Salty Matters

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

Evaporite interactions with magma Part 2 of 3: Nature of volatile exhalations from saline giants?

John Warren - Saturday, March 16, 2019

 

Introduction

This article discusses general mechanisms of earth-scale volatile entry into the ancient atmosphere during events that involved rapid and widespread heating of saline giants. It develops this notion by looking at whether volumes of volatiles escaping to the atmosphere are enhanced by either the introduction of vast quantities of molten material to a saline giant or the thermal disturbance of that salt basin by bolide impacts. This begins a discussion of the contribution of heated evaporites in two (or three if the Captitanian is counted as a separate event) of the world's five most significant extinction events. It also looks at possible evaporite associations with a substantial bolide impact that marks the end of the Cretaceous. The next article presents the geological details and implications of the various magma-evaporite-volatile associations tied to major extinction events.

As we have seen for evaporite interactions with giant and supergiant volumes of commodities in particular deposits, such as hydrocarbons, base metals (Cu, Pb-Zn and IOCG deposits) evaporites do not form a commodity accumulation. But if evaporites are involved in the accumulation and enrichment processes, the size and strength of the accumulation are much improved. Because of their high reactivity compared to the kinetic stability at and near  thelithosphere's surface across most other lithologies, evaporite act not as creators of enrichment but as facilitators of enrichment (Warren, 2016 Chapters 9, 10, 14, 15 and Salty Matters, March 31, 2017).


End-Permian event

The end-Permian extinction event, colloquially known as the Great Dying, occurred around 252 Ma (million years) ago, and defines the boundary between the Permian and Triassic geologic periods, as well as between the Palaeozoic and Mesozoic eras. It is the Earth's most severe extinction event, when up to 96% of all marine species, 70% of terrestrial vertebrate species disappeared (Table 1, Figure 1). It also involves the only known mass extinction of a number of insect species (≈25%). Some 57% of all biological families and 83% of all genera became extinct. The end-Cretaceous extinction, which marks the demise of dinosaurs, is less severe, although it probably has a stronger hold on the western zeitgeist, while on land, the end-Triassic event marks the ascendancy of the dinosaurs.


Suggested mechanisms driving the end-Permian extinction event include; massive volcanism centred on the Emeishan and Siberian Traps and the ensuing coal or gas fires and explosions, along with a runaway greenhouse effect that was triggered by temperature increases in marine waters (Figure 2). It also may have involved one or more large meteor impact events and a rise in oceanic water temperatures that drove a sudden release of methane from the sea floor due to methane-clathrate dissociation.

The end-Permian event follows on closely from the Capitanian (Emeishan) extinction event when in south China fusulinacean foraminifers and brachiopods lost 82% and 87% of species, respectively (Bond et al., 2015). Proximity in time of the two events may explain why the breadth of the end-Permian extinction event was so severe. The Earth's biota was still recovering from the Emeishan event when the vicissitudes of the End-Permian calamity further decimated the world's biota.

Both the Emeishan and end-Permian extinction events tie to elevated mercury levels in sediments that encompass their respective boundaries (Grasby et al., 2016). Astride both boundaries, the mercury stratigraphy shows relatively constant background values of 0.005–0.010 μg g–1. However, there are notable spikes in Hg concentration over an order of magnitude above background associated with the two extinction events. The Hg/total organic carbon (TOC) ratio shows similar large spikes, indicating that they represent a real increase in Hg loading to the environment. These Hg loading events are associated with enhanced Hg emissions created by the outflows of the Emeishan and end-Permian large igneous province (LIP) magmas.

Interestingly, there is indirect evidence for a synchronous antipodeal impact crater that some argue may have instigated the Siberian volcanism, in much the same way that the end-Cretaceous bolide impact on the Yucatan Peninsula is considered by some to be the antipodeal driver of the Deccan Trap volcanism (von Frese et al., 2009). Other contributing, but likely more gradual tiebacks to the Great Dying, include sea-level variations, increasing oceanic anoxia, increasing aridity tied to the accretion of the Pangean supercontinent, and shifts in ocean circulation driven by climate change (Figure 2).

End-Triassic event

The end-Triassic extinction event, some 201.3 Ma, defines the Triassic-Jurassic boundary. In the oceans, a whole class (conodonts) and 23-34% of marine genera disappeared. On land, all archosaurs other than crocodylomorphs (Sphenosuchia and Crocodyliformes) and Avemetatarsalia (pterosaurs and dinosaurs), some remaining therapsids, and many of the large amphibians became extinct. About 42% of all terrestrial tetrapods went extinct (Figure 3). This event vacated terrestrial ecological niches, allowing the dinosaurs to assume the dominant roles in the Jurassic period. It happened in less than 10,000 years and occurred just before the Pangaean supercontinent started to break apart (Tanner, 2018).


The extinction event marks a floral turnover as well. About 60% of the diverse monosaccate and bisaccate pollen assemblages disappear at the T-J boundary, indicating a significant extinction of plant genera. Early Jurassic pollen assemblages are dominated by Corollina, a new genus that took advantage of the empty niches left by the extinction.

Worldwide the end-Triassic extinction horizon is marked by perturbations in ocean and atmosphere geochemistry, including the global carbon cycle, as expressed by significant fluctuations in carbon isotope ratios (Korte et al., 2019). At this time the Central Atlantic Magmatic Province (CAMP) volcanism triggered environmental changes and likely played a crucial role in this biotic crisis (Schoene et al., 2010). Biostratigraphic and chronostratigraphic studies link the end-Triassic mass extinction with the early phases of CAMP volcanism, and notable mercury enrichments in geographically distributed marine and continental strata are shown to be coeval with the onset of the extrusive emplacement of CAMP (Percival et al. 2017; Marzoli et al., 2018). Sulphuric acid induced atmospheric aerosol clouds from subaerial CAMP volcanism can explain a brief, relatively cool seawater temperature pulse in the mid-paleolatitude Pan-European seaway across the T–J transition. The occurrence of CAMP-induced carbon degassing may explain the overall longterm shift toward much warmer conditions.

End-Cretaceous event

The end-Cretaceous extinction event defines Cretaceous-Tertiary (K–T) boundary, and was a sudden mass extinction event some 66 million years ago. Except for some ectothermic species, such as the leatherback sea turtle and crocodiles, no tetrapods weighing more than 25 kilograms survived. The K-T event marked the end of the Cretaceous period and with it, the entire Mesozoic Era, opening the Cenozoic Era.

A wide range of species perished in the K–T extinction, the best-known being the non-avian dinosaurs. It also destroyed a plethora of other terrestrial organisms, including certain mammals, all pterosaurs, some birds, lizards, insects, and plants. In the oceans, the extinction event killed off plesiosaurs and the giant marine lizards (Mosasauridae) as well as devastating fish, sharks, molluscs (especially ammonites, which became extinct) populations, and many species of plankton. It is estimated that 75% or more of all species on Earth vanished in the end-Cretaceous event.

In its wake, the same extinction event also provided evolutionary opportunities as many groups underwent remarkable adaptive radiation—sudden and prolific divergence into new forms and species within the disrupted and emptied ecological niches. Mammals in particular diversified in the Paleogene, evolving new forms such as horses, whales, bats, and primates. Birds, fish, and perhaps lizards also radiated in newly vacant niches.


In the geologic record, the K–T event is marked by a thin layer of sediment called the K–Pg (Cretaceous - Paleogene) boundary, that is found throughout the world in both marine and terrestrial rocks. The boundary clay shows high levels of the metal iridium and is widely interpreted as indicating the impact of a massive comet or asteroid 10 to 15 km (6 to 9 mi) wide some 66 million years ago (Figure 4a,b). The impact devastated the global environment, mainly through a lingering impact winter, which halted photosynthesis in plants and plankton.

The impact hypothesis, also known as the Alvarez hypothesis (Alvarez et al., 1980), was bolstered by the discovery of the 180-kilometer-wide (112 mi) Chicxulub crater in the Gulf of Mexico in the early 1990s, which provided conclusive evidence that the K–Pg boundary clay represented debris from an asteroid impact. In a 2013 paper, Paul Renne dated the impact at 66.043±0.011 million years ago, based on argon-argon dating (Renne, 2013). He went on to conclude that the main end-Cretaceous mass extinction event occurred within 32,000 years of this date. A 2016 drilling project into the Chicxulub peak ring, confirmed that the peak ring was comprised of granite, likely ejected within minutes from deep in the earth, but the well contained hardly any anhydrite/gypsum, the usual sulphate-containing seafloor rock across the region (Figure 4a, b). As we shall see in part 3, the missing CaSO4 was vaporised in the impact and dispersed as sulphurous aerosols into the atmosphere, causing longer-term deleterious effects on the climate and food chain. Another causal or contributing factors to the end-Cretaceous extinction event may have been the synchronous outflows of the Deccan Traps and other volcanic eruptions, so driving climate change, and possibly sea level change (von Frese et al., 2009).

Volatiles released when cooking saline giants and associated organic-rich sediments

Particular sets of assimilations and metamorphic alterations of evaporites occur within the explosive milieu associated with both igneous interactions and pressurised heating of salts tied to a bolide impact. Any carbonate and organic matter layers present in the saline sequence or adjacent strata generates additional volatiles that will quickly enter the earth's atmosphere. Figure 5 is a schematic of the estimated amount of volatiles released during contact metamorphism of different types of sedimentary rocks in contact with an igneous sill or magma body (after Ganino et al., 2009; Pang et al., 2013). More catastrophic volumes of similar volatile suites enter the atmosphere if a large bolide impacts a region underlain by a saline giant.


Hence, salty interactions must be considered and quantified when attempting to understand earth-scale environmental changes whenever large evaporite masses are caught up in regions of LIP emplacement or bolide impact. In such areas:

  • Basalt and granitoids do not release large volumes of volatiles, as compared to the amounts of volatiles that are released by the heating or assimilation of saliniferous country rock (heat transfer and hydrothermal circulation).
  • Most porous sandstones and organic-lean shales caught up in a contact aureole or consumed in a magma, release water vapour; a release that has little effect on global climate.
  • During desulphation of a magma, gypsum or anhydrite masses are assimilated into a rising magma chamber or the emplacement of a thick sill. If anhydrite beds are consumed (melted and absorbed) by a magma batholith, the reaction releases abundant SO2 constituting up to 47 wt% of the bedded sulphate (Gorman et al., 1984). Direct melting requires high temperatures (≈ 1300- 1400 °C). Such widespread desulphation of thick Devonian anhydrite beds occurred during the emplacement of the supergiant Noril'sk nickel deposit in Siberia (Black et al., 2014; Warren, 2016, Chapter 16).
  • But such elevated temperatures (≈1400°C) are not typical of most contact aureoles where a sill or dyke intrudes anhydritic country rock. However, similar high-volume SO2 releases can proceed at temperatures as low as 615°C if the anhydrite is impure and contains interlayers rich in organics and hydrocarbons (e.g., West and Sutton, 1954; Pang et al., 2013). This is especially so if the interacting calcium sulphate is gypsum (hydrated salt) rather than anhydrite. Experiments by Newton and Manning (2005) demonstrated that the solubility of anhydrite increases enormously with NaCl activity (salinity) in hydrothermal solutions at ≈600 to 800°C (Figure 6).


  • Pure limestone contains large amounts of CO2, but like anhydrite the thermal decomposition of limestone or dolomite into CaO, MgO and CO2 takes place at high temperatures (>950 °C) that are typical when blocks of sedimentary carbonate are assimilated into a magma chamber, but less typical of contact aureoles tied to dykes and sills. Impure limestones can release large amounts of CO2 (up to 29 wt%) during the formation of calc-silicates in the contact aureole at moderate temperatures of 450–500 °C. As early as 1940, Bowen documented the release of CO2 by decarbonisation reactions during progressive metamorphism of siliceous dolomites (Bowen, 1940)
  • Likewise, devolatilization of fine-grained calcareous and saline sedimentary rocks during contact metamorphism directly generates fluids rich in CO2 (i.e., decarbonisation) and SO2 (i.e., desulphatation), which in theory can enter the magmatic system.
  • When heated at a relatively low temperature (<300-400 °C), contact metamorphism and hydrothermal leaching of bituminous halite and organic-carbon-rich saline mudstones releases large volumes of chlorohalogens and methane (Visscher et al., 2004; Beerling et al., 2007). Halocarbon compounds (aka halogenated hydrocarbons) are chemicals in which one or more carbon atoms are linked by covalent bonds with one or more halogen atoms (fluorine, chlorine, bromine or iodine). Methyl chloride (CH3Cl) and methyl bromide (CH3Br) are commonplace halocarbons when a halite-dominant saline giant interacts with igneous sill emplacement. When thermally-derived chlorohalogens enter the upper atmosphere, they tend to be reactive and will degrade ozone.
  • Buring coal and coal gas release abundant CO2. Depending on its grade, coal can ignite at temperatures between 400-530°C. Methane will auto-ignite at temperatures around 550-600°C and in an oxygenated setting produces large volumes of carbon dioxide and water vapour. Flashpoints are much lower than these ignition temperatures.
  • Sulphidic (pyritic) sediments release abundant SO2 when heated at lower temperatures (<400°C).
  • Heating of hydrated salts at moderate temperatures (90-250°C) can release pressurised pulses of hypersaline chloride or sulphate brine, with the dominant ionic proportions dependent on predominant hydrated salt; e.g., carnallite incongruently alters as it releases an MgCl2 brine, gypsum incongruently alters as it releases a Ca-SO4 brine (see part 1). Such pressurised pulses are essential in the generation of explosive breccia pipes sourced at the sill penetration level in the hydrated evaporite interval (discussed in detail for the Siberian Traps in part 3).
  • Getting volatiles into the atmosphere

    When a saline giant is heated during emplacement of a large igneous province (LIP) or during the impact of a large bolide, it and adjacent carbonates and organic-rich mudstones release large volumes of volatiles that can have short and long term harmful effects on the Earth's biosystems (Black et al., 2012, 2014; Jones et al., 2016; Part 3 this series). The volume of volatiles released to the atmosphere by these interactions, especially sulphurous products (SO2, H2S), thermogenic CH4, organohalogens and CO2, are considered primary contributors to three or four of the major extinction events outlined in Figure 1, and perhaps others, as discussed in part 3.

    Height and volume of various volatile injections into the layers of Earth's atmosphere controls the longevity and intensity of climatic effects and are tied to the chemistry of particular volatiles (Figure 7; Textor et al., 2003; Robock, 2000). The low concentration of water in typical modern volcanic plumes results in the formation of relatively dry aggregates entering the atmosphere. More than 99% of these aggregates are frozen because of their fast ascent to low-temperature regions of the atmosphere. With increased salinities, the salinity effect increases the amount of liquid water attaining the stratosphere by one order of magnitude, but the ice phase is still highly dominant. Consequently, the scavenging efficiency for HCl is very low, and only 1% is dissolved in liquid water.


    Scavenging by ice particles via direct gas incorporation during diffusional growth is a significant process for volatile transport. The salinity effect increases the total scavenging efficiency for HCl from about 50% to about 90%. The sulfur-containing gases SO2 and H2S are only slightly soluble in liquid water; however, these gases are incorporated into ice particles in the atmosphere with an efficiency of 10 to 30%. Despite scavenging, more than 25% of the HCl and 80% of the sulphur gases reach the stratosphere during a more intense modern explosive eruption because most of the particles containing these species are typically lifted there by the force of the eruption (Figure 7b).

    Sedimentation of the particles tends to remove the volcanic gases from the stratosphere. Hence, the final quantity of volcanic gases injected in a particular eruption depends on the fate of the particles containing them, which is in turn dependent on the volcanic eruption intensity and environmental conditions at the site of the eruption.

    Today, volcanically-derived SO2 and H2S are the dominant sources for sulphur species in the atmosphere (Jones et al., 2016; Robock, 2000). Conversion of SO2 to aerosols is one of the critical drivers of climatic cooling during recent eruptions (Figure 7a; Robock, 2000). For SO2 to be effective in causing cooling in the atmosphere, escaping hydrogen sulphide quickly oxidises to SO2. Over hours to weeks following its eruptive escape the ongoing reaction of SO2 with atmospheric H2O forms a H2SO4 (sulphuric acid) aerosol, and this is a major cause of the acid rains tied to volcanism (Figure 7a, b).

    Tropospheric sulphate aerosols have an atmospheric lifetime of a couple of weeks due to the rapid incorporation as precipitation into the hydrological cycle (Figure 7b; Robock, 2000). However, if the intensity of the escaping volatile plume is capable of injecting sulphurous material above the tropopause into the stratosphere, then due to the lack of removal by precipitation, the lifetimes of sulphurous aerosols and the associated cooling effects are considerably extended (years rather than weeks: Figure 7a versus 7b).

    Modern eruptions

    World-scale cooling has been observed following a number of modest (by large igneous province standards) volcanic eruptions over the past few centuries (Figure 8; Bond and Wignall, 2014; Sigurdsson, 1990; and references therein). A recent example is provided by the Mount Pinatubo eruption of 1991, which injected 20 megatons of SO2 more than 30 km into the stratosphere. The result was a global temperature decrease approaching 0.5 °C for three years (although this cooling was probably exacerbated contemporaneous Mount Hudson eruption in Chile). One of the largest historical eruptions occurred in 1783-1784 from the Laki fissure in Iceland when a ≈15 km3 volume of basaltic magma was extruded, releasing ≈122 Mt of SO2, 15 Mt of HF, and 7 Mt of HCl. Laki’s eruption columns extended vertically up to 13 km, injecting sulfate aerosols into the upper troposphere and lower stratosphere, where they reacted with atmospheric moisture to produce ≈200 Mt of H2SO4. This aerosol-rich fog hung over the Northern Hemisphere for five months, leading to short-term cooling, and harmful acid rain in both Europe and North America. Additionally, HCl and HF emissions damaged terrestrial life in Iceland and mainland Europe, as this low-level fluorine-rich haze stunted plant growth and acidified soils.

    By causing or aiding in the collapse of food chains during the more intense sulphurous releases involved in the heating of large volumes of anhydrite held in ancient saline giants, vast quantities of acid rain may have killed much of the vegetation on land and photosynthetic organisms in the oceans during the three extinction events discussed in part 3.


    Halocarbons

    For halocarbons to form in a volcanic eruption requires the combination halogens with organic matter/methane or other hydrocarbons. We shall consider the levels and origins of two of the more common halocarbons in today's atmosphere; methyl chloride (CH3Cl) and methyl bromide (CH3Br) although many other species of halogenated hydrocarbons are present both naturally and anthropogenically (Schwandner, 2002; Visscher et al., 2004).

    The average Cl concentration of the Earth has been estimated to be 17 ppm (Worden, 2018 and references therein). Chlorine is the dominant anion in seawater, most modern and ancient evaporite beds and associated brines. Chlorine is present in most igneous rocks at low concentrations with little difference in level shown between granite and basic igneous rocks (both have a Cl- concentration of about 0.02%). However, igneous glass typically has higher Cl concentrations (≈0.08%). Chlorine is concentrated within any residual vapour phase during volcanic eruptions so can be independent of the volatiles created by heating of saline giants. Without the latter, the contribution of volcanically-erupted Cl to the atmosphere is still considerable. For example, the estimated current global volcanic emission of Cl is between 0.4 and 170 mt/year, while individual eruptions can produce hundreds of kilotons of Cl. For example, in 1980, St Helens emitted 670 kt of Cl into the atmosphere.

    In crystalline igneous rocks Br is found at low concentrations, typically <1 ppm in mid-ocean ridge basalts (MORB) (Worden, 2008 and references)). The average Br concentration of the Earth has been estimated to be 0.05 ppm. Chlorine/Bromine ratios are typically between 200 and 1000 in igneous rocks. Bromine is, however, found at relatively high concentrations (up to 300 ppm) in melt inclusions and matrix glass in acid igneous rocks since it is a highly incompatible element that does not easily sit within silicate, oxide or sulphide minerals. Bromine is concentrated within any residual vapour phase during volcanic eruptions. Based on experimentally-derived fractionation factors for halogens in volcanic materials, crustal average halogen concentrations, and measured amounts of Cl emitted from volcanoes, it can be concluded that the contribution of volcanically-erupted Br to the atmosphere is considerable. For example, the estimated current global volcanic emission of Br is between 2.6 and 78 kt while individual eruptions (e.g., St Helens in 1980) can emit 2.4–5.6 kt.

    The hinterlands of sedimentary basins that predominantly enriched in primary igneous rocks will provide only small quantities of Br into the sediment supply but rocks enriched in glass-bearing igneous rocks may supply relatively greater amounts of Br (Worden, 2018). Bromine is found in sedimentary basins as dissolved Br-, in solid solution in halite (NaClxBr1−x), or in less common salts resulting from potash-facies evaporites, such as sylvite. Bromine is also associated with organic-rich sediments, especially in marine settings, including organic-rich mudstone and coal. At a concentration of 65 mg/L, Br- is the second most abundant halogen in modern seawater.

    Organic matter and its more evolved forms –kerogen and hydrocarbons– are typical of most large evaporite basins. Mesohaline carbonates interlayered with anhydrite and halite beds can entrain high levels of organic matter to form high-yield source rocks, while the brine inclusions in some halites contain high amounts of volatile hydrocarbons and pyrobitumens. Evaporite beds composed of anhydrite or halite make excellent seals holding back large volumes of hydrocarbons (for literature documentation of these observations see Warren, 2016, Chapters 9 and 10). In combination, saline giants and their heat-responsive lithologies will contain vast volumes of potential volatiles, including halocarbons.

    Ozone (O3) destruction

    When halocarbons enter the stratosphere, they decimate the ozone layer, allowing harmful levels of ultraviolet (UV) radiation to reach the earth's surface (Figures 7a, 9a). Ozone is destroyed by the entry of a number of free radical catalysts into the stratosphere; today the most important catalysts are the hydroxyl radical (OH), nitric oxide radical (NO), chlorine radical (Cl) and the bromine radical (Br). Each radical is characterised by an unpaired electron in its molecular structure and is thus extremely reactive. All of these radicals have both natural and man-made sources; at present, most of the OH and NO in the stratosphere is naturally occurring, but human activity has drastically increased the levels of chlorine and bromine.

    The elements that form radicals in the stratosphere are found in stable organic compounds, especially halocarbons, which reach the stratosphere without being destroyed in the troposphere due to their low reactivity. Once in the stratosphere, the Cl and Br atoms are released from the parent halocarbon by the action of ultraviolet light.


    Ozone (O3) is a highly reactive molecule that quickly reduces to the more stable oxygen (O2) form with the assistance of a catalyst (radical). Cl and Br atoms destroy ozone molecules through a variety of catalytic cycles. The simplest example of such a reaction is when a chlorine atom reacts with an ozone molecule, taking an oxygen atom to form chlorine monoxide (ClO) and leaving behind an oxygen molecule (O2) (Figure 9b). The ClO can then react with another molecule of ozone, once more releasing the chlorine atom as ClO, so far yielding two molecules of oxygen. This ClO reaction can be repeated until the ClO is flushed from the stratosphere (Figure 9b, Fahey, 2007)

    Thus the overall effect of halocarbons entering the stratosphere is a decrease in the amount of ozone. A single chlorine radical can continuously destroy ozone for up to two years (this the time scale for its transport back down into the troposphere; Figure 7a). But there are other stratopheric reactions that remove CLO from this catalytic cycle by forming reservoir species such as hydrogen chloride (HCl) and chlorine nitrate (ClONO).

    Bromine radicals are even more efficient than chlorine at destroying ozone on a per-atom basis, but at present there is much less bromine than chlorine in the atmosphere. Laboratory studies have shown that fluorine and iodine atoms can participate in similar catalytic cycles. However, fluorine atoms react rapidly with water and methane to form strongly bound HF in the Earth's stratosphere, while organic molecules containing iodine react so quickly in the lower atmosphere that they do not reach the stratosphere in significant quantities.

    Halocarbon concentrations below the tropopause are always higher by several orders of magnitude than in the stratosphere, which contains the seasonally and locally variable ozone layer responsible for absorption of incident solar UV radiation (Schwandner, 2002). Penetration of the tropopause allows the ascent of long-lived halocarbons and today occurs primarily as a result of rising tropical air masses in a Hadley cell, rare turnover events, or large Plinian volcanic eruptions.

    Over the two to three years a chlorine or bromine radical can remain in the stratosphere, it reacts with ozone and converts it to oxygen. It has been estimated that a single chlorine atom can react with an average of 100,000 ozone molecules before it is removed from the catalytic cycle (Figure 8b. Other halocarbon-enabled reactions drive ozone destruction (these catalysts are derived from anthropogenic CFCs and other industrial halocarbons). Over the past half-century, our anthropogenic focus on ozone destruction from industrial chemicals has driven the public's understanding into to the much-needed legislated prevention of the entry of additional industrial halocarbons (especially CFCs) into the stratosphere.


    Implications

    However, there are additional deep-time implications for the health of the Earth's biota when natural events of the past drastically increased the amount of halocarbons entering the stratosphere, along with increased levels of sulphurous volatiles and greenhouse gases. We know modern volcanic exhalations containing relatively high levels of chlorine and bromine. But times of intense magmatic/volcanogenic or bolide heating of evaporites in a saline giant will contribute even greater volumes of halocarbons to the stratospheric levels of the atmosphere (Figure 10). If coals and peats are also present (typically not in the saline portion of the basin's sediment fill), then the heating of these additional organic-rich sediments will contribute even more carbon to the vast volumes of the halocarbons created by heating of the evaporites. Heating reactions in the saline giant and associated deposits can also supply elevated levels of the greenhouse gases CO2 and CH4. Explosive volcanism tied to the emplacement of LIPs in the region of a saline giant or the atmosphere-scale disturbance linked to the impact of a large bolide in an area underlain by a saline giant are efficient mechanisms to move large volumes of halocarbons, sulphurous volatiles and greenhouse gasses to the troposphere. The third article in this series will document the specific evaporite geology that contributed to four of the five major Phanerozoic extinction events (Figure 10).

    References

    Alvarez, L. W., W. Alvarez, F. Asaro, and H. V. Michel, 1980, Extraterrestrial Cause for the Cretaceous-Tertiary Extinction: Science, v. 208, p. 1095.

    Beerling, D. J., M. Harfoot, B. Lomax, and J. A. Pyle, 2007, The stability of the stratospheric ozone layer during the end-Permian eruption of the Siberian Traps: Philosophical Transactions of the Royal Society of London, ser. A, Mathematical and Physical Sciences, v. 365, p. 1843 –1866.

    Black, B. A., L. T. Elkins-Tanton, M. C. Rowe, and I. U. Peate, 2012, Magnitude and consequences of volatile release from the Siberian Traps: Earth and Planetary Science Letters, v. 317–318, p. 363-373.

    Black, B. A., E. H. Hauri, L. T. Elkins-Tanton, and S. M. Brown, 2014, Sulfur isotopic evidence for sources of volatiles in Siberian Traps magmas: Earth and Planetary Science Letters, v. 394, p. 58-69.

    Bond, D. P. G., I. Savov, P. B. Wignall, M. M. Joachimski, Y. Sun, S. E. Grasby, B. Beauchamp, and D. P. G. Blomeier, 2015, An abrupt extinction in the Middle Permian (Capitanian) of the Boreal Realm (Spitsbergen) and its link to anoxia and acidification: GSA Bulletin, v. 127, p. 1411-1421.

    Bond, D. P. G., and P. B. Wignall, 2014, Large igneous provinces and mass extinctions: An update, in G. Keller, and A. C. Kerr, eds., Volcanism, Impacts, and Mass Extinctions: Causes and Effects, v. 505, Geological Society of America, p. 0.

    Bowen, N. L., 1940, Progressive Metamorphism of Siliceous Limestone and Dolomite: Journal of Geology, v. 49, p. 225-274.

    Burgess, S. D., S. Bowring, and S.-z. Shen, 2014, High-precision timeline for Earth’s most severe extinction: Proceedings of the National Academy of Sciences, v. 111, p. 3316.

    Burgess, S. D., and S. A. Bowring, 2015, High-precision geochronology confirms voluminous magmatism before, during, and after Earth’s most severe extinction: Science Advances, v. 1, p. e1500470.

    Cao, C.-Q., D.-X. Yuan, H. Zhang, L. Xiang, Y.-C. Zhang, Y. Wang, J. Wang, S.-Z. Shen, L. Mu, Q.-F. Zheng, Y.-S. Wu, X.-D. Wang, J. Ramezani, S. A. Bowring, J. Chen, D. H. Erwin, S. D. Schoepfer, C. M. Henderson, and X.-H. Li, 2018, A sudden end-Permian mass extinction in South China: GSA Bulletin, v. 131, p. 205-223.

    Fahey, D. W., 2007, Twenty questions and answers about the ozone layer, Scientific Assessment of Ozone Depletion: 2006: Geneva, World Meteorological Organization, p. 1-38.

    Ganino, C., and N. T. Arndt, 2009, Climate changes caused by degassing of sediments during the emplacement of large igneous provinces: Geology, v. 37, p. 323-326.

    Gorman, J. A., E. U. Petersen, and E. J. Essene, 1984, Anhydrite equilibria and sulfide zonation in the Fowler massive sulfide body, Balmat, New York: AGU v. 65, p. 293.

    Grasby, S. E., B. Beauchamp, D. P. G. Bond, P. B. Wignall, and H. Sanei, 2016, Mercury anomalies associated with three extinction events (Capitanian Crisis, Latest Permian Extinction and the Smithian/Spathian Extinction) in NW Pangea: Geological Magazine, v. 153, p. 285-297.

    Gulick, S. P. S., P. J. Barton, G. L. Christeson, J. V. Morgan, M. McDonald, K. Mendoza-Cervantes, Z. F. Pearson, A. Surendra, J. Urrutia-Fucugauchi, P. M. Vermeesch, and M. R. Warner, 2008, Importance of pre-impact crustal structure for the asymmetry of the Chicxulub impact crater: Nature Geosci, v. 1, p. 131-135.

    Jablonski, D., G. Chaloner William, H. Lawton John, and M. May Robert, 1994, Extinctions in the fossil record: Philosophical Transactions of the Royal Society of London. Series B: Biological Sciences, v. 344, p. 11-17.

    Jones, M. T., D. A. Jerram, H. H. Svensen, and C. Grove, 2016, The effects of large igneous provinces on the global carbon and sulphur cycles: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 441, p. 4-21.

    Korte, C., M. Ruhl, J. Pálfy, C. V. Ullmann, and S. P. Hesselbo, 2019, Chemostratigraphy Across the Triassic–Jurassic Boundary - Chapter 10, in A. N. Sial, C. Gaucher, M. Ramkumar, and V. P. Ferreira, eds., Chemostratigraphy Across Major Chronological BoundariesChemostratigraphy Across Major Chronological Boundaries, Amreican Geophysical Union, Monograph 240, John Wiley and Sons, Inc, p. 185.

    Lowery, C. M., T. J. Bralower, J. D. Owens, F. J. Rodríguez-Tovar, H. Jones, J. Smit, M. T. Whalen, P. Claeys, K. Farley, S. P. S. Gulick, J. V. Morgan, S. Green, E. Chenot, G. L. Christeson, C. S. Cockell, M. J. L. Coolen, L. Ferrière, C. Gebhardt, K. Goto, D. A. Kring, J. Lofi, R. Ocampo-Torres, L. Perez-Cruz, A. E. Pickersgill, M. H. Poelchau, A. S. P. Rae, C. Rasmussen, M. Rebolledo-Vieyra, U. Riller, H. Sato, S. M. Tikoo, N. Tomioka, J. Urrutia-Fucugauchi, J. Vellekoop, A. Wittmann, L. Xiao, K. E. Yamaguchi, and W. Zylberman, 2018, Rapid recovery of life at ground zero of the end-Cretaceous mass extinction: Nature, v. 558, p. 288-291.

    Marzoli, A., S. Callegaro, J. Dal Corso, J. H. F. L. Davies, M. Chiaradia, N. Youbi, H. Bertrand, L. Reisberg, R. Merle, and F. Jourdan, 2018, The Central Atlantic Magmatic Province (CAMP): A Review, in L. H. Tanner, ed., The Late Triassic World: Earth in a Time of Transition: Cham, Springer International Publishing, p. 91-125.

    Newton, R. C., and C. E. Manning, 2005, Solubility of Anhydrite, CaSO4, in NaCl–H2O Solutions at High Pressures and Temperatures: Applications to Fluid–Rock Interaction: Journal of Petrology, v. 46, p. 701-716.

    Pang, K.-N., N. Arndt, H. Svensen, S. Planke, A. Polozov, S. Polteau, Y. Iizuka, and S.-L. Chung, 2013, A petrologic, geochemical and Sr-Nd isotopic study on contact metamorphism and degassing of Devonian evaporites in the Norilsk aureoles, Siberia: Contributions to Mineralogy and Petrology, v. 165, p. 683-704.

    Percival, L. M. E., M. Ruhl, S. P. Hesselbo, H. C. Jenkyns, T. A. Mather, and J. H. Whiteside, 2017, Mercury evidence for pulsed volcanism during the end-Triassic mass extinction: Proceedings of the National Academy of Sciences, v. 114.

    Raup, D. M., and J. J. Sepkoski, 1982, Mass Extinctions in the Marine Fossil Record: Science, v. 215, p. 1501.

    Renne, P. R., A. L. Deino, F. J. Hilgen, K. F. Kuiper, D. F. Mark, W. S. Mitchell, L. E. Morgan, R. Mundil, and J. Smit, 2013, Time Scales of Critical Events Around the Cretaceous-Paleogene Boundary: Science, v. 339, p. 684.

    Robock, A., 2002, The Climatic Aftermath: Science, v. 295, p. 1242.

    Rohde, R. A., and R. A. Muller, 2005, Cycles in fossil diversity: Nature, v. 434, p. 208-210.

    Schoene, B., U. Schaltegger, J. Guex, A. Bartolini, and T. J. Blackburn, 2010, Correlating the end-Triassic mass extinction and flood basalt volcanism at the 100 ka level: Geology, v. 38, p. 387-390.

    Schwandner, F. M., 2002, The Organic Chemistry of Volcanic Gases at Vulcano (Aeolian Islands, Italy): Doctoral thesis, Swiss Federal Institute of Technology Zurich (ETH), 144 p.

    Sepkoski, J. A., 2002, Compendium of Fossil Marine Animal Genera, in D. Jablonski, and M. Foote, eds., Bull. Am. Paleontol. no. 363 (Paleontological Research Institution, Ithaca, 2002).

    Sepkoski Jr., J. J., 1996, Patterns of Phanerozoic extinction: a perspective from global data bases, in O. H. Walliser, ed., Global Events and Event Stratigraphy.: Berlin, Springer-Verlag, p. 35-52.

    Sigurdsson, H., 1990, Evidence of volcanic loading of the atmosphere and climate response: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 89, p. 277-289.

    Tanner, L. H., 2018, Climates of the Late Triassic: Perspectives, Proxies and Problems, in L. H. Tanner, ed., The Late Triassic World: Earth in a Time of Transition: Cham, Springer International Publishing, p. 59-90.

    Textor, C., H.-F. Graf, M. Herzog, and J. M. Oberhuber, 2003, Injection of gases into the stratosphere by explosive volcanic eruptions: Journal of Geophysical Research: Atmospheres, v. 108.

    Visscher, H., C. V. Looy, M. E. Collinson, H. Brinkhuis, J. H. A. van Konijnenburg-van Cittert, W. M. Kürschner, and M. A. Sephton, 2004, Environmental mutagenesis during the end-Permian ecological crisis: Proceedings of the National Academy of Sciences of the United States of America, v. 101, p. 12952.

    von Frese, R. R. B., L. V. Potts, S. B. Wells, T. E. Leftwich, H. R. Kim, J. W. Kim, A. V. Golynsky, O. Hernandez, and L. R. Gaya-Piqué, 2009, GRACE gravity evidence for an impact basin in Wilkes Land, Antarctica: Geochemistry, Geophysics, Geosystems (DOI 10.1029/2008GC002149), v. 10.

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

    West, R. R., and W. J. Sutton, 1954, Thermography of Gypsum: Journal of the American Ceramic Society, v. 37, p. 221-224.

    Wignall, P. B., 2001, Large igneous provinces and mass extinctions: Earth-Science Reviews, v. 53, p. 1-33.

    Worden, R. H., 2018, Halogen Elements in Sedimentary Systems and Their Evolution During Diagenesis, in D. E. Harlov, and L. Aranovich, eds., The Role of Halogens in Terrestrial and Extraterrestrial Geochemical Processes: Surface, Crust, and Mantle: Cham, Springer International Publishing, p. 185-260.


     

    Comments
    Post has no comments.
    Post a Comment




    Captcha Image

    Trackback Link
    http://www.saltworkconsultants.com/BlogRetrieve.aspx?BlogID=13454&PostID=765911&A=Trackback
    Trackbacks
    Post has no trackbacks.

    Recent Posts


    Tags

    hydrogen Realmonte potash venice alkaline lake hydrothermal karst lot's wife hydrohalite well logs in evaporites lazurite Messinian wireline log interpretation 18O solikamsk 2 water in modern-day Mars halophile Clayton Valley playa: 13C sodium silicate Ganymede Hell Kettle epsomite methanotrophic symbionts trona non solar heating salts salt suture Deep Europe evaporite End-Cretaceous methane salt periphery Dead Sea caves source rock phreatic evaporite Calyptogena ponderosa Hadley cell: Warrawoona Group hydrothermal anhydrite Neoproterozoic Oxygenation Event sulfur CaCl2 brine stevensite well log interpretation Mixing zone Muriate of potash vadose zone SOP Precambrian evaporites palygorskite authigenic silica potash knistersalz evaporite dissolution causes of glaciation chert Great Salt Lake Lop Nor mummifiction Hadley Cell hydrothermal halite MOP bedded potash Pilbara Density log Gamma log Ure Terrace seal capacity sepiolite High Magadi beds evaporite karst basinwide evaporite gas in salt K2O from Gamma Log oil gusher waste storage in salt cavity End-Permian sinjarite evaporite-metal association auto-suture Noril'sk Nickel namakier bischofite saline giant RHOB Mega-monsoon Koeppen Climate sinkhole lithium brine black salt Proterozoic carbon cycle Prograde salt York (Whitehall) Mine Beebe hydrothermal field extrasalt supercritical halite halite-hosted cave Large Igneous Magmatic Province perchlorate subsidence basin cauliflower chert Platform evaporite Sumo Ceres Mesoproterozoic mass die-back Prairie Evaporite flowing salt Schoenite geohazard eolian transport carbon oxygen isotope cross plots sulfate Mars Boulby Mine Mulhouse Basin Lake Magadi sulphate anthropogenically enhanced salt dissolution intersalt doline salting-out climate control on salt phreatomagmatic explosion nacholite 13C enrichment Belle Isle salt mine deep seafloor hypersaline anoxic basin jadarite MgSO4 depleted Neoproterozoic stable isotope salt ablation breccia white smokers kainitite Koppen climate potash ore price Stolz diapir gassy salt endosymbiosis lithium carbonate gem meta-evaporite gas outburst dissolution collapse doline saline clay salt karst Kara bogaz gol zeolite Catalayud recurring slope lines (RSL) base metal dark salt lithium battery lunette causes of major extinction events Seepiophila jonesi Dead Sea saltworks brine lake edge hydrological indicator Turkmenistan gypsum dune hectorite Lake Peigneur DHAL African rift valley lakes MVT deposit GR log organic matter Jefferson Island salt mine brine evolution Paleoproterozoic Oxygenation Event NPHI log Pangaea deep meteoric potash CO2: albedo Ripon intrasalt North Pole Quaternary climate ancient climate Phaneozoic climate salt mine potash ore Hyperarid Lomagundi Event mine stability Bathymodiolus childressi carnallitite crocodile skin chert vanished evaporite natural geohazard halite Kalush Potash Belle Plain Member NaSO4 salts freefight lake Salar de Atacama mirabilite Thiotrphic symbionts Zaragoza Red Sea sedimentary copper End-Triassic Badenian Enceladus lapis lazuli silicified anhydrite nodules Sulphate of potash nuclear waste storage SedEx Evaporite-source rock association collapse doline Dallol saltpan LIP seawater evolution rockburst halotolerant marine brine Musley potash Zabuye Lake nitrogen ozone depletion salt leakage, dihedral angle, halite, halokinesis, salt flow, MgSO4 enriched Lamellibrachia luymesi Lop Nur supercontinent Magdalen's Road Weeks Island salt mine Ethiopia supercritical phase retrograde salt anthropogenic potash allo-suture Danakhil Depression, Afar antarcticite cryogenic salt Stebnyk potash halokinetic Stebnik Potash capillary zone HYC Pb-Zn water on Mars cryogenic spring salts McMurdo Sound phreatic explosion Archean anomalous salt zones blowout salt tectonics H2S halocarbon well blowout tachyhydrite halogenated hydrocarbon Deep seafloor hypersaline anoxic lake Corocoro copper snake-skin chert Ingebright Lake Five Island salt dome trend solar concentrator pans sulphur evaporite-hydrocarbon association vestimentiferan siboglinids methanogenesis salt trade extraterrestrial salt Neutron Log Atlantis II Deep Crescent potash brine pan Dead Sea karst collapse Karabogazgol McArthur River Pb-Zn DHAB astrakanite salt seal 18O enrichment hydrothermal potash circum-Atlantic Salt Basins silica solubility CO2 Patience Lake member magadiite SO2 dihedral angle

    Archive