Salty Matters

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

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

Avila-Salinas, W., 1990. Origin of the Copper Ores at Corocoro, Bolivia. In: L. Fontboté, G.C. Amstutz, M. Cardozo, E. Cedillo and J. Frutos (Editors), Stratabound Ore Deposits in the Andes. Special Publication No. 8 of the Society for Geology Applied to Mineral Deposits. Springer Berlin Heidelberg, pp. 659-670.

Bonatti, E., Fisher, D.E., Joensuu, O., Rydell, H.S. and Beyth, M., 1972. Iron-manganese-barium deposit from the north Afar rift (Ethiopia). Economic Geology, 67(6): 717-730.

Bull, S.W., 1998. Sedimentology of the Palaeoproterozoic Barney Creek formation in DDH BMR McArthur 2, southern McArthur basin, northern territory. Australian Journal of Earth Sciences: An International Geoscience Journal of the Geological Society of Australia, 45(1): 21-31.

Cooke, D.R., Bull, S.W., Donovan, S. and Rogers, J.R., 1998. K-metasomatism and base metal depletion in volcanic rocks from the McArthur basin, northern territory - Implications for base metal mineralization. Economic Geology, 93(8): 1237-1263.

Cox, D.P., Lindsey, D.A., Singer, D.A., Moring, B.C. and Diggles, M.F., 2007. Sediment-Hosted Copper Deposits of the World: Deposit Models and Database. USGS Open File Report 03-107, Version 1.3 (Available online at http://pubs.usgs.gov/of/2003/of03-107/.

Davidson, G.J., 1998. Alkali alteration styles and mechanisms, and their implications for a brine factory source of base metals in the rift-related McArthur Group, Australia. Australian Journal of Earth Sciences, 45(1): 33-49.

Davidson, G.J., 1999. Feldspar metasomatism along a Proterozoic rift-basin margin - "Smoke" around a base-metal "fire" (HYC deposit, Australia) or a product of background diagenesis? Geological Society of America Bulletin, 111(5): 663-673.

Davidson, G.J. and Dashlooty, S.A., 1993. The Glyde Sub-basin - A volcaniclastic-bearing pull-apart basin coeval with the McArthur River base-metal deposit, Northern Territory. Australian Journal of Earth Sciences, 40(6): 527-543.

Eldridge, C.S., Williams, N. and Walshe, J.L., 1993. Sulfur isotope variability in sediment-hosted massive sulfide deposits as determined using the ion microprobe SHRIMP: II. A study of the H.Y.C. deposit at McArthur River, Northern Territory, Australia. Economic Geology, 88(1): 1-26.

Entwistle, L.P. and Gouin, L.O., 1955. The chalcocite-ore deposits at Corocoro, Bolivia. Economic Geology, 50(6): 555-570.

Eugster, H.P., 1989. Geochemical environments of sediment-hosted Cu-Pb-Zn deposits. In: R.W. Boyle, A.C. Brown, C.W. Jefferson and E.C. Jowett (Editors), Sediment hosted stratiform copper deposits. Geological Association of Canada, Special Paper, pp. 111-126.

Flint, S., 1986. Sedimentary and diagenetic controls on red-bed ore genesis; the middle Tertiary San Bartolo copper deposit, Antofagasta Province, Chile. Economic Geology, 81(4): 761-778.

Flint, S.S., 1989. Sediment-hosted stratabound copper deposits of the Central Andes. Geological Association of Canada Special Paper, 36: 371-398.

Fusswinkel, T., Wagner, T., Wälle, M., Wenzel, T., Heinrich, C.A. and Markl, G., 2013. Fluid mixing forms basement-hosted Pb-Zn deposits: Insight from metal and halogen geochemistry of individual fluid inclusions. Geology, 41(6): 679-682.

Groves, D., I. and Bierlein, F., P. , 2007. Geodynamic settings of mineral deposit systems. Journal of the Geological Society, 164: 19-30.

Groves, D.I., Condie, K.C., Goldfarb, R.J., Hronsky, J.M.A. and Vielreicher, R.M., 2005. Secular changes in global tectonic processes and their influence on the temporal distribution of gold-bearing mineral deposits. Economic Geology, 100(2): 203-224.

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

Hinman, M., 1995. Base metal mineralisation at McArthur River: structure and kinematics of the HYC-Cooley zone at McArthur River. Australian Geological Survey Organisation, Record,1995/5.

Hitzman, M.W., Selley, D. and Bull, S., 2010. Formation of Sedimentary Rock-Hosted Stratiform Copper Deposits through Earth History. Economic Geology, 105(3): 627-639.

Holman, A.I., Grice, K., Jaraula, C.M.B. and Schimmelmann, A., 2014. Bitumen II from the Paleoproterozoic Here’s Your Chance Pb/Zn/Ag deposit: Implications for the analysis of depositional environment and thermal maturity of hydrothermally-altered sediments. Geochimica et Cosmochimica Acta, 139: 98-109.

Ireland, T., Bull, S.W. and Large, R.R., 2004a. Mass flow sedimentology within the HYC Zn-Pb-Ag deposit, Northern Territory, Australia: evidence for syn-sedimentary ore genesis. Mineralium Deposita, 39(2): 143-158.

Ireland, T., Large, R.R., McGoldrick, P. and Blake, M., 2004b. Spatial distribution patterns of sulfur isotopes, nodular carbonate, and ore textures in the McArthur River (HYC) Zn-Pb-Ag deposit, northern territory, Australia. Economic Geology, 99(8): 1687-1709.

Large, R.R., Bull, S.W., Cooke, D.R. and McGoldrick, P.J., 1998. A genetic model for the HYC deposit, Australia: Based on regional sedimentology, geochemistry, and sulfide-sediment relationships. Economic Geology, 93(8): 1345-1368.

Ljunggren, P. and Meyer, H.C., 1964. The copper mineralization in the Corocoro basin, Bolivia. Economic Geology, 59: 110-125.

Logan, G.A., Hinman, M.C., Walter, M.R. and Summons, R.E., 2001. Biogeochemistry of the 1640 Ma McArthur River (HYC) lead–zinc ore and host sediments, Northern Territory, Australia. Geochimica Cosmochimica Acta, 65: 2317-2336.

Logan, R.G., 1979. The Geology and mineralogical zoning of the HYC Ag-Pb-Zn deposit, McArthur River, NT Masters Thesis, Australian National University, Cannberra, Australia.

Logan, R.G., Murray, W.J. and Williams, N., 1990. HYC silver-lead-zinc deposit, McArthur River. In: F.E. Hughes (Editor), Geology of the mineral deposits of Australia and Papua New

Guinea. Monograph Series - Australasian Institute of Mining and Metallurgy, pp. 907-911.

McConachie, B.A. and Dunster, J.N., 1998. Regional stratigraphic correlations and stratiform sediment-hosted base-metal mineralisation in the northern Mt Isa Basin. Australian Journal of Earth Sciences: An International Geoscience Journal of the Geological Society of Australia, 45(1): 83-88.

Meyer, C., 1988. Ore deposits as guides to geologic history of the Earth. In: G.W. Wetherill and et al. (Editors), Annual review of earth and planetary sciences. Vol. 16. Annual Reviews Inc., pp. 147-171.

Mohr, M., Warren, J.K., Kukla, P.A., Urai, J.L. and Irmen, A., 2007. Subsurface seismic record of salt glaciers in an extensional intracontinental setting (Late Triassic of northwestern Germany). Geology, 35(11): 963-966.

Murray, W.J., 1975. McArthur River HYC lead-zinc-silver and related deposits. In: C.L. Knight (Editor), Economic geology of Australia and Papua New Guinea—metals. Australasian Institute of Mining and Metallurgy, Melbourne, pp. 329-338.

Perkins, W.G. and Bell, T.H., 1998. Stratiform replacement lead-zinc deposits: A comparison between Mount Isa, Hilton, and McArthur River. Economic Geology, 93(8): 1090-1212.

Pirajno, F. and Bagas, L., 2008. A review of Australia's Proterozoic mineral systems and genetic models. Precambrian Research, 166(1-4): 54-80.

Rose, A.W., 1976. The effect of cuprous chloride complexes in the origin of red-bed copper and related deposits. Economic Geology, 71(6): 1036-1048.

Ruffell, A.H., Moles, N.R. and Parnell, J., 1998. Characterisation and prediction of sediment-hosted ore deposits using sequence stratigraphy. Ore Geology Reviews, 12(4): 207-223.

Rutland, R.W.R.R., 1966. An unconformity in the Corocoro basin, Bolivia, and its relation to the copper mineralization. Economic Geology, 61: 962-964.

Savrda, C.E., Cook, R.B. and Petrov, A., 2006. Trace Fossil Preservation by Native Copper, Corocoro, Bolivia. Rocks & Minerals, 81(5): 362-363.

Schmidt, M., Al-Farawati, R. and Botz, R., 2015. Geochemical classification of brine-filled Red Sea Deeps. In: N.M.A. Rasul and I.C.F. Stewart (Editors), The Red Sea. Springer, Berlin, pp. 219-233.

Shockey, P.N. and Renfro, A.R., 1974. Copper-silver solution fronts at Paoli, Oklahoma. Economic Geology, 69: 266-268.

Singewald, J.T. and Berry, E.W., 1922. The geology of the Corocoro copper district of Bolivia. Johns Hopkins University studies in geology -- No. 1.. 117 p.

Walker, R.N., Gulson, B. and Smith, J., 1983. The Coxco deposit - a Proterozoic Mississippi Valley-type deposit in the McArthur River district, Northern Territory, Australia. Economic Geology, 78(2): 214 - 249.

Walker, R.N., Logan, R.G. and Binnekamp, J.G., 1977. Recent geological advances concerning the H.Y.C. and associated deposits, McArthur river, N.Y. Journal of the Geological Society of Australia, 24(7-8): 365-380.

Warren, J.K., 2000. Evaporites, brines and base metals: low-temperature ore emplacement controlled by evaporite diagenesis. Australian Journal of Earth Sciences, 47(2): 179-208.

Warren, J.K., 2010. Evaporites through time: Tectonic, climatic and eustatic controls in marine and nonmarine deposits. Earth-Science Reviews, 98(3-4): 217-268.

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

Warren, J.K. and Kempton, R.H., 1997. Evaporite Sedimentology and the Origin of Evaporite-Associated Mississippi Valley-type Sulfides in the Cadjebut Mine Area, Lennard Shelf, Canning Basin, Western Australia. In: I.P. Montanez, J.M. Gregg and K.L. Shelton (Editors), Basinwide diagenetic patterns: Integrated petrologic, geochemical, and hydrologic considerations. SEPM Special Publication, Tulsa OK, pp. 183-205.

Wilkinson, J.J., Stoffell, B., Wilkinson, C.C., Jeffries, T.E. and Appold, M.S., 2009. Anomalously Metal-Rich Fluids Form Hydrothermal Ore Deposits. Science, 323(5915): 764-767.

Williams, N., 1978. Studies of base metal sulphide deposits at McArthur River, Northern Territory, Australia I. The Cooley and Ridge Deposits. Economic Geology, 73(6): 1005 - 1035.

Williams, N. and Logan, R.G., 1986. Geology and evolution of the H.Y.C. stratiform Pb-Zn orebodies, Australia. Stanford Univ.Pub. Geol. Sci., 20: 57-60.

Wu, C., 2008. Bayan Obo Controversy: Carbonatites versus Iron Oxide-Cu-Au-(REE-U). Resource Geology, 58(4): 348-354.

 

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

 

Buick, R., and J. S. R. Dunlop, 1990, Evaporitic sediments of early Archaean age from the Warrawoona Group, North Pole, Western Australia: Sedimentology, v. 37, p. 247-277.

Buick, R., J. R. Thornett, N. J. McNaughton, J. B. Smith, M. E. Barley, and M. Savage, 1995, Record of emergent continental crust ≈3.5 billion years ago in the Pilbara Region: Nature, v. 375, p. 574 - 777.

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

Eriksson, P. G., W. Altermann, D. R. Nelson, W. U. Mueller, and O. Catuneanu, 2004, The Precambrian Earth - Tempos and Events: Developments in Precambrian Geology, Elsevier, 941 p.

Eugster, H. P., and B. F. Jones, 1968, Gels Composed of Sodium-Aluminum Silicate, Lake Magadi, Kenya: Science, v. 161, p. 160-163.

Flament, N., N. Coltice, and P. F. Rey, 2008, A case for late-Archaean continental emergence from thermal evolution models and hypsometry: Earth and Planetary Science Letters, v. 275, p. 326-336.

Grotzinger, J. P., and J. F. Kasting, 1993, New constraints on Precambrian ocean composition: Journal of Geology, v. 101, p. 235-243.

Habicht, K. S., and D. E. Canfield, 1996, Sulphur isotope fractionation in modern microbial mats and the evolution of the sulphur cycle: Nature, v. 382, p. 342-343.

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

Kah, L. C., J. K. Bartley, T. D. Frank, and T. W. Lyons, 2006, Reconstructing sea-level change from the internal architecture of stromatolite reefs: an example from the Mesoproterozoic Sulky Formation, Dismal Lakes Group, arctic Canada: Canadian Journal of Earth Sciences, v. 43, p. 653-669.

Kempe, S., and E. T. Degens, 1985, An early soda ocean?: Chemical Geology, v. 53, p. 95-108.

Lowe, D. R., 1983, Restricted shallow-water sedimentation of early Archean stromatolitic and evaporitic strata of the Strelley Pool Chert, Pilbara Block, Western Australia: Precambrian Research, v. 19, p. 239-283.

Lowe, D. R., 1994, Archean greenstone-related sedimentary rocks, in K. C. Condie, ed., Archean Crustal Evolution: Amsterdam, Elsevier, p. 121-170.

Lowe, D. R., and G. Fisher-Worrell, 1999, Sedimentology, mineralogy, and implications of silicified evaporites in the Kromberg Formation, Barberton Greenstone Belt, South Africa, in D. R. Lowe, and G. R. Byerly, eds., Geologic evolution of the Barberton Greenstone Belt, South Africa, Geological Society of America Special Paper, v. 329, p. 167-188.

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.

Maisonneuve, J., 1982, The composition of the Precambrian ocean waters: Sedimentary Geology, v. 31, p. 1-11.

Maliva, R. G., A. H. Knoll, and B. M. Simonson, 2005, Secular change in the Precambrian silica cycle: Insights from chert petrology: Geological Society of America Bulletin, v. 117, p. 835-845.

Nijman, W., K. H. de Bruijne, and M. E. Valkering, 1999, Growth fault control of Early Archaean cherts, barite mounds and chert-barite veins, North Pole Dome, Eastern Pilbara, Western Australia: Precambrian Research, v. 95, p. 245-274.

Perry, E. C. J., and L. Lefticariu, 2014, Formation and Geochemistry of Precambrian Cherts, in H. D. Holland, and K. K. Turekian, eds., Treatise on Geochemistry (2nd edition), Elsevier, p. 113-139.

Robert, F., and M. Chaussidon, 2006, A palaeotemperature curve for the Precambrian oceans based on silicon isotopes in cherts: Nature, v. 443 (7114), p. 969-972.

Rollinson, H., 2007, When did plate tectonics begin?: Geology Today, v. 23, p. 186-191.

Stern, R., 2007, When and how did plate tectonics begin? Theoretical and empirical considerations: Chinese Science Bulletin, v. 52, p. 578-591.

Sugitani, K., K. Mimura, K. Suzuki, K. Nagamine, and R. Sugisaki, 2003, Stratigraphy and sedimentary petrology of an Archean volcanic-sedimentary succession at Mt. Goldsworthy in the Pilbara Block, Western Australia: implications of evaporite (nahcolite) and barite deposition: Precambrian Research, v. 120, p. 55-79.

Tänavsuu-Milkeviciene, K., and J. F. Sarg, 2015, Sedimentology of the World Class Organic-Rich Lacustrine System, Piceance Basin, Colorado, in M. E. Smith, and A. R. Carroll, eds., Stratigraphy and Paleolimnology of the Green River Formation, Western USA: New York, Springer, p. 153-182.

van den Boorn, S. H. J. M., M. J. van Bergen, W. Nijman, and P. Z. Vroon, 2007, Dual role of seawater and hydrothermal fluids in Early Archean chert formation: Evidence from silicon isotopes: Geology, v. 35, p. 939-942.

Van Kranendonk, M. J., R. Hugh Smithies, A. H. Hickman, and D. C. Champion, 2007, Review: secular tectonic evolution of Archean continental crust: interplay between horizontal and vertical processes in the formation of the Pilbara Craton, Australia: Terra Nova, v. 19, p. 1-38.

Wacey, D., 2009, Early Life on Earth: A Practical Guide: Topics in Geobiology, 31, Springer.

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

Warren, J. K., 2016, Evaporites: A compendium (ISBN 978-3-319-13511-3) Published Feb. 22, 2016: Berlin, Springer, 1854 p.

Zegers, T. E., and P. E. van Keken, 2001, Middle Archean continent formation by crustal delamination: Geology, v. 29, p. 1083-1086.


 

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.

References

Alonso-Zarza, A. M., Y. Sánchez-Moya, M. A. Bustillo, A. Sopeña, and A. Delgado, 2002, Silicification and dolomitization of anhydrite nodules in argillaceous terrestrial deposits: an example of meteoric-dominated diagenesis from the Triassic of central Spain: Sedimentology, v. 49, p. 303-317.

Arakel, A. V., G. Jacobson, M. Salehi, and C. M. Hill, 1989, Silicification of calcrete in paleodrainage basins of the Australian arid zone: Australian Journal of Earth Sciences, v. 36, p. 73-89.

Arbey, N., 1980, Les formes de la silice et l’identification des évaporites dans les formations silicifiés: Bulletin, Centre Recherche Exploration–Production Elf- Aquitaine, v. 4, p. 309-365.

Birnbaum, S. J., and J. W. Wireman, 1985, Sulfate-reducing bacteria and silica solubility; a possible mechanism for evaporite diagenesis and silica precipitation in banded iron formations: Canadian Journal of Earth Sciences, v. 22, p. 1904-1909.

Bustillo, M. A., 2010, Chapter 3 Silicification of Continental Carbonates, in A. M. Alonso-Zarza, and L. H. Tanner, eds., Developments in Sedimentology, v. Volume 62, Elsevier, p. 153-178.

Bustillo, M. A., and A. Alonso-Zarza, 2007, Overlapping of pedogenesis and meteoric diagenesis in distal alluvial and shallow lacustrine deposits in the Madrid Basin, Spain: Sedimentary Geology, v. 198, p. 255-271.

Chafetz, H. S., and J. L. Zhang, 1998, Authigenic euhedral megaquartz crystals in a Quaternary dolomite: Journal of Sedimentary Research Section A-Sedimentary Petrology & Processes, v. 68, p. 994-1000.

Chowns, T. M., and J. E. Elkins, 1974, The origin of quartz geodes and cauliflower cherts through the silicification of anhydrite nodules: Journal Sedimentary Petrology, v. 44, p. 885-903.

Clayton, C. J., 1986, The chemical environment of flint formation in Upper Cretaceous chalks, in G. d. G. Sieveking, and M. B. Hart, eds., The Scientific Study of Flint and Chert: Cambridge, Cambridge University Press, p. 43-54.

Degens, E. T., and S. Epstein, 1962, Relationship between O18/O16 ratios in coexisting carbonates, cherts, and diatomites: Bulletin American Association Petroleum Geologists, v. 46, p. 534-542.

El Khoriby, M., 2005, Origin of the gypsum-rich silica nodules, Moghra Formation, Northwest Qattara depression, Western Desert, Egypt: Sedimentary Geology, v. 177, p. 41-55.

El Tabakh, M., K. Grey, F. Pirajno, and B. C. Schreiber, 1999, Pseudomorphs after evaporitic minerals interbedded with 2.2 Ga stromatolites of the Yerrida basin, Western Australia: Origin and significance: Geology, v. 27, p. 871-874.

Eugster, H. P., 1967, Hydrous sodium silicate from Lake Magadi, Kenya: precursors of bedded chert: Science, v. 157, p. 1177-1180.

Folk, R. L., and J. S. Pittman, 1971, Length-slow chalcedony; a new testament for vanished evaporites: Journal Sedimentary Petrology, v. 41, p. 1045-1058.

Gao, G., and L. S. Land, 1991, Nodular chert from the Arbuckle Group, Slick Hills, SW Oklahoma: a combined field, petrographic and isotopic study: Sedimentology, v. 38, p. 857-870.

Geeslin, J. H., and H. S. Chafetz, 1982, Ordovician Aleman ribbon cherts; an example of silicification prior to carbonate lithification: Journal of Sedimentary Petrology, v. 52, p. 1283-1293.

Goldberg, K., S. Morad, I. S. Al-Aasm, and L. F. De Ros, 2011, Diagenesis of Paleozoic playa-lake and ephemeral-stream deposits from the Pimenta Bueno Formation, Siluro-Devonian (?) of the Parecis Basin, central Brazil: Journal of South American Earth Sciences, v. 32, p. 58-74.

Gómez-Alday, J. J., F. Garcia-Garmilla, and J. Elorza, 2002, Origin of quartz geodes from Lano and Tubilla del Agua sections (middle-upper Campanian, Basque-Cantabrian Basin, northern Spain): isotopic differences during diagenetic processes: Geological Journal, v. 37, p. 117-134.

Guidry, S. A., and H. S. Chafetz, 2002, Factors governing subaqueous siliceous sinter precipitation in hot springs: examples from Yellowstone National Park, USA: Sedimentology, v. 49, p. 1253-1267.

Guthrie, G. D., D. Bish, and R. C. Reynolds, 1995, Modeling the X-ray diffraction pattern of opal CT: American Mineralogist, v. 80, p. 869-872.

Hay, R. L., 1968, Chert and its sodium-silicate precursors in sodium-carbonate lakes of east Africa: Contributions to Mineralogy and Petrology, v. 17, p. 255-274.

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

Heaney, P. J., 1995, Moganite as an indicator for vanished evaporites: a testament reborn?: Journal of Sedimentary Research A: Sedimentary Petrology & Processes, v. A65, p. 633-638.

Henchiri, M., and N. Slim-S'Himi, 2006, Silicification of sulphate evaporites and their carbonate replacements in Eocene marine sediments, Tunisia: two diagenetic trends: Sedimentology, v. 53, p. 1135-1159.

Hesse, R., 1989, Silica diagenesis: origin of inorganic and replacement cherts.: Earth Science Reviews, v. 26, p. 253-284.

Jones, B., and R. W. Renaut, 2007, Microstructural changes accompanying the opal-A to opal-CT transformation: new evidence from the siliceous sinters of Geysir, Haukadalur, Iceland: Sedimentology v. 54, p. 921-949.

Jones, B., R. W. Renaut, and M. R. Rosen, 1996, High-temperature (W901C) calcite precipitation at Waikite Hot Springs, North Island, New Zealand: Journal of the Geological Society of London, v. 153, p. 481-496.

Jones, B. F., S. L. Rettig, and H. P. Eugster, 1967, Silica in alkaline brines: Science, v. 158, p. 1310-1314.

Jones, J. B., and E. R. Segnit, 1971, The nature of opal. Part 1: Nomenclature and constituent phases: Journal of the Geological Society of Australia v. 18, p. 57-68.

Keene, J. B., 1983, Chalcedonic quartz and occurrence of quartzine (length-slow chalcedony) in pelagic sediments: Sedimentology, v. 30, p. 449-454.

Knauth, L. P., 1979, A model for the origin of chert in limestone: Geology, v. 7, p. 274-277.

Knauth, L. P., 1994, Petrogenesis of chert: Reviews in Mineralogy and Geochemistry, v. 29, p. 233-258.

Knauth, L. P., and S. Epstein, 1976, Hydrogen and oxygen isotope ratios in nodular and bedded cherts: Geochimica et Cosmochimica Acta, v. 40, p. 1095-1108.

Knauth, L. P., and D. R. Lowe, 2003, High Archean climatic temperature inferred from oxygen isotope geochemistry of cherts in the 3.5 Ga Swaziland Supergroup, South Africa: Geological Society of America Bulletin, v. 115, p. 566-580.

Krainer, K., and C. Spotl, 1998, Abiogenic silica layers within a fluviolacustrine succession, Balzano volcanic complex, Northern Italy - A Permian analogue for Magadi-type cherts: Sedimentology, v. 45, p. 489-505.

Mackenzie, F. T., and R. Gees, 1971, Quartz: synthesis at earth-surface conditions: Science, v. 173, p. 533-535.

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

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

 

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

 

Miehe, G., and H. Graetsch, 1992, Crystal structure of moganite: a new structure type for silica: European Journal of Mineralogy, v. 4, p. 693-706.

Milliken, K. L., 1979, The silicified evaporite syndrome; two aspects of silicification history of former evaporite nodules from southern Kentucky and northern Tennessee: Journal Sedimentary Petrology, v. 49, p. 245-256.

Muchez, P., P. Vanderhaeghen, H. El Desouky, J. Schneider, A. Boyce, S. Dewaele, and J. Cailteux, 2008, Anhydrite pseudomorphs and the origin of stratiform Cu–Co ores in the Katangan Copperbelt (Democratic Republic of Congo): Mineralium Deposita, v. 43, p. 575-589.

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

Peterson, M. N. A., and C. C. Von der Borch, 1965, Chert: modern inorganic deposition in a carbonate precipitating localitty: Science, v. 149, p. 1501-1503.

Pirajno, F., and K. Grey, 2002, Chert in the Palaeoproterozoic Bartle Member, Killara Formation, Yerrida Basin, Western Australia: a rift-related playa lake and thermal spring environment?: Precambrian Research, v. 113, p. 169-192.

Rodgers, K. A., and G. Cressey, 2001, The occurrence, detection and significance of moganite (SiO2) among some silica sinters: Mineralogical Magazine, v. 65, p. 157-167.

Scholle, P. A., and D. S. Ulmer-Scholle, 2003, A Color Guide to the Petrography of Carbonate Rocks: Grains, textures, porosity, diagenesis, v. 77: Tulsa, Okla, American Association of Petroleum Geologists Memoir, 459 p.

Siever, R., 1962, Silica solubility, 0°–200 °C., and the diagenesis of siliceous sediments: Journal of Geology, v. 70, p. 127-150.

Tucker, M. E., 1976a, Quartz replaced anhydrite nodules ('Bristol Diamonds') from the Triassic of the Bristol District: Geological Magazine, v. 113, p. 569-574.

Tucker, M. E., 1976b, Replaced evaporites from the late Precambrian of Finnmark, Arctic Norway: Sedimentary Geology, v. 16, p. 193-204.

Ulmer-Scholle, D. S., and P. A. Scholle, 1994, Replacement of evaporites within the Permian Park City Formation, Bighorn Basin, Wyoming, USA: Sedimentology, v. 41, p. 1203-1222.

Ulmer-Scholle, D. S., P. A. Scholle, and P. V. Brady, 1993, Silicification of evaporites in Permian (Guadalupian) back-reef carbonates of the Delaware Basin, west Texas and New Mexico: Journal of Sedimentary Petrology, v. 63, p. 955-965.

 

van den Boorn, S. H. J. M., M. J. van Bergen, W. Nijman, and P. Z. Vroon, 2007, Dual role of seawater and hydrothermal fluids in Early Archean chert formation: Evidence from silicon isotopes: Geology, v. 35, p. 939-942.

 

Warren, J. K., 2016, Evaporites: A compendium (ISBN 978-3-319-13511-3) Published Feb. 22, 2016: Berlin, Springer, 1854 p.

West, I., 1973, Vanished evaporites; significance of strontium minerals: Journal of Sedimentary Petrology, v. 43, p. 278-279.

West, I. M., 1964, Evaporite diagenesis in the lower Purbeck beds of Dorset [with discussion]: Proc. Yorkshire Geol. Soc., v. 34, p. 315-330.

Wheeler, W. H., and D. A. Textoris, 1978, Triassic limestone and chert of playa origin in North Carolina: Journal of Sedimentary Petrology, v. 48, p. 765-776.

Williams, L. A., and D. A. Crerar, 1985, Silica diagenesis, II. General mechanisms: Journal of Sedimentary Petrology, v. 55, p. 312-321.

Williams, L. A., G. Parks, and D. A. Crerar, 1985, Silica diagenesis, I. Solubility controls: Journal of Sedimentary Petrology v. 55, p. 301-311.

 


 

 

 

 

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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