Source: http://www.saltworkconsultants.com/blog/tag/13C/
Timestamp: 2019-04-21 10:37:30+00:00

Document:
Oxygen isotope determinations in evaporitic sediments are typically based on: 1) using oxygen held in the water molecule itself (H2O); 2) oxygen in the carbonate anion held in evaporitic dolomites or limestones or; 3) in sulphate from evaporitically precipitated gypsum or anhydrite. Oxygen measures on the water molecule can be co-associated with deuterium (D - heavy hydrogen) determinations. So isotopic sampling of evaporitic limestone and dolomites means carbon isotope values can be co-determined from the same mineral phase (CO3 source). Likewise, with the calcium sulphates, the sulphur isotope is always available for co-study (SO4 source in gypsum or anhydrite).
We have already discussed sulphur and carbon isotope variations in evaporitic settings in the previous two articles (30 April 2018 and 31 May 2018, respectively). So, in this article, we shall look at how oxygen isotope values vary with the co-associated deuterium, carbon and sulphur isotope phases. We focus on three sources for isotope samples (water molecules in a brine, evaporitic carbonate minerals, calcium sulphate minerals) and show that when oxygen values are co-plotted against deuterium, carbon or sulphur isotope values, it becomes a handy tool in defining depositional and diagenetic evolution in a range of evaporitic settings.
The stable isotope community has long known of the potentially extreme effects of evaporation on the isotopic composition of liquids and the residual enrichment of the heavier isotope in the remaining brine. After all, Urey himself applied this knowledge when he demonstrated the existence of deuterium through evaporative enrichment of liquid hydrogen (Urey et al., 1932). Enrichment in heavier isotopes in the residual brine is documented in settings as diverse as evaporating Dead Sea brines (Gat, 1984) and degassing epithermal systems (Zheng, 1990).
As any water (brine) evaporates there is a commensurate preferential escape of the lighter 16O water molecules, this leaves behind an increasing proportion of heavier water molecules containing 18O. Hence, with increasing degrees of evaporation the δ18O signature in the remaining water mass becomes increasingly positive (Figure 1). Co-variance of deuterium with increasing oxygen isotope values in a concentring brine is a long-established observation (Figure 2; Cappa et al., 2003), and defines a type of Raleigh fractionation or distillation.
There is another factor involved in the degree of enrichment of the heavier isotopes of oxygen or deuterium, and that is the humidity of the air above the evaporating brine. Humidity controls the extent of evaporative concentration, and there is a differential level of isotope enrichment in the residual brine tied to changing humidity (Figure 2). It is a response to the lowering of the evaporation rate with increasing humidity. The humidity effect in evaporative settings is documented both experimentally and in natural settings such as modern sabkhas and salinas (Chapter 2 in Warren 2016, for a summary of literature). As a general rule, the lower the humidity, the greater the degree of enrichment of the heavier isotope. Horton et al., (2016) show that δ18OSMOW values of saline lake waters from are often shifted by >+10‰ relative to source waters discharging into the lake (Figure 3, especially 3c).
Up until February 1979, the Dead Sea was a permanently stratified hypersaline water body (see Warren 2016, Chapter 4 for hydrological and sedimentological details). Both the upper and lower water masses were moderately enriched in δ18OSMOW (Figure 4: Gat 1984). After the overturn and mixing the surface waters, the degree of enrichment in δ18O in the surface waters constitutes a balance between the dilution by freshwater influx and the isotope fractionation (enrichment) which accompanies evaporative water loss and vapour exchange with the atmospheric moisture. Gat's modelling of the seasonal cycle and long-term trends of δ18OSMOW in response to the changes in the environmental parameters, shows that the dominant control on isotope enrichment in the surface waters, post overturn, is exercised by the salinity of the surface waters, through its effect on the vapour pressure gradient between the lake's surface and the atmosphere. Interestingly, before the overturn event the upper water mass was more homogenous in terms of salinity and isotope enrichment and its enriched isotope values mostly tracked those of the much more stable and somewhat more saline lower water mass.
Deuterium-oxygen isotope plots of water molecules can also be useful in studying the origin of hydrated salts such as gypsum, but only if there has been minimal postdepositional alternation of the primary precipitate. A classic paper focusing on the composition of structural water held in the gypsum lattice of Messinian (Late Miocene) evaporites of Sicily was published by Bellanca et al., 1986. In Sicily, there are two main types of texture in gypsum-dominated outcrops in the Messinian sub-basins of Sicily (laminated and massive). The laminar gypsum, locally known as balatino, is a shallow-water saltern deposit, the other is a massive form of gypsum typically interpreted as a diagenetic replacement of either primary gypsum of anhydrite.
The different isotopic compositions of hydration water in the two gypsum lithotypes are shown in Figure 6. Laminar gypsum shows a predominance of positive values for both oxygen (range-1.59‰ < δ18OSMOW < +6.02‰) and deuterium (range -7.3‰ < δD < +22.7‰), while both oxygen and deuterium ranges in the massive gypsum are negative (-4.21‰ < δ18OSMOW > -2.23‰; -40.9‰ < δD < -34.4‰).
In Figure 5 the majority of points representative of the laminar gypsum mother waters fall to the right of the meteoric water line of Craig ( 1961) and lie on a path characterised by a positive slope (δD = 3.97δ18OSMOW - 0.59) and includes the SMOW point. Such a distribution is consistent with an origin of the gypsum by direct pre­cipitation from an evaporating solution saturated with respect to gypsum and is close to those of mother waters in recent gypsum samples precipitated in Mediterranean salinas and, there­fore, suggest that the solutions from which the laminar gypsum precipitated were marine waters concentrated by evaporation. A few other examples show δ18O and δD values shifted towards negative values, which indicate stages of dilution with large masses of continental waters poured into the deposition basin during the crystallisation of gyp­sum (Bellanca et al., 1986).
The isotopic makeup of residual water molecules evolving into a brine is not the only phase affected by the chemical consequences of evaporation (Horton et al., 2016; Warren 2016). As any natural water evaporates, its chemistry changes, as concentrating dissolved phases and increasing alkalinity force changes in equilibrium conditions. One of the most obvious consequences of evaporation is the formation of sedimentary evaporites, including brine pool carbonates (e.g. calcite, aragonite, dolomite, trona). The coupled δ18O and δ13C enrichment during evaporation, and the precipitation of endogenic Holocene carbonates is documented and discussed at some length in a number of review papers (Horton et al., 2016; Pierre, 1988).
Horton et al. (2016) document a general tendency for calcites precipitated in lakes located in somewhat less humid climates to show enrichment in the heavier isotope. The observed average lake carbonate δ18OPDB values from the 57 lakes plotted in Figure 6 are more positive than the modelled summer month meteoric water derived calcite δ18O values (Horton et al., 2016). Lake calcites precipitating in humid environments generally plot closer to the 1:1 line, suggesting lakes in these environments are less impacted by evaporative modification. Yet, 46 of the 57 lake records analysed (i.e. 81%) plot to the right of the 1:1 line consistent with evaporative modification of lake water δ18O. Forty-two percent of the lake carbonate δ18O records are >5‰ shifted towards more positive δ18O values than would be expected for summer-month carbonate precipitates derived from unmodified local meteoric water. Although many lakes with vastly different modern aridity index values show similar offsets between modelled and observed δ18O, lakes from currently arid and semi-arid environments have a much larger average δ18O offset (5.4‰) than sub-humid and humid environment lakes (2.0‰).
The dolomite forming lakes of the Coorong region show a similar set of enrichment in both oxygen and carbon isotopes within that type of Holocene dolomite precipitating directly from evaporating surface brines (dolomite Type-A; Rosen et al., 1989; Warren 1990, 2000). The other type of Holocene dolomite in the Coorong lakes (dolomite-B) shows no noticeable C-O covariant trend related to Raleigh distillation (Figure 7a). Type-A dolomite has a heavier oxygen isotope signature than type-B and is 3 - 6‰ heavier in 13C (Figure 7a). Type-A dolomite also has distinct unit cell dimensions (Rosen et al., 1989).
Type A tends to be magnesium-rich with up to 3-mole percent excess MgCO3, while type-B is near stoichiometric or calcian-rich. Type-A dolomite typically occurs in association with magnesite and hydromagnesite, Type B with Mg-calcite. Transmission electron microscopy (TEM) shows that Type A dolomites have a heterogeneous microstructure due to closely spaced random defects, while type B dolomites exhibit a more homogeneous microstructure implying excess calcium ions are more evenly distributed throughout the lattice. TEM studies show that the two types of Coorong dolomite are distinct and are not intermixed with other mineral phases; they are primary precipitates, and not replacements and are not transitional (Miser et al., 1987).
Within the lake stratigraphy the dolomites occupy two distinct positions, Type A dolomites occur as surficial 'yoghurt' textured gels that in each water-filled winter season are washed and blown across the lake surface. By late spring and through summer these surface waters have dried up (summer salinities ≈ 120‰), and the lake sediment surface is a mud-cracked interval of massive carbonate (Warren, 1990; 2016). Type B dolomites occur in the laminated unit that underlies the laminated with signatures implying precipitation from waters with bicarbonates, perhaps showing a stronger strong input from organic materials and are especially prevalent in the more marginward part of the laminated fille where meteoric groundwaters are continually flowing into the edges of the lakes and mixing with lake pore brines.
Figure 7b places these two Coorong dolomites in the context of other areas of primary dolomite accumulations within Holocene carbonate depositional settings. Today sulphate-reducing bacteria or archeal methanogens have been called upon to explain the primary precipitation of dolomite in bacterial biofilms in almost all these other settings. It is not my intention to question the importance of bacterial metabolism in these other dolomite-accumulating settings, only to point out the bicarbonate from which the Coorong type A dolomites have precipitated show a positive and co-variant enrichment in both carbon and oxygen valued that are more typical of evaporative concentration. Evaporative enrichment in carbon values tied CO2 degassing in highly saline waters was documented in the Dead Sea by Stiller et al., 1985 and discussed in last month's article (31 May 2018).
Evaporitic carbonates especially when interbedded with calcium sulphate beds can also dissolve and alter (Warren, 2016; Chapter 7). Evaporite-derived dedolomites are often associated with evaporite dissolution breccias, which indicates the stratigraphic position of the now dissolved calcium sulphate bed that supplied the excess calcium needed to dedolomitise (Lee, 1994; Fu et al., 2008). Dedolomite under this scenario forms via the reaction of calcium sulphate-rich solutions with pre-existing dolomite to produce calcite with magnesium sulphate as a possible byproduct. The latter is rarely preserved, as it is highly soluble, and either remains as dissolved ions in the escaping waters or is quickly redissolved and flushed by through-flowing groundwaters (Shearman et al., 1961). The CaSO4 dissolution process is often driven by meteoric flushing of nearsurface oxidising waters and former ferroan dolomites are preferentially replaced. The resulting calcitised dolomites are outlined by intervals stained red with iron oxides and hydroxides.
With uplift-related (telogenetic) dedolomites the distribution and isotopic composition of dedolomite can reflect variations in the regional hydrology. This can be seen in the dedolomites of the Lower Cretaceous Edwards Group in the Balcones fault zone area of south-central Texas (Ellis, 1985, 1986). The Edwards Group consists of 120-180 metres of porous limestone and dolomite that accumulated on the Comanche shelf in shallow-water subtidal, intertidal, and supratidal marine environments. During early burial diagenesis, carbonate mud neomorphosed to calcitic micrite, aragonite and Mg-calcitic allochems were altered to calcite or were leached, and evaporites formed in tidal-flat sediments. Each of these phases had a characteristic stable isotope signature (Figure 8). Dolomite is widespread and formed in environments ranging from hypersaline to fresh-water as shown by the two isotope clusters in the Edwards dolomite (meteoric versus evaporitic reflux).
Late Tertiary faulting along the Balcones fault zone, tied to Jurassic salt withdrawal, initiated a circulating, fresh-water aquifer system to the west and north of a fairly distinct “bad-water line,” which roughly parallels the Balcones fault zone. To the south of the bad-water line, interstitial fluids remained relatively stagnant and contain over 1000 mg/l dissolved solids. Because of the differences in the chemistry of the interstitial fluids, post-faulting diagenesis in the two zones has been very different.
Water in the bad-water zone can be saturated with respect to calcite, dolomite, gypsum, celestite, strontianite, and fluorite, whereas water in the fresh-water zone is saturated only with respect to calcite. Due to the change in water chemistry, rocks in the fresh-water zone have been extensively recrystallised to coarse microspar and pseudospar, extensive dedolomitization has occurred, and late sparry calcite cements have precipitated. This creates a suite of covariant isotope trends and clusters with the dedolomite showing a distinctive set of carbon and oxygen values relate to soil water influences indicated by calcites with more negative carbon values (Figure 8 indicated by brown shading). In contrast, rocks in the bad-water zone retain fabrics associated with pre-Miocene diagenesis, and there is little or no evidence of widespread dedolomite, indicated by pink shading in Figure 8.
The importance of meteoric diagenesis in the formation of dedolomite in shallow, subsurface telogenetic environments is illustrated by the fact that the Edwards Group had a stable mineralogy of calcite and dolomite before the circulation of fresh water began and drove the precipitation of meteoric spar, microspar and dedolomite. Isotopic values for the dedolomites follow a similar trend to those of the microspars and pseudospars. As with the microspars and pseudospars formed by the entry of telogenetic water, it can be shown that dedolomites are in isotopic equilibrium with Edwards water on a regional scale, which supports the contention that the dedolomites are still forming from crossflows of present-day formation-water (Ellis, 1985).
Thus the δ34S and δ18O values of sulphate evaporites are directly related to the state of the aqueous sulphate reservoir wherever precipitation occurred. A plot of ancient marine CaSO4 evaporites shows the sulphur curve for seawater has varied across time from +30‰ in the Cambrian, to around +10‰ in the Permian and that it increased irregularly into the Mesozoic to its present value of +20‰. Oceanic oxygen isotope values show much less variability. Sulphur is largely resistant to isotopic fractionation during the increasing temperatures associated with burial alteration and transformation (Worden et al., 1997). All of these aspects are discussed in detail in the April 30, 2018 article.
With this knowledge of the relative lack of fractional in the subsurface compared to the much greater susceptibility of oxygen isotopes in the mesogenetic and telogenetic realms let us now look in more detail at the significance of oxygen variation in a variety of sulphate entraining settings.
Isotopically, the effects of dissolution and brine recycling in fracture-filling fibrous gypsum cements of various ages emplaced in a formation's burial evolution can be used define the sequential development of the superimposed diagenetic textures in the original gypsum unit (Figures 9, 10; Moragas et al., 2013). The upper Burdigalian Vilobí Gypsum Unit, located in the Vallès Penedès half-graben (NE Spain) and consists of a 60-m thick succession of laminated-to-banded primary and secondary gypsum. The unit is variably affected by Neogene extension in the western part of Mediterranean Sea. Tertiary extensional events are recorded in the evaporitic gypsum unit as six fracture sets and fills (faults and joints - S1 - S5), which can be linked with basin-scale deformation stages.
Combined structural, petrological and isotopic study of the unit by Moragas et al. (2013) established a chronology of fracture formation and infilling, from oldest to youngest as: (i) S1 and S2 normal faults sets with formation and precipitation of sigmoidal gypsum fibres; (ii) S3 joint sets with perpendicular fibres; (iii) S4 inverse fault sets, infilled by oblique gypsum fibres and associated with thrust-driven deformation of the previous fillings; and (iv) S5 and S6 joint sets tied to later dissolution processes and infilled by macrocrystalline gypsum cements likely related to the telogenetic realm. The fractures provided ongoing pathways for focused fluid circulation within the Vilobí Unit. The oxygen, sulphur and strontium isotope compositions of the original host rock and the various precipitates in the fractures imply ongoing convective recycling processes across the host-sulphates to the fracture infillings, as recorded by a general enrichment trend toward heavier S–O isotopes, from the oldest precipitates (sigmoidal fibres) to the youngest (macrocrystalline cements). The marine strontium signal is mostly preserved in the various postdepositional infillings, unlike the oxygen and to a lesser extent the sulphur isotope signals, which are evolving with the origin and temperature of the waters flowing in the fracture sets (Figure 10).
In any ancient silicified anhydrite nodule or bedded silicified succession, 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 Laño 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 (Figures 10, 11; Gómez-Alday et al., 2002).
Anhydrite nodules at Laño were produced by the percolation of saline marine brines, during a period corresponding to a depositional hiatus. They have δ34S and δ180 mean values of +18.8‰ and +13.6‰, respectively, both consistent with Upper Cretaceous seawater sulphate values. Higher δ34S and δ180 (mean values of + 21.2‰ and 21.8‰ 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 11a). Later calcite replacement and precipitation of geode-filling calcite in the siliceous nodules occurred in both sections, with δ13C and δ180 values indicating the participation of meteoric waters in both regions (Figure 11b). Synsedimentary activity of the Penacerrada diapir (Kueper salt - Triassic), which lies close to the Laño section, played a significant role in driving the local shallowing of the basin and in the formation of the silica in the 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.
This and the previous two articles have underlined the utility of stable isotope samples of brine or precipitates in better understanding the origin of a range of brines and their associated precipitates. But other than the sampling of water molecules in modern brines, the interpretation of all isotope values is equivocal without a petrographic understanding of how and when the sampled textures formed. Stable isotopes of evaporitic minerals with sulphur, carbon and oxygen are the mainstays of isotope work in the study of most evaporite basins, both modern and ancient. Other isotopes that may be useful are 11B and 37Cl, and we shall look at their application to evaporitic sediments in a later blog.
Bellanca, A., and R. Neri, 1986, Evaporite carbonate cycles of the Messinian, Sicily; stable isotopes, mineralogy, textural features, and environmental implications: Journal of Sedimentary Petrology, v. 56, p. 614-621.
Cappa Christopher, D., B. Hendricks Melissa, J. DePaolo Donald, and C. Cohen Ronald, 2003, Isotopic fractionation of water during evaporation: Journal of Geophysical Research: Atmospheres, v. 108.
Ellis, P. M., 1986, Post-Miocene carbonate diagenesis of the Lower Cretaceous Edwards Group in the Balcones fault zone area, south-central Texa, in P. L. Abbott, and C. M. Woodruff, eds., The Balcones escarpment, geology, hydrology, ecology and social development in central Texas, Geological Society of America, p. 101-114.
Fu, Q. L., H. R. Qing, K. M. Bergman, and C. Yang, 2008, Dedolomitization and calcite cementation in the Middle Devonian Winnipegosis Formation in Central Saskatchewan, Canada: Sedimentology, v. 55, p. 1623-1642.
Gat, J. R., 1984, The stable isotope composition of Dead Sea waters: Earth and Planetary Science Letters, v. 71, p. 361-376.
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.
Horton, T. W., W. F. Defliese, A. K. Tripati, and C. Oze, 2016, Evaporation induced 18O and 13C enrichment in lake systems: A global perspective on hydrologic balance effects: Quaternary Science Reviews, v. 131, p. 365-379.
Lee, M. R., 1994, Emplacement and diagenesis of gypsum and anhydrite in the late Permian Raisby Formation, north-east England: Proceedings - Yorkshire Geological Society, v. 50, p. 143-155.
Miser, D. E., J. S. Swinnea, and H. Steinfink, 1987, TEM observations and X-ray structure refinement of a twiined dolomite microstructure: American Mineralogist, v. 72, p. 188-193.
Moragas, M., C. Martínez, V. Baqués, E. Playà, A. Travé, G. Alías, and I. Cantarero, 2013, Diagenetic evolution of a fractured evaporite deposit (Vilobí Gypsum Unit, Miocene, NE Spain): Geofluids, v. 13, p. 180-193.
Pierre, C., 1988, Application of stable isotope geochemistry to the study of evaporites, in B. C. Schreiber, ed., Evaporites and hydrocarbons: New York, Columbia University Press, p. 300-344.
Rosen, M. R., D. E. Miser, M. A. Starcher, and J. K. Warren, 1989, Formation of dolomite in the Coorong region, South Australia: Geochimica et Cosmochimica Acta, v. 53, p. 661-669.
Shearman, D. J., J. Khouri, and S. Taha, 1961, On the replacement of dolomite by calcite in some Mesozoic limestones from the French Jura: Proceedings Geological Association of London, v. 72, p. 1-12.
Stiller, M., J. S. Rounick, and S. Shasha, 1985, Extreme carbon-isotope enrichments in evaporating brines: Nature, v. 316, p. 434.
Testa, G., and S. Lugli, 2000, Gypsum-anhydrite transformations in Messinian evaporites of central Tuscany (Italy): Sedimentary Geology, v. 130, p. 249-268.
Urey, H. C., F. G. Brickwedde, and G. M. Murphy, 1932, A Hydrogen Isotope of Mass 2: Phys. Rev., v. 39, p. 164.
Warren, J. K., 1990, Sedimentology and mineralogy of dolomitic Coorong lakes, South Australia: Journal of Sedimentary Petrology, v. 60, p. 843-858.
Warren, J. K., 2000, Dolomite: Occurrence, evolution and economically important associations: Earth Science Reviews, v. 52, p. 1-81.
Worden, R. H., P. C. Smalley, and A. E. Fallick, 1997, Sulfur cycle in buried evaporites: Geology, v. 25, p. 643-646.
In the next article we shall look at the utility of crossplots of carbon and oxygen isotopes. Stable oxygen isotope values (d18O) crossplotted with respect to carbon isotope values (d13C) from, the same sample creates one of the most widely applied proxies used to infer palaeo-environmental conditions (depositional and diagenetic) in Holocene and ancient carbonate sediments. This is in large part due to kinetic fractionations that occur during evaporation (Leng and Marshall, 2004). It has long been known that as any liquid evaporates, the residual fluid becomes enriched in the less abundant heavy isotope(s) (see Horton et al., 2016 for detailed discussion).
Over the Phanerozoic the standard paradigm for interpreting variations in variations in 13C values from modern and ancient marine carbonate is based on an integration of our understanding of the carbon cycle with the following arguments. Most of the carbon in Earth’s near-surface systems is stored in sedimentary rocks with only about 0.1% in living organisms and the atmosphere-hydrosphere (Figure 1). Oxidized carbon occurs primarily as marine carbonates and reduced carbon as organic matter in sediments. In the carbon cycle, CO2 from the oceans and atmosphere is transferred into sediments as carbonate carbon (Ccarb) or organic carbon (Corg), the former of which monitors the composition of the oceans (Figure 1). The cycle is completed by uplift and weathering of sedimentary rocks and by volcanism, both of which return CO2 to the atmosphere.
It is assumed that the carbon isotopic ratio in calcareous shells of marine organisms is in equilibrium with that of seawater. So as more carbon 12 is held in biomass during times of high primary productivity, and increased burial of organic carbon, calcareous (CaCO3) skeletal materials become enriched in carbon 13. In contrast during periods of low biological productivity and decreased burial of organic carbon, for example following mass extinctions, marine calcareous skeletal materials become enriched in carbon 12.
Hence plotting variations in carbon isotopes in marine carbonates and organic matter over time offers a way to trace the growth of the crustal reservoir of reduced carbon (Des Marais, 1997). That is, the relative abundance of carbon isotopes is controlled chiefly by: 1) equilibrium isotopic effects among inorganic carbon species, 2) fractionation associated with the biochemistry of organic matter, and 3) the relative rates of burial of carbonate and organic carbon in sediments (Condie 2016).
Because organic matter preferentially incorporates 12C over 13C, there should be an increase in the 13C/12C ratio (as measured by δ13C) in buried carbon with time, and indeed this is what is observed (Des Marais, 1997; Worsley & Nance, 1989). δ13Corg increases from values < -40‰ in the Archaean to modern values of -20 to -30‰. On the other hand, seawater carbon as tracked with δ13Ccarb remains roughly constant with time, with δ13Ccarb averaging about 0%.
δin represents the isotopic composition of carbon entering the global surface environment comprised of the atmosphere, hydrosphere, and biosphere. The right side of the equation represents the weighted-average isotopic composition of carbonate (δ13Ccarb) and organic (δ13Corg) carbon buried in sediments, and fcarb and forg are the fractions of carbon buried in each form (fcarb = 1 - forg). For timescales longer than 100 Myr, δin = -5‰, the average value for crustal and mantle carbon (Holser et al., 1988). Thus, where values of sedimentary δ13Ccarb and δ13Corg can be measured, it may be possible to determine forg for ancient carbon cycles. Higher values of δ13Ccarb indicate either a higher value of forg or a greater negativity of average δ13Corg.
During the Phanerozoic, there are several peaks in δ13Ccarb, the largest at about 110, 280, 300, 400, and 530 Ma (Figure 3). These peaks are widely interpreted to reflect an increase in burial rate of organic carbon (Des Marais et al., 1992; Frakes et al., 1992). This is because organic matter selectively enriched in 12C depletes seawater in this isotope, raising the δ13C values of seawater. In the late Paleozoic (300-250 Ma), the maxima in δ13Ccarb correspond to the rise and spread of vascular land plants, which provided a new source of organic debris for burial (Condie 2106, Berner, 1987, 20 01). Also conducive to preservation of organic remains at this time were the vast lowlands on Pangea, which appear to have been sites of widespread swamps where bacterial decay of organic matter is minimized. The drop in δ13Ccarb at the end of the Permian is not understood. Perhaps, large amounts of photosynthetic O2 generated by Carboniferous forests led to extensive forest fires that destroyed large numbers of land plants in the Late Permian (Condie, 2016). However, the reasons for the oscillations in δ13Ccarb are not yet unequivocally resolved and, as in all sciences, the tenet "...perceived correlation does not necessarily equate to causation"must always be at the forefront in the scientific mindset.
Across the Precambrian and the Phanerozoic, the initiation of glaciation on a global scale, as in the Cryogenian ‘Snowball Earth’, has been interpreted to be dependent on parameters like the latitudinal extent of continents and oceanic circulations (Figure 4; Condie, 2016). The main drive for an onset of global glaciation is believed to be the lowering of atmospheric CO2. It likely also requires a continental landmass to be covering one of the earth's polar positions. More recently, cooling related to an increase in the earth's albedo due to widespread evaporites (saline giants) has been added to the list of possible drivers to the onset of glaciation.
Climate modelling studies imply that CO2 concentrations as low as 100–150 ppm are required to initiate global glaciation (e.g. Liu et al., 2013; Feulner and Kienert, 2014). One potential cause of lowered CO2 is drawdown of CO2 during intense silicate weathering in equatorial regions (Hoffman and Schrag, 2002; Goddéris et al., 2003). Photosynthesis provides another mechanism for CO2 drawdown, via conversion of CO2 to O2 and rapid burial of organic carbon, which is reflected in a positive δ13C excursion for carbonates (Pierrehumbert et al., 2011). Additionally, long term cloud cover (Feulner et al., 2015), fluctuations in atmospheric-ocean heat transport, the earth's albedo, or solar luminosity (Pierrehumbert et al. (2011) are also proposed as potential causes of the onset of glaciation (ice-house mode climate).
In a recent paper, Schmid 2017 focused on the cause of the Bitter Springs carbon isotope anomaly, she argues the cause of the pre-glacial, globally recognised, carbon and oxygen isotope variations in carbonate sediments tied to the Bitter Springs anomaly is a response to widespread fractional evaporation of dissolved CO2. This carbon isotope anomaly ties to a well defined correlation with the distribution of Neoproterozoic evaporite basins. She also shows volcanism occurred during the onset of the Bitter Springs Stage (811–788 Ma) and associated widespread evaporite distribution across Australia.
Schmid (op. cit.) argues that the albedo effect began with of the widespread deposition of Rodinian supercontinent evaporites in very shallow marine to epicontinental sedimentary successions beginning ≈810 Ma, increased siliciclastic redbed weathering. This and continuing evaporite deposition and exposure between ≈780 and 720 Ma drove a worldscale increase in Earth's albedo. Such highly reflective salt deposits defined a saline giant across an area that today covers one-third of the Australia continent. Thus, this and other penecontemporaneous saline giants over the Rodinian supercontinent played a potentially significant role in the onset of atmospheric cooling via a significant increase in albedo (Figure 5). These salt beds occur in periods that typify the onset of local (750 Ma) and then global glaciation (720 Ma).
Schmid (2017) goes on to note that the degree of evaporation in the Bitter Springs group sediments is related to the δ13C signature in variably concentrated waters (Figure 6). That is the Tonian Bitter Springs Group (≈830–750 Ma), within the Amadeus Basin in central Australia consists of thick halite and sulphate evaporite accumulations and associated carbonates. The deposition of halite occurred in shallow marine, lagoon (salina) environment (Gillen Formation), and developed into sulphate-dominated supratidal sabkha during sea level regression (Johnnys Creek Formation). The overall regression was interrupted by a transgressive phase lasting at least 20 Ma and leading to deposition of basin-wide stromatolitic dolostone (Loves Creek Formation). The salinity and high evaporation is reflected in positive δ13C in the intercalated carbonates (+4 to +6‰ VPDB) of the evaporitic units, while the shallow marine stromatolitic incursion of the Loves Creek Formation (−2‰ δ13C) show typical marine carbonate isotopic values (Figure 7).
This salinity controlled isotopic separation supports the observations of Stiller et al. (1985) who noted extreme enrichment of 13C in the dissolved inorganic carbon pool in evaporating brines up with δ13C values of up to + 16.5‰ under natural abiotic, oxic conditions in Dead Sea evaporation ponds (Figure 7). The systematic increase in 13C values in highly evaporated waters from the various bittern ponds of the Dead Sea Saltworks is thought to result from a nonequilibrium gas-transfer isotope fractionation. The process of ongoing evaporation leads to CO2 loss within the evaporative brine as less and less gas can held in solution (see Warren 2016, Chapter 9). CO2 exchange in a concentrating surface brine occurs directly between the water column and air, resulting in direct CO2 loss through evaporation. In a sabkha environment. CO2 is released from the hypersaline groundwater through sediments before being released to air as evaporites may form intrasediment precipitates. Overall, atmospheric CO2 uptake in hypersaline settings fed by shallow marine water is diminished compared to the normal marine settings.
Precipitated carbonates modern salinas and sabkhas are mainly aragonite, and formed in association with such evaporative brine, are consistently13C enriched, as seen in nearby Solar Lake and Sabkha Gavish (Figure 2; Stiller et al., 1985; Schidlowski et al., 1984). In a similar fashion, Palaeoproterozoic interbedded shallow marine carbonates, redbeds and evaporites have values up to δ13C + 17.2‰ (Melezhik et al., 1999). Permian and Triassic (Schmid et al., 2006a) redbeds and evaporite sequences also have 13C-rich carbonates (up to +7‰) and enrichment is partly attributed to evaporation and associated CO2 loss (Beauchamp et al., 1987). In modern oceans, atmospheric CO2 is consumed by biological activity and carbonate production originates from mainly marine organisms, leading to near atmospheric to organic negative δ13C signatures in the precipitated sediment(Andersson, 2013).
If increasing salinity leads to unfavourable conditions for photosynthesising organisms to survive (Lazar and Erez, 1992), carbonate through to bittern precipitation becomes increasingly abiotic and evaporation driven, especially at the upper end of the evaporation series. The loss of Ca during evaporation of a brine, via aragonite and calcium sulphate precipitation, leads to an increase in Mg/Ca ratio and an increase in residual brine density. This can result in primary dolomite precipitation or widespread reflux dolomitisation (Schmid et al., 2006, Warren 2000, 2016.
In summary, the typical δ13C signature in normal marine carbonate sediment across much of geological time centres around 0 ‰ and ranges between a few parts per mille on either side of the zero line reflecting precipitation by calcifying and photosynthesising organisms (e.g. algae), while abiotic, evaporation induced carbonates tend to have δ13C values above +1‰. More positive δ13C values (+4 to +6‰) tend to typify dominantly abiotic carbonates (and local methanogenic carbonates with even more positive values) and support the notion of evaporation-driven 13C-enrichement in times of widespread evaporitic epeiric and basinwide carbonates. In the Precambrian, widespread marine stromatolitic units such as, algal Loves Creek Formation reflects δ13C values for biogenic carbonate precipitation under shallow marine, non-hypersaline conditions. The change from a shallow hypersaline lagoon towards evaporitic mudflats and salterns suggests an increase in aridity and continentality/hydrographic isolation, with associated more positive δ13C values.
The Bitter Springs Group chemostratigraphy has been correlated globally and the negative excursion was named previously after this unit (Bitter Springs Stage anomaly). However, the mechanism of evaporation-driven fractionation of δ13C is different from the commonly proposed inorganic-organic carbon fractionation, and challenges the views on interpreting global chemostratigraphic anomalies or excursion and their cause. Evaporite basins covered vast regions worldwide prior to the Sturtian glaciation, e.g. the Australian evaporites would have covered a third of the continent. The light surface of evaporites and associated carbonates would have had a high albedo and effectively cause less surface heat absorption. This subsequently would have triggered temperature decrease on a continental and possibly global scale. The Schmid paper hypothesises that the deposition of evaporites worldwide would have contributed to global cooling starting ≈100 Ma prior to Snowball Earth and would have played an important role in the onset of global glaciation.
Andersson, A.J., 2013. The oceanic CaCO3 cycle. In: T. Holland (Editor), Treatise on Geochemistry, 2nd ed. Elsevier, pp. 519-542.
Beauchamp, B., Oldershaw, A.E. and Krouse, H.R., 1987. Upper Carboniferous to Upper Permian 13C-enriched primary carbonates in the Sverdrup Basin, Canadian Arctic: comparisons to coeval western North American ocean margins. Chem. Geol. , 65: 391-413.
Berner, R.A., 1987. Models for carbon and sulfur cycles and atmospheric oxygen; application to Paleozoic geologic history. American Journal of Science, 287: 177-196.
Berner, R.A., 2001. Modeling atmospheric O2 over Phanerozoic time. Geochimica et Cosmochimica Acta, 65: 685-694.
Condie, K.C., 2016. Earth as an Evolving Planetary System (3rd edition). Elsevier, 350 pp.
Des Marais, D.J., 1997. Isotopic evolution of the biogeochemical carbon cycle during the Proterozoic Eon. Organic Geochemistry, 27(5): 185-193.
Des Marais, D.J., Strauss, H., Summons, R.E. and Hayes, J.M., 1992. Carbon isotope evidence for the stepwise oxidation of the Proterozoic environment. Nature, 359: 605-609.
Feulner, G., Hallmann, C. and Kienert, H., 2015. Snowball cooling after algal rise. Nat. Geosci. , 8: 659-662.
Feulner, G. and Kienert, H., 2014. Climate simulations of Neoproterozoic snowball Earth events: similar critical carbon dioxide levels for the Sturtian and Marinoan glaciations. Earth Planet. Sci. Lett., 404: 200-205.
Frakes, L.A., Francis, J.E. and Syktus, J.L., 1992. Climate modes of the Phanerozoic. Cambridge University Press, New York, 274 pp.
Goddéris, Y., Donnadieu, Y., Nédélec, A., Dupré, B., Dessert, C., Grard, A., Ramstein, G. and François, L.M., 2003. The Sturtian ‘snowball’ glaciation: fire and ice. Earth Planet. Sci. Lett. , 211: 1-12.
Hoffman, P.F. and Schrag, D.P., 2002. The snowball Earth hypothesis: testing the limits of global change. Terra Nova, 14: 129-155.
Holser, W.T., Schidlowski, M., Mackenzie, F.T. and Maynard, J.B., 1988. Geochemical cycles of carbon and sulfur. In: C.B. Gregor, R.M. Garrels, F.T. Mackenzie and J.B. Maynard (Editors), Chemical cycles in the evolution of the earth. John Wiley, New York, pp. 105–173.
Horton, T.W., Defliese, W.F., Tripati, A.K. and Oze, C., 2016. Evaporation induced 18O and 13C enrichment in lake systems: A global perspective on hydrologic balance effects. Quaternary Science Reviews, 131: 365-379.
Lazar, B. and Erez, J., 1992. Carbon geochemistry of marine-derived brines: I. 13C depletions due to intense photosynthesis. Geochim. Cosmochim. Acta, 56: 335-345.
Leng, M.J. and Marshall, J.D., 2004. Paleoclimate interpretation of stable isotope data from lake sediment archives. Quaternary Science Reviews, 23(811-831).
Liu, Y., Peltier, W.R., Yang, J. and Vettoretti, G., 2013. The initiation of Neoproterozoic ‘‘snowball” climates in CCSM3: the influence of paleocontinental configuration. Climate Past, 9: 2555-2577.
Melezhik, V.A., Fallick, A.E., Medvedev, P.V. and Makarikhin, V.V., 1999. Extreme 13Ccarb enrichment in ca. 2.0 Ga magnesite-stromatolite-dolomite-‘red beds’ association in a global context: a case for the world-wide signal enhanced by a local environment. Earth Sci. Rev., 48: 71-120.
Pierrehumbert, R.T., Abott, D.S., Voigt, A. and Koll, D., 2011. Climate of the neoproterozoic. Annu. Rev. Earth Planet. Sci., 39: 417-460.
Schidlowski, M., Matzigkeit, U. and Krumbein, W.E., 1984. Superheavy organic carbon from hypersaline microbial mats; Assimilatory Pathway and Geochemical Implications. Naturwissenschaften, 71(6): 303-308.
Schmid, S., 2017. Neoproterozoic evaporites and their role in carbon isotope chemostratigraphy (Amadeus Basin, Australia). Precambrian Research, 290: 16-31.
Schmid, S., Worden, R.H. and Fisher, Q., 2006. Carbon isotope stratigraphy using carbonate cements in the Triassic Sherwood Sandstone Group: Corrib Field, west of Ireland. Chem. Geol., 225: 137-155.
Stiller, M., Rounick, J.S. and Shasha, S., 1985. Extreme carbon-isotope enrichments in evaporating brines. Nature, 316: 434.
Swart, P.K., 2015. The geochemistry of carbonate diagenesis: The past, present and future. Sedimentology, 62(5): 1233-1304.
Warren, J.K., 2000. Dolomite: Occurrence, evolution and economically important associations. Earth Science Reviews, 52(1-3): 1-81.
Worsley, T.R. and Nance, R.D., 1989. Carbon redox and climate control through Earth history: A speculative reconstruction. Paleogeography, Paleoclimatology, Paleoecology, 75: 259-282.

References: v. 
 v. 
 v. 
 v. 
 v. 
 v. 
 v. 
 v. 
 V. 
 v. 
 v. 
 v. 
 v. 
 v. 
 v. 
 v. 
 v. 
 v.