzipping (rifting-breakup) of the Pangean supercontinent (Figure 18; Marzoli et al., 2018). CAMP extends across the former Pangaea from modern central Brazil northeastward some 5000 km across western Africa, Iberia, and northwestern France, and from Africa westward for 2500 km through eastern and southern North America and as far west as Texas and the Gulf of Mexico (Figure 18 - dashed red line). The Province is composed of basic igneous rocks emplaced in a combination of shallow intrusions and erupted large lava flow fields extending over a land surface area in excess of 10 million km2. During its emplacement, sill intrusions into evaporites are particularly widespread in the vast Amazonas and Solimões intracratonic basins (≈1 ×106km2), representing up to 70% of the total CAMP sill volume (Svensen et al., 2018).


Sedimentary rocks intruded by sills in the Amazonas and Solimões basins include a lower (Ordovician–Mississippian) and upper (Pennsylvanian–Permian) Paleozoic series (Milani and Zalán, 1999). The lower Paleozoic series consists of sandstones and shales, some of which are particularly organic-rich (total organic content up to 8wt.%; Milani and Zalán, 1999; Gonzaga et al., 2000). The upper Paleozoic series is dominated by evaporite and carbonate deposits of varying abundances, interlayered with clastics. Sills are widespread within the upper Paleozoic evaporitic sequence, extending almost continuously from the western margin of the Solimões Basin to the eastern margin of the Amazonas Basin (Fig.19c). Sills within the lower Paleozoic unit are restricted to the eastern part of the Amazonas basin. As illustrated in Fig.19c, high-Ti sills are found only in the lower Paleozoic series. Let's look now at the saline geology of the region and then at the effect its assimilation had on sill geochemistry.


Saline geology

A significant, as yet poorly delineated, set of variable hydrated potash salts and sylvinites occur in bedded halite in the Amazon Basin, Brazil (Figures 19a, 20; Szatmari et al. 1979). The Amazon Basin is about 2,100 km long and 300 km wide, it is an intracratonic sag basin atop an aulacogen between the Guyana and Guaporé cratons (Figure 19b). The basin fill contains a number of stacked mega-sequence cycles (as defined by wireline interpretation) ranging in age from Lower Ordovician (Autaz Mirim Member of Trombetas Formation) to Lower Permian (Figure 19b; Andirá Formation; Gonzaga et al., 2000). The basin has a widespread Upper Cretaceous cover (Alter do Chão Formation) and was affected by widespread tholeiitic magmatic activity at the end-Triassic (e.g.Penatecaua dolerites of the CAMP), making seismic-based hydrocarbon exploration difficult, especially as much of the basin still lies beneath thick tropical jungle. Since the recognition of a widespread igneous overprint of the Palaeozoic sedimentary succession in the 1970s, hydrocarbon exploration efforts have been subdued (Thomaz-Filho et al., 2008). However, in the past few years, SRTM studies are proving useful, in front of seismic surveys and drilling, in the general identification of geological features in the Amazon Basin (Ibanez et al., 2016)

Th Amazonas-Solimoes intracratonic sag basin is developed on the same scale as the Alberta basin of Canada and entrains the Carboniferous (Pennsylvanian, ≈305 Ma) saline Nova Olinda Formation. It is made up of a large laterally extensive set of cyclic evaporite beds, dominated by interbedded combinations of anhydrite, shale and halite (Figures 20c, 21). These evaporites occur within the Carboniferous-Permian megasequence, known as the Tapajós Group, which can be up to 1600m thick (Milani and Zalan, 1999). The lowest part of the megasequence is a blanket of eolian sandstones (Monte Alegre Formation), which is covered by marine-influenced carbonates and evaporites (Itaituba and Nova Olinda Formations, respectively), along with subordinate sandstones and shales (Figure 19c). The Tapajós megacycle is closed by a suite of Permian continental redbeds (Andirá Formation) of Permian age. Subsequent east-west regional extension facilitated a pervasive intrusion of magmatic bodies during the end-Triassic to Early Jurassic (Penatecaua dolerites and equivalents).

Individual halite beds in the Nova Olinda evaporite cycles are 20-80 m thick, while the Nova Olinda Fm. has an average thickness of 900m. Because of the high levels of entrained anhydrite beds in the Nova Olinda Fm., evaporite layers are not halokinetic, but are subject to collapse and flow about the basin margin, especially in areas of intense meteoric dissolution (Figure 20).


Early Petrobras drilling programs conducted in the Amazon Basin from 1953 to 1963, defined the presence of halite but did not appreciate that persistent sylvinite/carnallite beds cap a number of the beds of NaCl in The Nova Olinda Formation. During the late 1960s and 1970s, higher-resolution gamma-ray logging tools were used, along with better mud technology and associated narrower calliper measures. This work identified a number of (0.5 - 2m thick) layers of sylvinite, within the halites (Szatmari et al. 1979). For example, the fifth and seventh depositional cycles define isolated salt sub-basins that accumulated significant potash salts in Fazendinha and Arari regions (Figure 20). KCl contents of these beds are between 28-33% in beds some 2.47-2.65 m thick (Garrett, 1995). The average ore depth at Fazendinha, the larger of the known potash areas, is 1,050m (Figure 20). Much of the halite and potash distribution is controlled by the underlying rift-basin architecture (Figure 19b). Potential potash reserves poorly defined, but are interpreted to be large (Szatmari et al., 1979; Garrett, 1995).

Based on its texture, structure and chemistry, the potash intersection in the Amazon Basin is divided into three distinct zones, called informally, lower (milky or white sylvinite), middle (sulphates) and upper zones (red sylvinite) (Figure 20). The lower zone (milky-white sylvinite zone) contains sylvinite, with halite and subordinate intercalated kieserite and anhydrite beds. The lower potash zone is persistent within the basin and so covers an extensive area, whereas the upper potash zone is patchier. The greater extent of the lower potash zone is perhaps because it is the best isolated from any dissolution driven by circulation of undersaturated pore fluids through the overburden.

The middle zone is composed of a combination of sulphate and chloride salts and is informally termed the sulphate zone. It hosts a variety of K, Mg and sulphate minerals that include a number of hydrated salts. Typical mineral assemblages encompass sylvinite, sylvite, and langbeinite (K2SO4.2MgSO4) as well as the hydrated salts; polyhalite (K2SO4.2MgSO4.2CaSO4.2H2O), kainite (MgSO4.KCl.3H2O) and kieserite (MgSO4.H2O). The sulphate distribution in this unit changes from anhydrite and polyhalite in the west (Fazendinha) to langbeinite and kainite in the east (Faro area). Towards the basin centre, chloride beds replace marginal sulphate beds in the sulphate unit. A gradual increase in potash concentration from west to east is interpreted by Sad et al., 1982, as indicating the inflow direction was from the basin's western boundary.

The upper potash zone consists of coarsely-crystalline red sylvinite, with thin halite and anhydrite laminations. This level includes the best K2O grades drilled so far, averaging 23% K2O (between 33% to 16%). Red sylvinite is interpreted as a second generation product formed diagenetically by incongruent leaching of primary carnallite, but, as yet no carnallite (KCl.MgCl2.6H2O) has been identified in the upper unit.

The potash zone is overlain by impermeable coarsely-crystalline halite, with minor shale intercalations in a zone up to 25 m thick, in turn, overlain by impermeable shale beds some 20 m thick. It is underlain by an impervious, at times sparry, halite interval some 70m thick (Figure 20). At the time it was described (1970s-mid 1980s) little was known of the significance of halite crystal textures in terms of their primary versus diagenetic signatures. Such a study of the nature of the halites enclosing the potash zone in the Amazon basin would aid in the definition of an ore genesis model. We do know that a single potash zone does not extend across the basin. This is seen in a compilation of existing Petrobras wells in the Amazon Basin, which intersect the Nova Olinda Fm. Instead, potash salts accumulated in a series of sumps atop a persistent thick halite unit (Figure 20).

Elevated sulphate content in the potash zone of the Amazon Basin reflects the MgSO4-enriched nature of the world ocean during the Carboniferous. Potentially high levels of sulphate in proximity to adjacent sylvinite ore targets will complicate the processing of potential ore (see Warren 2016, Chapter 11). But in terms of supplying high levels of volatiles during sill intrusions, it is highly likely the various hydrated sulphate salts in the potash zone focused sill emplacement and contributed to elevated levels of halocarbons and sulphurous gases escaping into the earth's atmosphere at the end Triassic. As yet, no phreato-magmatic pipes have been documented in the Nova Olinda, but as the sourcing evaporite unit lie a kilometer beneath the surface and the dense tropical Amazon Jungle, this is not surprising. Increasing future use of STRM data may help solve this (Ibanez et alo., 2016)

 

Saline sediment-sill interaction

Sills from the Amazonas Basin have previously been described as low-Ti tholeiitic basalts and andesitic basalts De Min et al., 2003), and sills from both basins are generally characterised by a mineral assemblage of clinopyroxene, plagioclase, Fe–Ti oxides, rare olivine and orthopyroxene and accessory quartz-feldspar intergrowths. Recent studies report the presence of high-Ti sills in the eastern part of the Amazonas Basin (Figures 18, 21; Davies et al., 2017; Heimdal et al., 2018, 2019; Marzoli et al., 2018), but no high-Ti occurrences have been observed in the Solimões Basin. 


High-precision U–Pb dates from four dolerites from the Amazonas and Solimões basins overlap in age, with U–Pb ages for low-Ti dolerites of 201.525 ±0.065 (Amazonas Basin) and 201.470 ±0.089 (Solimões Basin), and for high-Ti dolerites in the Amazonas Basin of 201.477 ±0.062 and 201.364 ±0.023 Ma (Figure 18; Davies et al., 2017; Heimdal et al., 2018). This suggests that low-and high-Ti CAMP magmatism were active simultaneously, although low-Ti magmatism likely started earlier.

Detailed studies of CAMP sill geochemistry showing likely assimilation of chloride salts from the Nova Olinda evaporites are published in Heimdal et al., 2019, and summarised in this section. They show the bulk of e dolerites as sampled in the wells, illustrated in Figure 22, are characterised by phenocrysts of clinopyroxene and plagioclase in subophitic to intergranular textures, Fe–Ti oxides, and rare olivine and orthopyroxene. A different mineralogical assemblage (microphenocrysts of alkali-feldspar, quartz, biotite and apatite) is found in small independent domains, localised within the framework of coarser plagioclase and clinopyroxene laths. These fine-grained evolved domains crystallised in late-stage, evolved melt pockets in the interstitial spaces between earlier crystallised coarser grained crystals.


The majority of the studied dolerites are generally evolved tholeiitic basalts and basaltic andesites with low TiO2 concentrations (<2.0 wt.%). Four samples have high TiO2 concentrations (>2.0 wt.%), and are found in the eastern part of the Amazonas Basin (Figure 20a, c).

Whole-rock major and trace element and Sr-Nd isotope geochemistry of both low- and high-Ti sills is similar to that of previously published CAMP rocks from the two magma types. Low-Ti sills show enriched isotopic signatures (143Nd/144Nd201Ma from 0.51215 to 0.51244; 87Sr/86Sr201Ma from 0.70568 to 0.70756), coupled with crustal-like characteristics in the incompatible element patterns (e.g. depletion in Nb and Ta). Unaltered high-Ti samples show more depleted isotopic signatures (143Nd/144Nd201Ma from 0.51260 to 0.51262; 87Sr/86Sr201Maf from 0.70363 to 0.70398).

Low-Ti dolerites from both the Amazonas and Solimões basins contain biotite with extremely high Cl concentrations (up to 4.7 wt.%). They show that there is a strong correlation between host-rock lithology and Cl concentrations in biotite from the dolerites, and interpret this to reflect large-scale crustal contamination of the low-Ti magmas by halite-rich evaporites (Figure 21). The findings of Heimdal et al. (2019) support the hypothesis that sill-evaporite interactions increased volumes of volatile released during the emplacement of CAMP, and underlines the case for the active involvement of this LIP in the end-Triassic extinction event.


End-Cretaceous extinction event - Saline interactions driven by a bolide impact)

About 66 million years ago, at the end of the Cretaceous, one or possibly multiple large asteroids collided with the Earth. Paul Renne dated this impact at 66.043±0.011 million years ago on the Yucatan Peninsula, based on argon-argon dating (Renne, 2013). He went on to conclude that the main end-Cretaceous mass extinction event occurred within 32,000 years of this date. The bolide produced a crater some 150x180 km in diameter named the Chicxulub impact structure (Figure 23). Worldwide, a record of this event is evidenced by an iridium-enriched interval, in what is now called the Cretaceous-Tertiary Boundary Clay (KTBC) (Alvarez et al., 1980).

Other authors favouring additional bolide impacts at the end of the Cretaceous, such as Lerbekmo (2014) and Chaterjee (1997), have argued that some 40,000 years later, a much larger meteorite struck the shelf of the India-Seychelles continent, which was drifting northward in the southern Indian Ocean, producing a crater, some 450x600 km across, named the Shiva impact (Lerbekmo, 2014; Chaterjee, 1997). If a bolide-related feature, the Shiva crater was split by subsequent plate tectonism and today is not widely recognised by the scientific community as a K-T impact site.

As for any sound scientific hypothesis, there are ongoing arguments for the Chicxulub site being the "smoking gun" for the end-Cretaceous extinction event, many of these arguments and the supporting literature is discussed in (Kring, 2007). I shall focus on the saline geology of the Yucatan impact site, but recognise the arguments of some authors that the Shiva site is closely linked in time with the extrusion of the Deccan Traps. More importantly, voluminous Deccan Traps eruptions and intrusions had likely already degraded the end-Cretaceous atmosphere. A large bolide crashing into an anhydrite saltern in the palaeo-Gulf of Mexico was perhaps the coup de grâce for many already -stressed late Mesozoic communities (Wang et al., 2018)


Saline Geology of the Yucatan site

As it is covered by a Tertiary-age sediment carapace, there are no current evaporite outcrops on the Yucatan Peninsula. However, the region is underlain by thick Cretaceous anhydrite beds and has a nearby giant oil field, Cantarell, reservoired in a carbonate breccia trap possibly related to the impact (Grajales-Nishimura et al., 2000). Ongoing petroleum exploration means a number of exploration wells sample the Cretaceous geology of the Yucatan Peninsula (inset in Figure 24). Regionally, Cretaceous (Albian) saltern anhydrite beds extend from Guatemala, across the Yucatan Peninsula and north possibly to Veracruz. Depositionally similar, back-reef saltern beds typify the early Cretaceous (Albian) Ferry Lake Anhydrite, which extends across the onshore northern, and offshore eastern, Gulf of Mexico (Pittman, 1985; Petty, 1995; Loucks and Longman, 1982).

Pemex wells drilled on the Yucatan Peninsula, penetrate some 1300 –3500 m of bedded Tertiary, Cretaceous, and Jurassic strata (Figure 24; Ward et al., 1995). Palaeozoic metamorphic rocks are intersected at 2418 m in well Y4 and at 3202 m in well Y1. ‘‘Volcanic rock/andesite,’’ now broadly interpreted as an ‘‘impact-melt rock’’ or suevite is intersected in the lower parts of wells Y6 and C1. Based on the well geology there are seven major biostratigraphic-lithostratigraphic units in the Mesozoic section overlying basement rocks in the vicinity of the Chixulub impact site (Units A-F; Ward et al., 1995 and references therein). The regional depositional setting is typical of a Cretaceous carbonate platform, which at times became sufficiently isolated to deposit stacked anhydrite saltern beds in a rudistid back-reef setting (Warren, 2016; Chapter 5).

Unit A consists of red and grey sandstone, shale, and silty dolomite near the base of wells Y1, Y2, and Y4. This unit is Jurassic to Early Cretaceous in age (López Ramos, 1975).

Unit B is predominantly dolomite in its lower part, becomes rich in intercalated anhydrite and dolomite upward. Rock salt was cored in this unit in T1 at 2378–2381 m. Nummoloculina sp. was identified in Y2, suggesting an Albian age.

Unit C is predominantly shallow-water limestone in the lower part, becoming more dolomitic upward. At the base of unit C in wells Y1 and Y2 is a horizon with the large benthic orbitulinid foraminifer Dicyclina schlumbergeri? (Figure 23\4). Nummoloculina (N. heimi?) also occurs in the lower part of this unit in cores Y1 and Y4. Nearer the platform margin (Y4), the upper part of this unit contains a rudist limestone, but in other wells the rocks reflect more restricted depositional environments across the platform interior. Shallow-subtidal to intertidal dolomite makes up most of this section in Y5A, where anhydrite is interlayered with dolomite in the upper parts of the unit in Y1, Y2, and T1. The fossil assemblage indicates an Albian-Cenomanian age for unit C.

Unit D is predominantly somewhat deeper-subtidal limestone and marl, with horizons containing abundant tiny, mainly trochospiral planktic foraminifers as seen in samples from Y1, Y2-Y4, and Y5A.

Unit E consists of shallow-platform limestones with intervals containing abundant small planktic foraminifers. The unit contains rudist-bearing limestones considerd by López Ramos (1975) as Turonian, and a similar age is indicated by the presence of Marginotruncana pseudolinneiana and Dicarinella imbricata in samples from Y1, Y2, Y4, and Y5A.

Unit F consists of dolomitized shallow-platform limestone with benthic foraminifers. Abundant textularid and miliolid foraminifers are at the top of unit F (Fig. 2). The presence of Marginotruncana schneegansi and Globotruncana fornicata in well Y5A suggests a Santonian age for that part of this unit.

Unit G is a thick interval of breccia with abundant sand- to gravel-sized angular to subrounded fragments of dolostone, anhydrite, and minor limestones suspended in a dolomicrite matrix. The poorly-sorted fabric is similar to that of debris-flow deposits. López Ramos (1973) reported marl and limestone intercalations within the thick breccia from 1090 to 1270 m in well C1 (Figs. 1 and 2). In addition, Y4 and Y4 contain dolomite that may separate an upper breccia with rare or no planktic foraminifers from a lower breccia with abundant planktic foraminifers. Core in Y2 is composed of finely crystalline anhydrite, possibly also representing a less disturbed sedimentary layer or anhydrite block within the breccia interval.

Clasts of carbonate rocks in these breccias are fragments of many different kinds of dolostone and limestone, with different diagenetic histories. Anhydrite fragments typically make up 15%–20% of the breccia; much of the anhydrite is composed of tiny angular cleavage splinters. Some breccia layers contain grey-green fragments of altered volcanic ‘‘glass’’ and spherules. Other minor but significant constituents of the breccia are fragments of melt rock and basement as seen in Y6 (1295.5–1299 m), Y6 (1377–1379.5 m), and C1 (1393–1394 m). In addition, Hildebrand et al. (1991) found shocked quartz from Y6 (1208 –1211 m), and Sharpton et al. (1994) reported shock-deformed quartz and feldspar grains and melt inclusions in the dolomite-anhydrite breccia.

Planktic and benthic foraminifers are present in the breccia matrix and include Abathomphalus mayaroensis, Globotrun-canita conica, Rosita patelliformis, Pseudoguembelina palpebra, Racemiguembelina fructicosa, and Hedbergella monmouthensis, which indicate a late Maastrichtian (end-Cretaceous) age for formation of the breccia (Ward et al., 1995).

Climatic outcomes of the Yucatan impact

Widespread Jurassic anhydrites, hydrocarbon reservoirs and source rocks surround the Yucatan impact region; their vaporisation on bolide impact and rapid entry into the upper atmosphere added a good deal to the ensuing climatic mayhem (Figure 25) As we discussed for LIP emplacement, anhydrite decomposes at high temperatures, to form SO2 gas, CaO, and oxygen. Thermodynamic calculation and extrapolation using the free energy of formation of anhydrite and its reaction products as a function of temperature up to 1120°C (Robie et al., 1979), give an equilibrium pressure of 1 bar SO2 over the reaction:

2CaSO4 = 2CaO + 2S02 + O2

at a temperature around 1500°C (Brett, 1992; Yang and Ahrens, 1998). Experimental studies by Rowe et al. (1967) indicate that anhydrite decomposes in an open crucible above 1200°C. Temperatures higher than 1500°C are well in the range of temperatures of material subjected to strong shock in large bolide impacts, and at higher temperatures the equilibrium pressure would be considerably higher. Because the system is open, SO2 and oxygen would escape to the atmosphere as they did in the laboratory crucible of Rowe et al. (1967) and would continue to do so as long as post-impact temperatures were elevated.

Published discussions of the impact site geology all consider anhydrite as the evaporite mineralogy, with minor volumes of halite (well T1 in Figure 24). This lower salinity end of the evaporite series is typical of mega-sulphate settings, worldwide (Warren, 2016; Chapter 5) In addition, there is no evidence for hydrated potash salts in the region and this too is typical of starn salterns in a meg-sulphate basin. There is, however, the additional possibility that not all of the saltern gypsum had converted to anhydrite at the time of the impact. If so, this would have further detabilised and volatised the various lithologies at the site of the impact.


Intercalated carbonates, kerogens and other organic sediments at the collision site contributed additional CO2, CH4, H2O, and halocarbons to the atmosphere, as well as vast quantities of heat and particulates. The following discussion of the various contributors to climatic changes, driven by the Chicxulub impact, is taken mostly from Kring, 2007 (and contained references).

Acid rain; Because the Chicxulub impact occurred in a region with anhydrite, sulphurous vapour was injected into the stratosphere, producing sulphate aerosols and eventually sulphuric acid rain. Estimates of the amount of S liberated vary, consensus ranges from 7.5 × 1016 to 6.0 × 1017 g S, which would have produced 7.7 × 1014 to 6.1 × 1015 mol of sulphuric acid rain. In addition, the earth’s atmosphere was shock-heated by the impact event, producing nitric acid rain as well. Independent of the geology of the impact siter, the earth's atmosphere is heated when pierced by a bolide as the vapour-rich plume expands out from an impact site, and ejected debris rains through the atmosphere. In a Chicxulub-sized impact event, the ejecta debris is, estimated to produce ≈1×1014 mol of NOx in the atmosphere and, thus, ≈1×1015 moles of nitric acid rain. Impact-generated wildfires may have produced an additional ≈3×1015 mol of nitric acid. Sulphuric and nitric acid rain fell over a few months to a few years (Figure 25a).

Wildfires; Evidence of impact-generated fires is recovered from K/T boundary sequences worldwide in the form of fusinite pyrolitic polycyclic aromatic hydrocarbons, carbonised plant debris, and charcoal. The distribution of the fires is still poorly understood and may have had a restricted geographic distribution limited to the vicinity of the impact event, produced not by impact ejecta but by the direct radiation of the impact fireball which had a plasma core with temperatures over 10,000 °C. Several additional parameters influence the outcome (e.g., the trajectory of the impacting object, its speed, and mass of the ejecta). The amount of soot recovered from K/T boundary sediments (imply that the fires released ≈104 GT of CO2, ≈102 GT CH4 and 103 GT CO, which is equal to or larger than the amount of CO2 produced from vapourised target sediments. This likely had a severe effect on the global carbon cycle (Figure 25a).

Dust and aerosols in the atmosphere; Calculations suggest that dust and sulphate aerosols from the impact event, and soot from post-impact wildfires, caused surface temperatures to fall by preventing sunlight from reaching the surface where it was needed for photosynthesis. The base of the marine food chain, composed of photosynthetic plankton, collapsed. Slight increases or decreases in average water temperatures cannot extinguish photosynthetic plankton, nor the presence or absence of organisms higher up the food chain. Photosynthesisers are primarily affected by the availability of their energy source, light. Consequently, the loss of photosynthetic plankton following the Chicxulub impact event is evidence that sunlight was significantly blocked, whether it was by dust, soot, aerosols, or some other agent.

The timescale for particles settling through the atmosphere range from a few hours to approximately a year (Figure 25a, b). The time needed for the bulk of the dust to settle out of the atmosphere is ambiguous, however, because the size distribution of the dust is unclear. Some sites seem to be dominated by spherules ≈250 μm in diameter, which would have settled out of the atmosphere within hours to days. However, if there is a substantial amount of submicron material, then it may remain suspended in the atmosphere for many months. Soot, if it were able to rise into the stratosphere, would have taken similarly long times to settle. Soot that only rose into the troposphere, however, would have been flushed out of the atmosphere promptly by rain.

The dust, aerosols, and soot caused surface cooling after the brief period of atmospheric heating that immediately followed the impact. The magnitude of that cooling is unclear, however, because the opacity generated by the three components is uncertain and their lifetime in the atmosphere is also uncertain. Nonetheless, significant decreases in temperature of several degrees to a few tens of degrees have been proposed for at least short periods. Short-term cooling likely had a severe effect on the global carbon cycle, in what is popularly termed a “nuclear winter’ scenario (Figure 25).

Ozone destruction; Ozone-destroying Cl and Br is produced from the vaporised projectile, vaporised target lithologies, and biomass burning. Over five orders of magnitude more Cl than is needed to destroy today's ozone layer was injected into the stratosphere, compounded by the addition of Br and other reactants. The affect on the ozone layer may have lasted for several years, although it is uncertain how much of an effect it had on surface conditions. Initially, dust, soot, and NO2 may have absorbed ultraviolet radiation, and sulphate aerosols may have scattered the radiation. The settling time of dust was probably rapid relative to the time span of ozone loss, but it may have taken a few years for the aerosols to precipitate.

Greenhouse gases; Water and CO2 were produced from Chicxulub's target lithologies and the projectile, which could have potentially caused greenhouse warming after the dust, aerosols, and soot settled to the ground. Significant CO2, CH, and H2O were added to the atmosphere. Some of these components came directly from target materials. These include carbonates, which liberate CO2 when vaporised, and also includes hydrocarbons, the remainder of which has subsequently migrated into cataclastic dykes beneath the crater and impact breccias deposited along the Campeche Bank (e.g. Cantarell field). Water was liberated from the saturated sedimentary sequence and the overlying ocean (the lesser of the two sources).

The residence times of gases like CO2 are greater than those of dust and sulphate aerosols, so greenhouse warming may have occurred after a period of cooling. Estimates of the magnitude of the heating vary considerably, from an increase of global mean average temperature of 1 to 1.5 °C (based on estimates of CO2 added to the atmosphere by the impact) to ≈7.5 °C (based on measures of fossil leaf stomata).

Local and regional effects; The local and regional effects of the impact were enormous. Tsunamis radiated across the Gulf of Mexico, crashing onto nearby coastlines, and also radiated farther across the proto-Caribbean and Atlantic basins. Tsunamis were 100 to 300 m high when they crashed onto the gulf coast and ripped up sea floor sediments down to water depths of 500 m. The Gulf of Mexico region was also affected by the high-energy deposition of impact ejecta, density currents, and seismically-induced slumping of coastal sediments following magnitude 10 earthquakes. Tsunamis may have penetrated more than 300 km inland. The local landscape (both continental and marine) was buried beneath a layer of impact ejecta that was several hundred meters thick near the impact site and decreased with radial distance. Peak thicknesses along the crater rim may have been 600 to 800 m. Along the Campeche bank, 350 to 600 km from Chicxulub, impact deposits of ≈50 to ≈300 m are logged in the Cantarell boreholes.

Impact events also produce shock waves and air blasts that radiate across the landscape. Wind speeds over 1000 km/h are possible near the impact site, although they decrease with distance from the impact site. The pressure pulse and winds can scour soils and shred vegetation and any animals living in nearby ecosystems. Estimated radii of the area damaged by an air blast range from ≈900 to ≈1800 km.

Significant heat would have been another critical regional effect. Core temperatures in the plume rising from the crater were over 10,000° C, possibly high enough to generate fires out to distances of 1500 to 4000 km. The intense thermal pulse would have been relatively short-lived (5 to 10 min). Additional heating and spontaneous wildfires were ignited when impact ejecta fell through the atmosphere (3 to 4 days; Figure 25a).

The end-Cretaceous bolide impact had both short and long term effects on the Earth's climate and its atmospheric temperatures (Figure 25b). Over hours to days following impact, there was severe atmospheric heating as ejecta rained down through the atmosphere. This was following by a period of weeks to years of cooler temperatures as the atmosphere was polluted by SO2, NOx and soot from the impact preventing sunlight reaching the surface (nuclear winter scenario). Then, across time frame of decades to millennia, after the atmosphere cleared, increased CO2 levels drove a period of global warming. The legacy of the impact and the biotal recovery over the next few hundred thousand years is documented in a recent paper by Lowery et al., 2018. They showed that life reappeared in the basin just years after the impact and a high-productivity ecosystem was established within 30 kyr.

Extinction events intensified by heating evaporites

Evaporite salts are more chemically reactive at earth surface conditions than other sediments. Subsurface evaporites are prone to dissolution, alteration and reprecipition from the time they first precipitated and throughout their subsurface journeys in the diagenetic and metamorphic realms (Warren 2016). The same is true, but perhaps more so, if bedded salts are exposed to a heat source outside the normal geothermal gradient experienced in burial. Additional heat can come for the emplacement of igneous sills, magma bodies or the hot hydrothermal circulation it drives. Or it can come from near instantaneous heating to thousands of degrees associated with a bolide impact. Volatile products that result from this heating, as they enter the earth's atmosphere, can be inimical to life and include vast volumes of halocarbons, SO2, methane and CO2. Methane and CO2 come from kerogens and hydrocarbons stored in intercalated mudstones and limestones while volatilisation of carbonates can supply CO2.

The reactivity of evaporites and the vast volumes of volatiles released explains the intimate association of saline giants, heating and the three most devastating of the five major Phanerozoic extinction events.

Interestingly, two other events on the list of the "big five;" the Emeishan and late Devonian events (Figure 1) also have possible associations with heated evaporites. The Emeishan LIP intersects the edge of the anhydrite-rich Sichuan basin, while the 120km-diam., Late Devonian, Woodleigh bolide impacted the intracratonic Silurian Yaringa Fm. salts (including potash beds) on the coast of West Australia (SaltWork GIS database version 1.8 overlays, Chen et al., 2018; Glikson et al, 2005). But, before definitive conclusions can be made, more work is required to better tie down impact age, actual geographic extent of LIP emplacement, extent of evaporite breccias and evaporite volumes.

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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.

References

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Silica mobility and replaced evaporites: 3 - Archean cherts

John Warren - Sunday, August 28, 2016

Introduction

The two previous articles on silica mobility in evaporitic settings emphasised Phanerozoic examples and discussed silica textures largely tied to the replacement of sulphate evaporite nodules. This article will extend the time frame back to the Archean and also discuss scale controls on massive marine-derived evaporite beds in the early earth. The next article after this focuses on the Proterozoic. In order to extend our discussion into saline Precambrian successions, we must consider changes in ionic proportions and temperatures of the world’s oceans that this involves, and also include the background context of biological evolution of silica-extracting organisms.

Chert deposits clearly preserve a record of secular change in the oceanic silica cycle cross the Precambrian and the Phanerozoic (Maliva et al., 2005), with the chert nodule-evaporite association most obvious in alkaline brine-flushed areas in Phanerozoic sediments (previous 2 articles). Many silicified Phanerozoic evaporite examples co-occur with significant volumes of salts deposited in marine-fed megahalite and megasulphate basins. The evolutionary radiation of silica-secreting organisms across a deep time background is reflected in the transition from abiogenic silica deposition, characteristic of marine and nonmarine settings in the Archean and Proterozoic eons, to the predominantly biologically-controlled marine silica deposits of the Phanerozoic.

Silica levels in the Archean ocean

Estimated silica concentration in Precambrian seawater is 60 ppm SiO2 or more, while silica concentration of much of the modern ocean is controlled by silica-secreting organisms at values of 1 ppm or less to a maximum of 15 ppm (Perry and Lefticariu, 2014). There is no conclusive fossil evidence that such organisms were present in the Precambrian in sufficient abundance to have had a significant influence on the silica cycle, although some later Neoproterozoic protists likely had scales that were siliceous, and Ediacaran sponges certainly produced siliceous spicules. This contrasts with the Phanerozoic, during which the appearance of radiolaria and diatoms changed the locus of silica precipitation (both primary and replacement) from the peritidal and shallow shelf deposits characteristic of the Neoproterozoic, Mesoproterozoic, and much of the Paleoproterozoic, to the deep ocean biogenic deposits since the mid to late Phanerozoic. Comparative petrography of Phanerozoic and Precambrian chert shows an additional early change in nonbiogenic chert deposition occurred toward the end of the Paleoproterozoic era and was marked by the end to widespread primary and early diagenetic silica precipitation in normal marine subtidal environments (Table 1; ca. 1.8 Ga Maliva et al., 2005). Interestingly, the Precambrian transition corresponds to the onset of a plate tectonic regime resembling that of today (Stern, 2007). It was also the time when sulphate levels in the world’s oceans had risen to where gypsum became a primary marine evaporite, as evidenced by large silicified anhydrite nodules (with anhydrite relics) in the late Paleoproterozoic Mallapunyah Fm in the McArthur Basin, Australia (Warren, 2016). Paleoproterozoic early diagenetic “normal marine” cherts generally formed nodules or discontinuous beds within carbonate deposits with similar depositional textures. It seems these “normal marine” cherts formed primarily by carbonate replacement with subsidiary direct silica precipitation. In saline settings cauliflower cherts are also obvious from this time onwards.

 

Some of these Paleoproterozoic peritidal cherts were associated with iron formations and are distinctly different from younger cherts and appear to have formed largely by direct silica precipitation at or just below the seabed. These primary cherts lack ghosts or inclusions of carbonate precursors, have fine-scale grain fracturing (possibly from syneresis), exhibit low grain-packing densities, and are not associated with unsilicified carbonate deposits of similar depositional composition (Perry and Lefticariu, 2014). Cherts in some Paleoproterozoic iron formations (e.g., the Gunflint Formation, northwestern Lake Superior region) are composed of silica types similar to those in Phanerozoic sinters (e.g., the Devonian Rhynie and Windyfield chert sinters, Scotland, both of which preserved fine-scale cellular detail of Devonian plants, fungi and cyanobacteria, as well as elevated gold levels in the fault feeder system). Such “normal marine cherts lie outside the evaporite focus of this series of articles and for more detail the reader is referred to Perry and Lefticariu, 2014 and references therein.

Archean crustal tectonics and silicification of world-scale evaporites

Archean evaporites were not deposited as saline giants within subsealevel restricted basins created by sialic continent-to-continent proximity setting. In the greenstone terranes that typified the early Archean these tectonic settings simply could not yet exist (Warren, 2016, Chapter 2). Stern (2007) defines plate tectonics as the horizontal motion of Earth’s thermal boundary layer (lithosphere) over the convecting mantle (asthenosphere), and so it is a world-scale system or set of processes mostly driven by lithosphere sinking (subduction pull). He argues that the complete set of processes and metamorphic indicators, associated with modern subduction zones, only became active at the beginning of the Neoproterozoic (≈ 1 Ga). Stern interprets the older record to indicate a progression of tectonic styles from active Archaean tectonics and magmatism (greenstone belts), to something akin to modern plate tectonics at around 1.9 Ga (Figure 1). If so, then modern world-scale plate tectonics only began in the early Neoproterozoic, with the advent of deep subduction zones (blueschists) and associated powerful slab pull mechanisms. Flament et al. (2008) argue that the world’s continents were mostly flooded (mostly covered with shallow ocean waters) until the end of the Archaean and that only 2–3 % of the Earth’s area consisted of emerged continental crust by around 2.5 Ga (aka “water-world”).


It is very likely that the Archaean Earth’s surface was broken up into many smaller plates with volcanic islands and arcs in great abundance (greenstone terranes). Small protocontinents (cratons) formed as crustal rock was melted and remelted by hot spots and recycled in subduction zones. There were no large continents in the Early Archaean, and small protocontinents were probably the norm by the MesoArchaean, when the higher rate of geologic activity (hotter core and mantle) prevented crustal segregations from coalescing into larger units (Figures 1 and 3 ). During the Early-Middle Archaean, Earth’s heat flow was almost three times higher than it is today, because of the greater concentration of radioactive isotopes and the residual heat from the Earth’s accretion, hence the higher ocean temperatures (Figure 2; Eriksson et al. 2004). At that time of a younger cooling earth there was considerably greater tectonic and volcanic activity; the mantle was more fluid and the crust much thinner. This resulted in rapid formation of oceanic crust at ridges and hot spots, and rapid recycling of oceanic crust at subduction zones with oceanic water cycling through hydrothermally active zones somewhat more intensely than today (Zegers and van Keken 2001; Ernst 2009; Flament et al. 2008).


In the Pilbara craton region of Australia significant crustal-scale delamination occurred ≈ 3.49 Ga, just before the production of voluminous TTG (tonalite, trondhjemite, and granodiorite) melts between 3.48 and 3.42 Ga and the accumulation sonic evaporites (Figure 3; Zegers and van Keken 2001). Delamination resulted in rapid uplift, extension, and voluminous magmatism, which are all features of the 3.48–3.42 Ga Pilbara succession. As the delaminated portion was replaced by hot, depleted mantle, melts were produced by both decompressional melting of the mantle, resulting in high-MgO basalts (this is the Salgash Subgroup in the Pilbara craton), and melting of the gabbroic and amphibolitic lower crust, so producing TTG melts. Partial melting of the protocrust to higher levels can be envisaged as a multistep process in which heat was conducted to higher levels and advection of heat occurs by intrusion of partial melts in subsequently higher levels (indicated by purple arrows in Figure 3). TTG melt products that were first intruded were subsequently metamorphosed and possibly partially melted, as can be inferred from the migmatitic gneisses of the Pilbara. This multistep history explains the complex pattern of U-Pb zircon ages of gneisses and granodiorites found within the Pilbara batholiths and the range in geochemical compositions of the Pilbara TTG suite.


Key to the formation of early Archaean evaporites, which indicate a sodium bicarbonate ocean at that time (see next section), is the observation that crustal delamination and the creation of TTG melts led to up to 2 km of crustal uplift (Figure 3). This would have driven some regions of what were submarine sedimentary systems into suprasealevel positions in the Archean waterworld, so creating the potential for hydrographically-isolated subsealevel marine seepage sumps in those portions of the uplifted crust above the zones of delamination. It also explains the centripetal nature of much shallow marine sedimentation of that time. This is cardinal at the broad tectonic scale when comparing the distribution of Archaean and Phanerozoic evaporites (Warren, 2016). Most Archaean evaporite are remnants that are pervasively silicified and underlain by layered igneous complexes, which were dominant across the greenstone seafloor and are associated with bottom-nucleated baryte beds tied to hydrothermal seeps.

Felsic protocontinents (suprasealevel cratons) hosting silicified evaporite remnants probably formed atop Archaean hot spots from a variety of sources: mafic magma melting more felsic rocks, partial melting of mafic rock, and from the metamorphic alteration of felsic sedimentary rocks. Although the first continents formed during the Archaean, rock of this age makes up only 7% of the world’s current cratons; even allowing for erosion and destruction of past formations, evidence suggests that only 5–40 % of the present volume continental crust formed during the Archaean. 

Archean oceans and silicified sodic evaporites 

Chert styles and occurrences in saline settings across deep time clearly show that we cannot carry Phanerozoic silica mobility models in saline lacustrine or CaSO4 evaporite associations directly across time into the deep Precambrian. Rather, comparisons must be made in a context of the evolution of the earth’s atmosphere and associated ocean chemistry, both of which are in part related to the earth's tectonic evolution.

Levels of early Archaean sulphate in the world ocean were probably less than a few percent of the current levels and probably remained so until the evolution of an oxygen-reducing biota into the Proterozoic (Habicht and Canfield 1996; Kah et al. 2004; Warren, 2016). Grotzinger and Kasting (1993) argue that high levels of atmospheric CO2 meant HCO3/Ca ratios were much higher in the Archaean and the Paleoproterozoic oceans than today. All the calcium in seawater was deposited as marine cement-stones and other alkaline earth precipitates well before bicarbonate was depleted and there was no Ca left over to precipitate as gypsum. The early Archaean waterworld ocean was likely a Na–Cl–HCO3 sea, and not the Na–Cl ocean of today (Kempe and Degens 1985; Maisonneuve 1982). This early Archaean hydrosphere had a chemistry similar to that found in modern soda lakes like Lake Magadi and Lake Natron (pathway I brines) and hence the term “soda-lake oceans.” This rather different marine brine chemistry would have precipitated halite and trona/nahcolite, not halite/gypsum. It probably meant that if gypsum/anhydrite did ever precipitate directly from evaporating Archaean seawater it did so only in minor amounts well after the onset of halite precipitation.

 

The case for nahcolite (NaHCO3) as a primary evaporite (Figure 4a-d), along with halite, in the 3.42 Ga rocks of the Barberton greenstone belt was first documented by Lowe and Fisher-Worrell,1999), both the nahcolite and the halite are silicified. Beds of these silicified sodic evaporite define 5 types of precipitates: (1) large, pseudohexagonal prismatic crystals as much as 20 cm long that increase in diameter upward; (2) small isolated microscopic pseudohexagonal crystals; (3) small, tapering-upward prismatic crystals as much as 5 cm long; (4) small acicular crystallites forming halos around type 1 crystals; and (5) tightly packed, subvertical crystal aggregates within which individual crystals cannot be distinguished. Measurement of interfacial angles between prism and pinacoid faces on types 1 and 2 crystals show four interfacial angles of about 63° and two of about 53°. The morphologies and interfacial angles of these crystals correspond to those of nahcolite, NaHCO3 (Figure 4e). There is no clear evidence for the presence of gypsum in these beds. Sugitani et al. (2003) reported silicified nahcolite (the high CO2 form of sodium carbonate salts; see Warren, 2016, chapter 2) in ≈ 3.2 Ga rocks in the northern part of the Eastern Pilbara block, Western Australia (Figures 4, 5). Coarse, upward-radiating, silicified evaporite crystals in the ca. 3.47–3.46 Ga Strelley Pool Chert (Lowe 1983) show the same habit, geometry, and environmental setting as silicified nahcolite pseudomorphs in the Kromberg Fm. in the Barberton belt, South Africa, and also probably represent silicified NaHCO3 precipitates (Lowe and Tice 2004). Depositional reconstructions in both regions imply a strong hydrothermal association to the silicification of the evaporites in both regions as do bottom-nucleated baryte layers that define seafloor seeps fed by hydrothermal waters moving up faults (Figure 4f; Nijman et al., 1999; van den Boorn et al., 2007).

The pervasive presence of type 1 brines as ocean waters in the early Archean, along with elevated silica levels in most surface ocean waters, compared to the Phanerozoic, implies a significant portion of Archean cherts may also have had a volcanogenic sodium silicate precursor, much like the silicification seen in the modern African rift valley lakes (Eugster and Jones, 1968 and article 1 in this series of articles on silica mobilisation). So in order to decipher possible evaporite-silicification associations we must include aspects of hydrothermal fluid inherent to the Archean, as well as the likely higher surface temperatures that typified highly reducing (anoxic) waters of the early Archean ocean (Figure 3).


Archean evaporite deposition and silicification

Worldwide, the most widespread Archaean depositional environment, especially in early Archaean greenstone terranes, was the mafic plain environment (Condie 2016; Lowe 1994). In this setting, large volumes of basalt and komatiite were erupted to form widespread mostly submarine mafic plains characteristic by ubiquitous pillow structures in the lava interlayers. A second significant sedimentary environment was a deepwater, nonvolcanic setting, where chemical and biochemical cherts, banded iron formation, and carbonate laminites were deposited. The typical lack of evaporite indications in these mostly deepwater sediments indicates an ongoing lack of hydrologic restriction while the sediments were accumulating (waterworld association). The third association, a greywacke-volcanic association becomes more widespread in later Archaean greenstones, which typically sit stratigraphically atop mafic plain units. This association is composed chiefly of greywackes and interbedded calc-alkaline volcanics, hydrothermal precipitates and, in some shallower parts, silicified evaporites. It was perhaps mostly an island arc system and dominantly more open marine as it typically lacks widespread indicators of former marine evaporites. However, more locally it also preserves fluvial and shallow-marine detrital sediments, that were probably deposited locally in Archaean pull-apart basins, and associated with mineralogically mature sediments (quartzarenites, etc.). These more continental associations typified the shallowest to emergent parts of these continental rifts.

Unlike the other two early Archean  greenstone terranes this third terrane type can in places, such as the Pilbara, be tied to sedimentary indicators of a surfacing seafloor, indicated by particular chert and volcaniclastic layers showing mud cracks, wave ripples, tidalites interbedded with hyaloclastics, vuggy cherts, banded iron formations, carbonates and thick now-dissolved and altered type 1 evaporite masses (breccias), perhaps residues of beds formerly dominated by sodium carbonate and halite salts (Figure 5). The Warrawoona Group, preserves many such silicified examples that retain fine detail of primary textures such as mud cracks, oolites, and evaporite crystal casts and pseudomorphs, all indicating shallow-water to emergent deposition atop the mafic plain. In terms of crystal outlines there few if any casts of possible gypsum crystals, more typically, they indicate bladed pseudo-hexagonal, bottom-nucleated nahcolite, trona and in some instances, halite pseudomorphs (Figure 4).

Depositionally, to acquire the needed high salinities, these cherty evaporite units must have risen, at least locally, to shallow near-sealevel depths and at time become emergent, allowing local hydrographically-isolated lacustrine/rift evaporite subaqueous deposition or precipitation of local seepage drawdown salts. Associated primary-textured carbonate and baryte layers interbedded with the cherts are typically minor, bottom-nucleated baryte textures that may likely indicate hydrothermal vent deposits (Figure 4f; Nijman et al., 1999).

Inherent high solubility of any sodium bicarbonate and/or halite salts in what was a hotter burial system, more strongly influenced by hydrothermal circulation than today, meant most of the original sodic evaporite salts were not preserved, unless silicified in early burial. But their presence as silicified pseudomorphs in less-altered greenschist terranes intercalated with volcanics (Figure 4), such as in the Yilgarn, Pilbara and Kaapvaal cratons, clearly shows two things; (1) at times in the early Archaean waterworld there was sufficient hydrographic restriction to allow marine sodian carbonate and sodian chloride evaporites to form and (2) this marine restriction/seepage inflow was probably driven by ongoing volcanism and associated uplift, with evaporites restricted to particular basinwide stratigraphic indicator levels. In the East Pilbara, the early Archaean evaporite stratigraphic level is the Strelley Pool chert, in the Warrawoona group (Figure 5). This is also the level with some of the earliest indications of cellular life-forms (Wacey 2009).

For the original sodic evaporites, it marks the hydrological transition from open marine seafloor to a restricted hydrographically-isolated marine-fed sump basin, surrounded by granite-cored highs with the required uplift likely driven by delamination at the level of the mantle transition (Figures 1 and 3). Given the intimate association of chemical sediments to volcanism in early Archaean greenstone basins, and the sodium bicarbonate ocean chemistry then, compared to the Phanerozoic evaporite hydrochemistries, we can expect a higher proportion of CO2 volatilisation, a higher boron content (tourmalinites) in early Archaean, and a higher level of silicification.

Is the present the key to the past?

The study of silicified evaporites and associated sediments, formed in the early stages of the Earth’s 3.5 Ga sedimentary record, shows that not only has ocean chemistry evolved (see August 24, 2014 blog), the earth’s lithosphere/ plate tectonic character has also evolved (Eriksson et al. 2013). The further back in time, the less reliable is the application of the current plate tectonic paradigm with its strongly lateral movements of crustal blocks and associated plate-scale evaporite basin controls. Phanerozoic evaporites, and the associated silicified sulphate nodules, define a marine-fed seep system where subsealevel continental rifts and continent-continent collision belts favour the formation of mega-evaporite basins (Warren, 2010). Instead, in a substantial portion of the earlier part of the 2 billion year earth history that is the Archaean, shows early-earth evaporite deposition was favored by hydrographic isolation created by strong vertical movement of earth’s crust related to upwelling mantle plumes and crustal delamination with more intense hydrothermal circulation and silicification. There is still no real consensus as to actual time when plate tectonics, as it operates today, actually began, but there is consensus that the present, in terms of plate tectonics, plate-edge collision and evaporite distribution, is not the key to much of the Archaean (Stern 2007; Rollinson 2007).

Uplift and the local accumulation of sodium carbonate Archean evaporites occurred in a depositional setting that was dominated by volcaniclastics,hydrothermal vents and extensional tectonics. Tectonic patterns in these settings have a strongly vertical flavor. In contrast, Phanerozoic salts formed from marine waters with a NaCl dominance with minor bicarbonate compared to calcium, and located mostly in subsealevel sumps formed at interacting sialic plate margins where the dominant tectonic flavor is driven the lateral movement of plates atop a laterally moving asthenosphere and the relative proportion of vilified salts is lower.

Whatever and wherever the onset of Archaean evaporite deposition, all agree that the mechanisms and aerial proportions world-scale plate tectonics were different in early earth history compared to the Phanerozoic. The current argument as to how different is mostly centred on when earth-scale plate tectonic processes became similar to those of today. Given much higher crustal heat flows, it is likely that hydrographically isolated subsealevel depressions, required to form widespread marine evaporites were more localized in the Archaean than today and were more susceptible to hydrothermal alteration, metamorphism and silicification. Appropriate restricted brine sumps would have tended to occur in magmatically-induced uplift zones atop incipient sialic segregations, with crestal subsealevel grabens, which were hydrographically isolated by their surrounds created by supra-sealevel uplift. Once deposited, the higher heat flow in Archaean crust and mantle would also have meant any volumetrically significant evaporites masses were more rapidly recycled, silicified and replaced via diagenetic and metamorphic processes than today.

Some authors have noted that there are no widespread marine evaporites in the Archaean and in the sense of actual preserved salts, this is true. But when one considers that the Archaean crust was much hotter than today and hydrothermal circulation was more active and pervasive, then widespread burial preservation of the primary salts seems highly unlikely. Even in the Neoproterozoic, lesser volumes of the original salt masses remain (Hay et al. 2006). The lack of preserved salts in earlier Precambrian strata is perhaps more a matter of great age, polycyclic metamorphic alteration and the typical proximity to shallow hydrothermal fluids in emergent evaporite forming regions of the Archean waterworld. However we must also ask if the onset of modern styles of plate tectonics also played a role in the relative absence of preserved saline giants in strata older than 1Ga, In the next article we shall look how cooling and the onset of sialic plate tectonics similar to today, altered the types, styles and distributions of silicified and other evaporite salts as the world's oceans moved toward a chemistry more akin to that of today.

References

 

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Lowe, D. R., and M. M. Tice, 2004, Geologic evidence for Archean atmospheric and climatic evolution: Fluctuating levels of CO2, CH4, and O2 with an overriding tectonic control: Geology, v. 32, p. 493-496.

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Silica mobility and replaced evaporites: 2 - replaced CaSO4

John Warren - Sunday, July 31, 2016

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

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

Types and traits of authigenic silica and cherts

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

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


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

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


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

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

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

 

Silicification of calcium sulphate nodules and isolated crystals

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


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

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


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

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


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

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

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


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

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

 

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

  

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

  

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

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


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

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

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

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

John Warren - Friday, April 29, 2016

 

Introduction

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

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

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

 

The Rift

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

The Salt

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


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

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


The Deeps

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


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

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


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


Brine Chemistry in Red Sea DHALS

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


Mineralization in Red Sea DHALS

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

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

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


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

 

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

 

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

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

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

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

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

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

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

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

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

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

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

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

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

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

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

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

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


 

 

 

 

 

 


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