Geological Digressions

Mineralogy of carbonates: sabkhas

Whiskey is for drinkin’; water is for fightin!

The Architecture of Connected Holes; A Different Way to Look at the Liquid Earth

A misspent youth serves to illustrate groundwater flow

Contrails, analogies, and visualizing groundwater flow

Springs and seeps

References

M.S. Bedinger and J.R. Harrill. 2012. Groundwater Geology and Hydrology of Death Valley National Park, California and Nevada. National Park Service, U.S. Dept. of Interior, Natural Resource Technical Report 2012/652. A comprehensive account of DV. Includes a discussion of the influence of glacial-interglacial changes in climate during the Pleistocene, the succession of lakes during pluvial dominant climates, and the alternation of evaporite – non-evaporite conditions.

W.R. Belcher, M.S. Bedinger, J.T. Back and D.S.Sweetkind, 2009. Interbasin flow in the Great Basin with special reference to the southern Funeral Mountains and the source of Furnace Creek Springs, Death Valley, California, U.S. Journal of Hydrology, v. 369, p. 30-43

W.R. Belcher and D.S.Sweetkind, 2010, Death Valley regional groundwater flow system, Nevada and California: hydrogeologic framework and transient groundwater flow model. U.S.G.S. Professional Paper 1711.

C. Del Pardo, B.R. Smith-Konter, L.F.Serpa, C. Kreemer, G, Blewitt, W.C. Hammond. 2012. Interseismic deformation and geologic evolution of the Death Valley Fault Zone. Journal of Geophysical Research, v. 117

Li, J., Lowenstein, T.K. and Blackburn, I.R. 1997. Responses of evaporite mineralogy to inflow water sources and climate during the past 100 k.y. in Death Valley, California. Geological Society of America Bulletin, 109, p. 1361-1371. The paper that identifies three chemical groups of inflow waters.

T.K. Lowenstein and F. Risacher, 2009. Closed basin brine evolution and the influence of Ca-Cl inflow waters: Death Valley and Bristol Dry Lake, California, Qaidam Basin, China, and Salar de Atacama, Chile. Aquatic Geochemistry, v. 15, p. 71-94. Focuses on the Ca-Cl waters originating from a magma chamber beneath DV. Also describes briefly the changes wrought by recurring pluvial periods during the Pleistocene.

Mineralogy of evaporites: saline lake brines

Leave a Reply

Gypsum books, several cm across, precipitated from supersaturated brines within fine-grained sediment and exumed by wind ablation. Altiplano, Chile

Gypsum books, several cm across, precipitated from supersaturated brines within fine-grained sediment and exhumed by wind ablation. Altiplano, Chile

The evolution of brines that produce saline lake evaporites.

This is part of the How To…series on evaporites

Seawater is generally considered to be isochemical because its composition varies little from sea to sea. There are some changes in saturation levels of calcite and aragonite in deep ocean waters (below their compensation depths), and salinity in basins having more restricted seawater inflow (e.g. Mediterranean Sea, Arabian Gulf), but the range of composition variation is limited. This means that evaporation of seawater follows a predictable chemical and thermodynamic path.

Not so terrestrial waters (rivers, groundwater) whose starting compositions are remarkably varied. Most natural waters contain Na in abundance plus an array of cations and anions like K⁺, Ca²⁺, Mg²⁺, Cl^–, SO₄^2-, HCO₃^–, CO₃^2-, and SiO₂. But the starting concentrations of these ions varies greatly meaning that the resulting evaporite mineralogy is also varied. Enter Lawrence Hardie and Hans Eugster whose chemical modelling in the late 1960s and 70s showed us how to make sense of this natural variability and complexity. Their publications are still cited 50 years later (a few are listed below); they are compulsory reading for any student who is interested in evaporites and geochemistry.

Hardie and Eugster (1970) calculated evaporation paths and mineral products for natural waters having different compositions at atmospheric pressure and 25^oC. The waters are assumed to be in equilibrium with CO₂ (pCO₂ is constant) – this is important for CO₃^2- and HCO₃^– and calculation of pH. Some of the relevant calculations have been described in a previous post on carbonate geochemistry – the same principles apply to evaporites.

Equilibrium constants for each mineral (common evaporite minerals from saline lakes are listed below),
Ion activities and activity coefficients,
Ion activity products (IAP),
Degree of mineral saturation,
Ionic strength, low values in dilute water, increasing with evaporation,
Note that complex ion pairs are not included in the calculations.

As water evaporates, the activities of cations and anions increases. If the ion activity product equals the equilibrium constant for a mineral, then the solution has become saturated and that mineral will precipitate; at this point the IAP remains constant. The important results in this process are, using calcite as an example:

During precipitation, both Ca²⁺ and CO₃^2- are removed from the original water in equal proportions,
Calcite will continue to precipitate until one or both Ca²⁺ and CO₃^2- have been depleted.
It is likely that the original concentrations of Ca²⁺ and CO₃^2- are not equal. If, in this example Ca²⁺ > CO₃^2-, then calcite will precipitate until the CO₃^2- is used up, leaving excess Ca²⁺ available for some other mineral like gypsum.
If on the other hand Ca²⁺ < CO₃^2-, then the Ca²⁺ will be depleted first leaving none for gypsum.
Water is continually removed during evaporation and its concentration decreases; if the process proceeds to desiccation then the concentration of water is zero.

As evaporation proceeds the model predicts which minerals are likely to precipitate and which will not, depending on the starting composition of the water. The succession of mineral precipitates defines the evaporation path for that particular water.

Two reasonably typical examples are shown below (from Hardie and Eugster, 1970, Figure 2). Here, the apices of each triangle plot the changes in ion concentrations as evaporation proceeds. For the example of calcite precipitation, the path moves away from the Ca and CO₃^2- – HCO₃^– apex towards the Cl apex.

Note that the evaporation paths turn abruptly at the point where precipitation depletes the relevant ions.

Example (a): Sulphate rich water that evolves to a Cl brine via precipitation of gypsum.

Example (b): Carbonate rich – sulphate poor water; no gypsum precipitates and the system evolves to a CO3–Cl brine. If calcite precipitates, the brine will likely become rich in Na and K. Halite saturation will be reached if evaporation continues.

In their model, brine evolution curves are complicated by SiO₂ and precipitation of sepiolite (illustrated in the flow-path diagrams below). Sepiolite is a hydrated Mg-silicate; its precipitation will influence the Mg path. Precipitation of sepiolite also releases H⁺ and is therefore an important part of pH buffering (along with the carbonate equilibria). As pH is lowered, CO₃^2- will decrease. Therefore, there is potential for reversal of the CO₃^2- enrichment trends.

Hardie and Eugster (1970) identified 4 important groups of natural waters; Eugster and Hardie (1978) added a fifth group. Note (1) not all waters will fit neatly into one of these groups, and (2) the ions that define each group are dominant –subordinate cation and anion species will also be present:

Na–CO₃-Cl
Na–CO₃-SO₄–Cl (e.g. Lake Magadi, Kenya)
Na–(Ca)–SO₄-Cl (e.g. Great Salt Lake, Utah)
Mg–Na–(Ca)–SO₄-Cl (Poison Lake, Washington)
Ca–Mg–Na–(K)–Cl (Bristol Dry Lake, Mohave Desert, California; Dead Sea)

Table of common evaporite minerals in saline lakes; click on the image to enlarge

In all waters, alkali metal carbonates are the first to precipitate, particularly calcite or aragonite depending on the Mg/Ca ratio. This is a critical first stage in evaporation because it determines the subsequent precipitation sequence of minerals as Ca²⁺ and CO₃^2- are removed from solution. This is the Calcite Divide, separating waters that become HCO₃ rich or HCO₃ poor (keep in mind that continued evaporation increases the overall ionic strength). Depending on the SO₄ concentration, excess Ca results in gypsum precipitation that, in turn, creates the next geochemical divide – the gypsum divide that determines whether the brines evolve as SO₄ rich – Ca poor, or SO₄ poor – Ca rich. One of the last minerals to precipitate is halite from Cl-rich brines.

Brine evolution is shown diagrammatically below (modified from Figures 7 & 8 in Hardie and Eugster, 1970; Warren, 2016, Fig. 2). The two pathways represent waters that include sepiolite precipitation, and those lacking significant aqueous SiO₂. For the case where sepiolite precipitates, there is an additional path towards gypsum saturation because of the pH effect on CO₃-HCO₃ (noted above), such that the brine becomes enriched in Ca.

You should also check the flow diagram in Warren (2016) Figure 2, that shows in greater detail complications such as changes in Mg/Ca ratios, variations in HCO₃, in addition to sepiolite production.

In this post, I have focused on the work of Hardie and Eugster because it has been pivotal in guiding more recent research. Of course, there have been modifications and improvements to their models, and you can access these recent advances in the literature cited below.

Links to related topics

Mineralogy of evaporites: The rise of diapirs

Mineralogy of evaporites: salt tectonics

Mineralogy of evaporites: Saline lakes

Mineralogy of evaporites: Death Valley hydrology

Mineralogy of evaporites: Marine basins

Mineralogy of carbonates: sabkhas

References

M. Babel and B.C. Schreiber, 2014. Geochemistry of evaporites and evolution of seawater. Treatise on Geochemistry. Elsevier, p. 483-560. Deals with marine evaporites but many aspects relevant to non-marine. Encyclopedic with extensive list of references

D.M. Deocampo and B.F. Jones. 2014. Geochemistry of Saline Lakes. Treatise on Geochemistry. Elsevier, p. 437-469. Encyclopedic with extensive list of references

H.P.Eugster, 1980. Geochemistry of evaporitic lacustrine deposits. Annual Review of Earth & Planetary Sciences, v. 8, p. 35-63.

L.A. Hardie and H.P. Eugster, 1970. The evolution of closed basin brines. Mineralogical Society of America, Special Paper v. 3, p. 273-290. 50 years old but still an iconic paper and a must-read.

J. Warren, 2016. Evaporites. In W.M. White (Ed.) Encyclopedia of Geochemistry. Springer International, p. 1-8. Concise summary of saline lake and marine brine.

Mineralogy of evaporites: Saline lakes

Leave a Reply

Saline lake evaporites – some general physical and chemical conditions

This is part of the How To…series on evaporites

Evaporites are rocks formed by precipitation of water-soluble salts under conditions of intense evaporation. The overarching condition is that evaporation of surface waters exceeds inflow or recharge, such that mineral saturation or supersaturation is maintained. From a geological perspective this usually requires basins to be closed. This is fairly easy to visualize in small land-locked basins like salt lakes, playa lakes, and salars that are bound by topographic barriers or where surface drainage is minimal or non-existent – in such cases new water may flow into the basin but none leaves, except by evaporation (a kind of Hotel California of evaporites).

Evaporites formed in oceanic systems are quite different. Iconic examples like the Permian Zechstein Sea indicate such basins were huge. These basins must still be hydrologically ‘closed’ such that surface evaporation exceeds seawater inflow, which means that mixing of normal seawater and saturated water must be kept to a minimum. Thus, we might expect such conditions on shallow platforms that are protected by a barrier of some kind (like the Late Miocene-Pliocene Messinian Mediterranean Sea), or in deeper basins where salinity stratification is maintained. Notably, there are no modern analogues for these types of evaporite deposit.

The notes here are a brief (and incomplete) introduction to saline lake evaporites. The literature I have drawn on (below) is more encyclopedic.

Tectonics

Tectonics ultimately controls the conditions in which evaporites form: topography, climate ( humidity and precipitation), groundwater flow and surface drainage, bedrock composition and its weathering products. Tectonic-geomorphic controls on closed-basin evaporites are a lesson in contrasts:

Australian basins, like the ephemeral Kati Thanda – Lake Eyre, or the acidic salt flats in Yilgarn, have formed on ancient, stable, deeply weathered Precambrian cratons;
The Death Valley salt lakes occur in a region of active crustal extension – the Basin and Range Province. In northeast Africa, Magadi Lake is one of a string of salt lakes in the East African Rift.
Dead Sea fills a pull-apart basin along the Dead Sea Transform Fault System – the lithospheric-scale strike-slip boundary between the African and Arabian plates;
Salars of the Chilean Altiplano-Atacama region are nestled against volcanic complexes, including calderas, associated with Andean subduction.

Climate

For evaporites to accumulate, evaporation must balance or exceed water inflow from precipitation, surface runoff, or groundwater seepage. This does not mean that it can never rain in such places, but that floods or deluges must be short-lived. Even the Atacama Desert receives a few drops of rain or rare flash floods.

All evaporites occur in arid climates. We commonly associate aridity with warmth – Kati Thanda – Lake Eyre, Dead Sea and Death Valley are but three (of many) regions associated with subtropical deserts. Evaporites can also accumulate in cold, even frigid climates; good examples can be found in the Dry Valley lakes of Antarctica (common gypsum and halite) where mean annual temperatures range from -15^oC to -30^oC. Likewise, many salars of the Chilean Atliplano are at elevations of about 4000 m, where summer temperatures are cool and humidity is close to zero. While temperature may not determine where an evaporite deposit will form, higher temperatures do increase the rates of evaporation and mineral precipitation. Temperature will also control the local micro- and macro-biota, both of which can influence mineral precipitation and the REDOX conditions of lake waters.

Hydrology and hydrogeology

Basins that lack outgoing surface drainage are called endorheic basins. Such basins are surrounded by drainage divides and the area of drainage can be significantly greater than that of the enclosed lake. Water enters these basins as surface runoff (rain or snow melt), groundwater seepage, direct precipitation and in some cases hydrothermal flow. Except for precipitation, all these fluid inputs must contain sufficient solute. It is the balance among all these inputs against evaporation that maintains levels of mineral saturation-supersaturation or desiccation.

The elevated topography ensures that groundwater flow is towards the lake. Groundwater will exit at the shoreline and if there is sufficient hydraulic drive (depending on the amount of topographic relief) seepage may occur across the lake floor. Surface runoff, occasionally manifested as flash floods, will also deposit clastic sediment along the lake margins. For example, alluvial fans may interfinger with evaporite deposits.

Evaporite composition

Oceanic evaporites are formed from seawater, the composition of which has probably not varied much over geological time. In contrast, and as Hardie and Eugster (1970) state in one of their iconic papers on saline lakes, the “composition of brines … exhibits a bewildering range”. There is an equally impressive array of evaporite minerals – several of the publications listed below tabulate the mineral species. The original composition of saline lake waters is governed by several factors:

Most natural waters contain plenty of Na, plus a relatively limited array of other cations and anions like K⁺, Ca²⁺, Mg²⁺, Cl^–, SO₄^2-, HCO₃^–, CO₃^2-, and some SiO₂.
Bedrock-soil composition through which waters seep. Surface weathering and soil development, particularly soils influenced by microbes, will be strong determinants of lake water compositions. For comparison, limestones will produce calcium and carbonate rich waters with slightly elevated pH values (pH 8-9), whereas rocks containing sulphides will produce more acidic waters where HCO₃^– and CO₃^2- are absent (unstable). Those having a surfeit of alkali metal carbonates (most commonly Na, K and Li) can have pH values >10. Lithium is commercially important in brines beneath several salars of Chile, Bolivia and Argentina.
Residence time of groundwater: this is basically the time taken from infiltration at the surface, to surface seepage at some location down hydraulic gradient. Longer residence times usually give rise to higher solute concentrations.
Hydrothermal fluid input can introduce ion species less abundant than in more local bedrock, as well as influencing lake water temperature.
Large saline lakes may be become salinity and density stratified. Upper layers will tend to be more oxygen-rich because of regular mixing; deeper layers more anoxic. Salinity stratification is an important control on REDOX conditions throughout lake waters.
REDOX conditions: Oxidation can play a significant role in fluid composition where sulphides are present in bedrock (commonly as mineralization). Sulphides readily oxidize to sulphate, releasing cations like iron (Fe²⁺ becomes Fe³⁺) and acidic waters (this is one of the main problems of acid-mine drainage). In larger lakes, the upper layers can support phototrophic and other organisms (e.g. algae, diatoms), that ultimately sink through the brine layer to the lake floor where decay consumes any remaining oxygen. Here, the salinity boundary acts as a chemocline, a chemical boundary, below which the redox conditions change drastically. Under these conditions, metal oxides like Fe₂O₃ may dissolve and any sulphate is reduced to sulphide; this may happen at or below the sediment-water interface where metal cations like Fe, Cu, Mn, Pb and Zn are stabilized as their respective sulphides. The deeper anoxic brines may also support methanogenic or anaerobic bacteria.
Cycle of wetting and drying: Maintenance of evaporite conditions requires intense evaporation over relatively long periods. However, successive wetting and drying cycles can influence brine and mineral compositions. Under shallow conditions, wetting (e.g. rain) can dissolve some minerals and temporarily change the saturation state of lake and interstitial fluids; for example, halite dissolves readily, potentially removing Na and Cl from the site. Drying eventually results in the kind of evaporative pumping witnessed on sabkhas, moving water and solute from the watertable through the capillary zone. In larger lakes, extended dry periods may shift brine stratification boundaries in concert with lower lake levels and migrating shorelines. Brief periods of precipitation will increase mixing and oxygenation of upper lake layers, perhaps even diluting the deeper brines. Any sediment introduced to the lakes, particularly bottom-hugging sediment gravity flows, may also modify brine stratification and disturb chemoclines.

The next post deals with closed-basin brine composition and evolution

The photo of Death Valley salt polygons is from Marli Miller’s beautiful image collection.

Links to related posts

Mineralogy of evaporites: salt tectonics

Mineralogy of evaporites: Saline lake brines

Mineralogy of evaporites: Death Valley hydrology

Mineralogy of evaporites: Marine basins

Mineralogy of carbonates: sabkhas

References

K.C. Bennison and B.B. Brown. 2015. The evolution of end-member continental waters: The origin of acidity in southern Western Australia. G.S.A. Today, v. 25, No. 6, p. 4-10.

D.M. Deocampo and B.F. Jones. 2014. Geochemistry of Saline Lakes. Treatise on Geochemistry. Elsevier, p. 437-469.

L.A.Hardie. 1991. The significance of evaporites. Annual Review Earth & Planetary Sciences, v. 19, p. 131-168. Click on the OA full text icon top right.

J. Warren, 2016. Evaporites. In W.M. White (Ed.) Encyclopedia of Geochemistry. Springer International, p. 1-8. Concise summary of saline lake and marine brine.

Mineralogy of carbonates; Sabkhas

Leave a Reply

Modern marine sabkhas – transitions between carbonates and evaporites

This is part of the How To…series on carbonate rocks

The development of sedimentary facies models based on modern depositional analogues flourished in the 1960s and 70s on the heels of foundational work by 1950s icons (I’m loath to name them because I’m bound to miss someone out). Modern sabkhas were an important part of this trend as a worthy and convenient analogue for many ancient marine evaporite deposits, convenient in the sense that the analogue model was sometimes applied to evaporites even when evidence for “shallow marine conditions” were lacking. It is now firmly established that not all evaporites require a seawater precursor.

Sabkhas and the depositional models derived from them are important because:

they are part of a sedimentologist’s toolbox for interpreting ancient environments (sedimentary facies, stratigraphy, and geochemistry),
sabkha diagenesis includes rapid precipitation (and dissolution) of halite at the surface, gypsum-anhydrite in the shallow subsurface, and carbonate hardgrounds,
they are one of the few sedimentary environments where recent dolomite precipitates,
they help define the excursions of paleo-shorelines (regression, transgression),
they place narrow boundary conditions on paleoclimates,
it is possible to track ancient seawater compositions, and
they provide a kind of go-between for marine carbonate and evaporite facies.

The classic areas of study (one might refer to it informally as the ‘type area’) are the shallow coasts of Abu Dhabi (United Arab Emirates, until 1971 referred to as Trucial states) and Qatar. They are the coastal and inland extensions of the shallow, ramp-like Arabian Gulf. Despite its geological and ecological importance, only about a third of the original Abu Dhabi sabkha coast has escaped urban and industrial encroachment (Lokier, 2013).

Sabkhas are important components of coastal plains, comprising supratidal flats and salt marshes that periodically are flooded by spring tides. They are characterised by a mix of evaporite minerals (predominantly gypsum, anhydrite and halite) and carbonates, and hence are generally restricted to hot arid climates (Abu Dhabi and Qatar sabkhas straddle the Tropic of Cancer). In places, the sabkhas extend inland for 30km; they are remarkably flat having slopes less than 0.1^o (Lokier et al. 2018), which means that even minor excursions of relative sea level will have a profound effect on shorelines and associated facies.

Modern sabkhas occur at several other locations (e.g. Sinai Peninsula, and Coorong Lagoon south of Adelaide) but we will focus on Abu Dhabi. Here the coast is bordered by numerous small islands, separated by tidal channels and enclosing lagoons floored by carbonate mud, ooid and skeletal carbonate shoals, and beaches. Small reefs have grown on the seaward margins of some islands. Inland, beach ridges as old as 4000 years provide good evidence for ancient shorelines, successively abandoned during coastal progradation.

A typical transect across the Abu Dhabi coastal plain encounters sedimentary facies that are also represented in shallow stratigraphic sections (revealed in cores). I have drawn on the excellent descriptions in classic papers by Evans et al (1964), Illing et al (1965), Kendall and Skipwith (1969), Bathurst (1976, Chapter 4), and more recent contributions by Paul and Lokier (2017), and Lokier et al. (2017).

Lower intertidal and subtidal areas are floored primarily by skeletal and pelloidal sands, drained by tidal channels. Local sheltered spots contain aragonite muds. There is a bustling benthic community of molluscs, particularly grazing Cerithiid gastropods and burrowing crabs, echinoderms, foraminifera and calcareous algae.

An important early diagenetic component of this facies belt is a hardground, composed mainly of skeletal debris and cemented by high-Mg calcite, aragonite and some dolomite. Exposed hardgrounds are frequently broken into polygons that may have been caused by desiccation. They occur at the surface in the modern intertidal zone but can be traced landward beneath the sabkha surface. Dating of cemented shells by C¹⁴ indicates ages up to 7000 years BP beneath the sabkha. Hardground formation seems to have kept pace with the migrating shoreline during late Holocene progradation.

The upper intertidal realm is characterised by a belt of algal mats up to 600 m wide (also referred to as microbial or cyanobacteria mats, and stromatolites). Seaward and landward boundaries of this facies belt are well defined: the former by grazing Cerithiids and crabs that cannot thrive in the elevated salinities higher up-slope; the latter by intense desiccation that characterises the sabkhas. The leathery mats are commonly arranged as decimetre-sized polygons with up-turned and overturned margins, long considered to be the product of desiccation. This hypothesis has been challenged by Lokier et al (2017) who observed breakage during competitive mat growth, such that polygon margins are progressively displaced, overthrust or overturned. Overgrowth by a new generation of microbes can be recognized by the small discordances between mat layers. The polygons sometimes contain gypsum crystal mush. Incipient aragonite and gypsum cements are found beneath some mats.

A range of mat morphologies are present, depending on the regularity of tidal flooding and desiccation. Flatter, or pustular forms are more common in lower intertidal zones. Crinkly and tufted forms tend to occupy upper intertidal regions where desiccation is more intense. Mats that are completely dried are reworked during spring tides.

The sabkha facies extends inland from the upper intertidal zone. It tends to be flat, featureless, and devoid of vegetation. Halite crusts are common; large gypsum crystals commonly poke through the surface sediments. The halite is ephemeral – it dissolves during spring tides and any brief periods of rain.

There is an important relationship between the mineralogy of sabkhas, evaporite crystal growth, and depth to the permanent watertable; note that in general, the watertable will rise towards the surface at it approaches the shoreline. These phreatic brines provide the solute for precipitation of gypsum, anhydrite, and halite. Transfer of solute to the overlying (unsaturated) vadose zone and sabkha surface is by evaporative pumping. Thus, most of the gypsum and anhydrite is precipitated in the capillary zone above the watertable, the crystals displacing sediment as they grow (the capillary zone is not permanently saturated with groundwater but is wetted at grain contacts).

Sabkha structures that have reasonable preservation potential are Tepees – conical or inverted ‘V’ structures in broken, cemented crusts and hardgrounds. The crusts are cemented by aragonite and high Mg calcite, but also form in surficial halite. Tepee structures and their associated polygons form by expansion of rapidly cementing surface sediment; the expansion causes breakage that leads to upturned, overturned, and overthrust crust fragments. Tepees formed of carbonate crusts are commonly peritidal, at the junction between the watertable and sediment surface. Those formed in halite are supratidal (like the examples shown here). Tepee structures have been recognized in many ancient carbonates.

The watertable beneath Abu Dhabi sabkhas is usually less than 1 metre below the surface. Groundwater here is highly saline – some measured salinities are 10-20 times that of seawater. The recharge to groundwater needs to balance that lost by intense evaporation. Early workers thought that periodic surface flooding by seawater, or perhaps landward subsurface seepage of seawater would suffice. However, hydrogeological analysis by Sanford and Wood (2001) showed that these processes were relatively minor. Two other processes dominate recharge: upward seepage of groundwater from deeper permeable units, and downward infiltration from rainfall – the latter being most significant. Rainfall also has the potential to change evaporite chemistry – it will dissolve halite (halite has low preservation potential) and may create (temporary) unsaturated conditions for gypsum and anhydrite, resulting in a degree of dissolution.

Stratigraphy

A typical stratigraphic column for Abu Dhabi sabkhas is shown below. The profile (dug to the watertable) records sabkha progradation over intertidal algal mats and cerithiid lags, and subtidal (lagoonal) mixed carbonate plus siliciclastic silts and muds. Displacive gypsum and anhydrite crystals up to 10 cm long occur in several layers. A hardground at the base of this section is the buried equivalent of the hardground presently forming in the intertidal-subtidal facies belts.

The muds, silts and algal mats account for about 60-70% of the local stratigraphy. During compaction and lithification, the thickness of these lithologies will be significantly reduced to the extent that the other layers may appear over-represented. For example, the mat layer, about 150 mm thick would, upon compaction, be represented by a layer no more than a few millimetres thick. In comparison, the thickness of the cerithiid layer may change very little. The halite crust may not be preserved at all!

The sabkha photos generously contributed by Stephen Lokier, are indicate on each figure.

References

G. Evans, C.G.St.C. Kendall and P. Skipwith. 1964. Origin of the coastal flats, the Sabkha of the Trucial Coast, Persian Gulf. Nature, v. 202, p. 759-761.

L.V. Illing, A.J. Wells and J.C.M. Taylor. 1965. Penecontemporary dolomite in the Persian Gulf. In, Dolomitization and limestone diagenesis, a symposium. SEPM Special Publication v.13, p. 89-111.

C.G.St.C. Kendall and P.A.D.E. Skipwith, 1969. Holocene shallow water carbonate and evaporite sediments of Kor al Bazam, Abu Dhabi, southwest Persian Gulf. AAPG Bulletin, v. 53, p. 841-869.

S.W.Lokier. 2013. Coastal Sabkha Perservation in the Arabian Gulf. Geoheritage, v. 5, p.11-22.

S.W. Lokier et al. 2017. A new model for the formation of microbial polygons in a coastal sabkha setting. The Depositional Record, v.3, p. 201-208. Open Access.

S.W. Lokier, W.M. Court, T. Onuma and A. Paul. 2018. Implications of sea level rise in a modern carbonate ramp setting. Geomorphology, v. 304, p. 64-73.

A. Paul and S.W. Lokier, 2017. Holocene marine hardground formation in the Arabian Gulf: Shoreline stabilization, sea level, and early diagenesis in the coastal sabkha of Abu Dhabi. Sedimentary Geology, v. 352, p. 1-13.

W.E. Sanford and W.W. Wood. 2001. Hydrology of the coastal sabkhas of Abu Dhabi, United Arab Emirates. H 358-366.ydrogeology Journal, v.9, p.

Links to other posts in this series:

Mineralogy of carbonates; skeletal grains

Mineralogy of carbonates; non-skeletal grains

Mineralogy of carbonates; lime mud

Mineralogy of carbonates; cements

Mineralogy of carbonates; sea floor diagenesis

Mineralogy of carbonates; Beachrock

Mineralogy of carbonates; deep sea diagenesis

Mineralogy of carbonates; Karst

Mineralogy of carbonates; Burial diagenesis

Mineralogy of carbonates; Neomorphism

Mineralogy of carbonates; Pressure solution

Mineralogy of carbonates: Stromatolite reefs

Mineralogy of carbonates; Burial diagenesis

Leave a Reply

Variables of temperature, pH, and organic solvents in sandstone-limestone burial diagenesis. From Surdam et al. 1989

Modified from Surdam et al, 1989

This is part of the How To…series on carbonate rocks

Some general statements about the diagenesis of limestones during burial.

As sediment is buried, the combined effects of temperature, pressure and changing fluid composition act to lithify, overprint, and even obliterate all previous diagenetic histories. In limestones, the original sediment components and any early cements formed during seafloor or shallow meteoric diagenesis, are cemented, replaced by more stable carbonate phases, or recrystallized. Burial diagenesis is governed primarily by:

Increasing temperature as a function of the local geothermal gradients,
Increasing hydrostatic pressures with depth and sediment/water loads,
Reactions resulting from organic maturation where pH buffering and changes in pCO₂drastically shift the stability of carbonates,
Reactions involving silicates (especially clays) that also change fluid composition and pCO₂,
Tectonically induced faulting and folding that can alter permeability pathways, and
Tectonic uplift that reduces ambient temperatures and pressures and exposes indurated rock to meteoric conditions.

Continue reading →

Mineralogy of carbonates; Karst

Leave a Reply

Karst landscapes – limestone dissolution, saturation and kinetics.

This is part of the How To…series on carbonate rocks

Mountains in traditional Chinese painting are commonly depicted as pinnacles appearing out of some ethereal mist, bound by precipitous faces; symbols of some heightened awareness, an expression of deep time. And in the fertile valleys below an alternative, ephemeral human presence, almost an afterthought.

As surreal and metaphorical as these iconic images seem, they are rooted in real-world landscapes, the karst of southern China’s Guizhou, Guangxi, Yunnan and Chongqing provinces that in 2007 were designated a UNESCO World Heritage site. Its cones, pinnacles, sinkholes, caves, bridges and Shillin (Stone Forests) developed in thick Devonian to Triassic limestone. This is one of the classic regions for tropical and subtropical karst.

Karst surface and subsurface structures are sculpted in limestone, dolostone and evaporites like gypsum and anhydrite. In all cases, the primary process is dissolution in the meteoric vadose and shallow phreatic zones. Therefore, karst is best developed in humid climates: prime examples include the tropical-subtropical landscapes of South China, temperate New Zealand (on Oligocene limestones), and the glacial – post-glacial of Ireland (the Burrens, that are underlain by Carboniferous limestone).

Karst formation provides a good opportunity to examine two competing diagenetic processes – dissolution and precipitation in terms of solution saturation, chemical kinetics, and groundwater flow.

Dissolution controlled by calcite saturation

Dissolution of limestone at the Earth’s surface is summarized in the following reaction (keeping in mind that all the carbonate equilibria are involved depending on pH, temperature, and concentrations or activities):

CaCO₃ + H₂O + CO₂(aqueous) → Ca²⁺ + 2HCO₃^– (1)

An important determinant for calcite dissolution is the degree of saturation. Dissolution will take place in solutions that are undersaturated, precipitation in solutions that are over- or supersaturated with respect to calcite. For calcite, the degree of saturation in a solution is calculated by comparing its ion-activity product with the solubility product (i.e. the ion product if that solution were at equilibrium under the same conditions of temperature and pressure). The ratio between these two activity products is called the saturation (Ω). Ω values less than one indicate undersaturation, values greater than one over-saturation; a value of one indicates the solution is in equilibrium with solid calcite.

Rainwater contains dissolved CO₂ and carbonic acid; the average pH is 5.5 to 5.8. As it filters through soils it may pick up additional CO₂ from plant decay. Dissolution of limestone in the vadose and shallow phreatic zones proceeds rapidly because the water is highly undersaturated. As residence time increases in the phreatic zone, so too do the concentrations of dissolved carbonate species; there is a concomitant decrease in undersaturation and as a consequence, a decrease in the rate of limestone dissolution. At some point in this process, the saturation approaches one and dissolution ceases.

Dissolution controlled by kinetics

At this stage in our deliberations we should remind ourselves that the discussion of limestone solubility and groundwater saturation is based on thermodynamic parameters such as ion activity and energy transfer, that together determine whether a chemical reaction will proceed. What thermodynamics doesn’t do is describe the paths which these reactions take. This is the role of Chemical Kinetics. What does this mean?

Kinetics deals primarily with two parameters: the rate at which reactions take place, and the path that chemical species take to form a reaction. Reactions in aqueous solutions involve collisions between at least two ion species. They may combine directly such as:

A + B (reactant ions) → C (product)

or via a smallish number of intermediate steps until the final product is formed (these intermediate steps are called elementary reactions). Each reaction step requires that the ions have a certain amount of energy before it can proceed – this is the activation barrier.

A → A* (fast)

B → B* (slow)

A* + B* → C where A* and B* are short-lived intermediate species.

The overall rate of a reaction is determined by the slowest intermediate reaction – the one that finds it most difficult to reach its activation energy (in this case the reaction involving B).

Knowing something about the kinetics of calcite dissolution and precipitation can help us decipher which processes are important in diagenesis, whether it is cementation on the seafloor or the formation of karst.

We can now look at the picture of limestone dissolution in a different context, summarized in the diagram below. Dissolution is rapid at high degrees of undersaturation because:

There are very few ion species competing for space on active crystal faces, and
Dissolved mass is moved rapidly away. From a chemical kinetic perspective, the reaction is controlled by the rate at which the dissolved mass is removed from calcite crystal surfaces and transferred to some other site (transport-controlled reactions); in groundwater systems this depends on groundwater flow rates. Thus, the reaction is probably a relatively simple A + B → C type.

As the concentration of dissolved species increases (i.e. greater degrees of saturation) the number of molecular collisions also increases at calcite crystal surfaces. Thus, at low levels of undersaturation the rate of dissolution is controlled more by what is happening at the crystal surface – it is a surface controlled reaction and sensitive to factors such as adsorption of ion species on the surface, and dehydration of adsorbed species. For example, Ca²⁺ and CO₃^2- in solution are surrounded by water molecules and for them to combine at the crystal surface, they need to shed this water (dehydrate). Under these conditions, slow intermediate reactions will determine the overall rate of dissolution.

Precipitation

At some point in their seepage journey karst fluids are capable of precipitating calcite, commonly as drip cements in caves as water filters through the vadose zone (providing us with the spectacle of stalactites and stalagmites), and in tufas where groundwaters emerge as springs. For this to happen the solutions must be over- or supersaturated with respect to calcite. However, we also know that as saturation levels approach one there is no further dissolution and therefore no mechanism to increase dissolved carbonate species to levels of oversaturation. Some other process must intervene here to produce supersaturated conditions.

It is generally understood that the partial pressure of CO₂ in water passing through the vadose zone is greater than that in cave atmospheres. As water enters a cave, the various carbonate equilibria will accommodate this change in pCO₂ by degassing CO₂, reducing the concentration of H₂CO₃, and pushing reactions (1) and (2) to the left.

CO₂(gas) ← CO₂ (aqueous) (2)

CaCO₃ + H₂O + CO₂(aqueous) ← Ca²⁺ + 2HCO₃^– (1)

From a kinetic perspective, calcite precipitation under these conditions is surface controlled, aided in part by the availability of nucleation sites on the crystal surface and the delivery or removal of ion species by fluid flow.

Links to other posts in this series:

Mineralogy of carbonates; skeletal grains

Mineralogy of carbonates; non-skeletal grains

Mineralogy of carbonates; lime mud

Mineralogy of carbonates; cements

Mineralogy of carbonates; sea floor diagenesis

Mineralogy of carbonates; Beachrock

Mineralogy of carbonates; deep sea diagenesis

Mineralogy of carbonates; skeletal grains

References and useful texts

D. Ford and P. Williams. 2007. Karst Hydrogeology and Geomorphology. John Wiley & Sons.

J.W. Morse and F.T. Mackenzie 1990. Geochemistry of sedimentary carbonates. Developments in Sedimentology 48. Elsevier, Amsterdam, 707 p.

P.A. Domenico and F.W. Schwartz, 1997. Physical and Chemical Hydrogeology, 2^nd Ed. John Wiley & Sons. 506 p. This book focuses on groundwater but has an excellent section on aqueous chemistry.

USGS Glossary of karst terminology. 1972. PDF version

W.B. White, 2015. Chemistry and karst. Acta Carsologica, v.44/3, p. 349-362. Full text available

Mineralogy of carbonates; diagenetic settings

Leave a Reply

Carbonate diagenesis; How limestones form.

This is part of the of How To…series… on carbonate rocks

Of all the common rock-forming minerals, carbonates are the most reactive chemically. The transformation of loose sediment to hard limestone involves chemical reactions that, depending on the conditions (ionic concentrations, pH, degree of saturation, temperature) promote precipitation or dissolution of minerals, most commonly calcite, Mg-calcite, aragonite and dolomite. These reactions take place at the surface (e.g. sea floor) and at all stages during sediment burial and uplift. Limestone diagenetic pathways are complicated; this is part of the attraction for those who study them (notwithstanding the opportunity to conduct field work in places like Bahamas).

Some general requirements for diagenesis to proceed are:

The thermodynamic stability and metastability of precipitating phases are determined by pressure, temperature and chemical composition of the fluids (including partial CO₂ pressures, pH, and Mg/Ca ratios).
Reactions take place in water: sea water, modified sea water, fresh water and saline brines. These fluids are never static; they flow, delivering new solute (ions in solution) to sites of precipitation, and removing dissolved solids to other sites in the permeable sediment.
Carbonate diagenesis at all burial depths is, like any other rock type, governed by subsurface fluid flow and evolving fluid compositions. Fluid flow itself is governed by hydraulic gradients that are generated by topography, sediment compaction and tectonic loads.
If the conditions change, any phase that becomes thermodynamically unstable will dissolve. This applies particularly to metastable phases like aragonite and high Mg-calcite, where changes in the fluid environment can render them unstable and prone to dissolution. Fluid composition will change when, for example, shallow sea floor carbonates are exposed to meteoric fresh water as sea level falls, or during burial where original seawater is modified by reactions involving carbonates, siliciclastics and, importantly organic matter.

Geologists like to subdivide things and the carbonate diagenetic realm is no different (it helps to simplify a complex world). Three diagenetic environments are frequently cited, with a fourth located at the transition to metamorphism (each environment will be treated in greater detail in separate articles), and depicted in the diagram below:

Seafloor environments (almost syndepositional), including the first few centimetres or metres of burial.
Shallow burial – meteoric environments
Deep burial environments

Sea floor diagenesis

Precipitation of aragonite and high-Mg calcite cements on the sea floor or in the first few centimetres to metres beneath it, is common in tropical settings, less so in cooler waters. It is the region influenced by ambient seawater compositions; it is also referred to as the marine phreatic zone.

This, the earliest stage of carbonate diagenesis is promoted by low sedimentation rates, high levels of mineral supersaturation in seawater and rapid exchange of atmospheric CO₂ with aerated seawater. Microorganisms like bacteria and photosynthesizing algae also contribute, either as mediators or by direct precipitation. Some microorganisms such as endolithic bacteria and algae have a dual role in that their substrate boring activities tend to destroy primary grains but leave micrite rinds and infills in their wake.

Seafloor cementation can take place from the supratidal-intertidal zone (marine vadose zone) (e.g. crusts, beachrock), reef and shallow subtidal platform (mainly within the photic zone), to the outer platform and slope (beyond the photic zone). At greater depths, first aragonite (the aragonite compensation depth – ACD – is 2-3km) then calcite (CCD is 4-6km) begin to dissolve because of high CO₂ partial pressures.

Meteoric environments

Platform and reef deposits exposed by a fall in relative sea level will be subjected to an influx of fresh water. Carbonate deposits subjected to meteoric conditions undergo significant changes to their mineralogy and texture.

In the meteoric environment, groundwater flow through permeable aquifers is driven largely by gravitational potential energy, usually referred to as topography-driven flow. Hydrogeologists have long recognized three distinct zones within meteoric settings; James and Choquette (reference at bottom of the page) included these zones in their early model of carbonate diagenesis:

The watertable, below which all porosity is completely saturated; this is the phreatic zone.
Above the watertable is the unsaturated or vadose zone where pore spaces are mostly air-filled but are periodically wetted by rising watertables (may be seasonal) and infiltration of surface water.
Where aquifers intersect the coast there is a zone of fresh water and seawater mixing. The location of the phreatic mixing zone in relation to the shoreline depends on the hydraulic gradient (or hydraulic head) in the aquifer and how far the aquifer extends offshore. It is not uncommon for fresh water to flow 100s, even 1000s of metres offshore, beneath a platform or shelf.

Aragonite and high-Mg calcite are unstable in fresh water which means that any clasts (skeletal fragments, ooids, foram tests) and cements that contain these minerals will begin to dissolve. In some cases, secondary porosity will form. Also common is the replacement of clasts and cements by low-Mg calcite.

Karst landscapes form during prolonged exposure to meteoric conditions where even the low-Mg calcites dissolve.

Deep burial environments

Deep burial usually refers to the interval below the influence of marine phreatic and meteoric fluids (10s to 100s of metres deep) to several 1000m depth (depending on the geothermal gradient). The principle physical process is compaction that rearranges sediment frameworks and drives interstitial fluids to other parts of the sedimentary basin. The influence of temperature on the promotion and rates of chemical reactions becomes increasingly important with depth.

Fluid composition is primarily modified seawater, modification that can eventually produce saline brines. Many different reactions come to play as burial depth and temperature increase. Compaction enhances pressure solution where significant volumes of rock are dissolved and the solute transferred to other parts of the sedimentary basin; what remains are stylolites. Other reactions involve dehydration of clays, clay mineral transformations (e.g. kaolinite-smectite), and perhaps most significantly, the diagenesis of organic matter. All these reactions contribute to changes in pH and alkalinity that effect calcite and dolomite stability.

Carbonate diagenesis is dominated by precipitation of calcite and ferroan calcite that replace aragonite and high-Mg calcite. Recrystallization of calcite spar tends to mask and even obliterate original depositional fabrics. Dolomitization is also common during burial diagenesis and like calcite overprints earlier grain and cement fabrics (i.e. those formed in the sea floor and meteoric environments.

Companion diagenetic environments

It is important to recognize that the diagenesis of carbonates during burial is not divorced from broadly similar processes taking place in siliciclastic rocks. Precipitation of quartz and clays, and dissolution of feldspar are important determinants for evolving fluid compositions that significantly effect carbonate mineral stability (mostly low Mg-calcite). Organic matter in carbonates and siliciclastics has a profound effect on deep burial diagenetic pathways. Complex organic compounds like kerogens begin to break down at about 60^o C. The rate of organic diagenesis increases markedly at 80^o C – the lower burial temperature limit of the oil-generation window. Important byproducts of these reactions are organic acids that modify pH and control alkalinity. Note too that pH buffering and alkalinity are also influenced by silicate transformations, particularly those involving smectite, kaolinite and illite. The diagram above (from Surdam and others, 1989) summarizes the progression of diagenetic reactions commonly observed in siliciclastic rocks, in relation to organic maturation, organic acids and pH buffering.

Links to other posts in this series:

Mineralogy of carbonates; non-skeletal grains

Mineralogy of carbonates; lime mud

Mineralogy of carbonates; cements

Mineralogy of carbonates; sea floor diagenesis

Mineralogy of carbonates; Beachrock

Mineralogy of carbonates; deep sea diagenesis